Close
About
FAQ
Home
Collections
Login
USC Login
Register
0
Selected
Invert selection
Deselect all
Deselect all
Click here to refresh results
Click here to refresh results
USC
/
Digital Library
/
University of Southern California Dissertations and Theses
/
Spatial and temporal evolution of magmatic systems in continental arcs: a case study of dynamic arc behaviors in the Mesozoic Sierra Nevada, California
(USC Thesis Other)
Spatial and temporal evolution of magmatic systems in continental arcs: a case study of dynamic arc behaviors in the Mesozoic Sierra Nevada, California
PDF
Download
Share
Open document
Flip pages
Contact Us
Contact Us
Copy asset link
Request this asset
Transcript (if available)
Content
Spatial and temporal evolution of magmatic systems in continental arcs:
a case study of dynamic arc behaviors in the Mesozoic Sierra Nevada,
California
by
Katie E. Ardill
A Dissertation Presented to the
FACULTY OF THE USC GRADUATE SCHOOL
UNIVERSITY OF SOUTHERN CALIFORNIA
In Partial Fulfillment of the
Requirements for the Degree
DOCTOR OF PHILOSOPHY
(GEOLOGICAL SCIENCES)
August 2020
Copyright 2020 Katie Ardill
ii
Acknowledgements
The research in this dissertation was funded by a National Science Foundation grant EAR-1624847 to Scott
Paterson and EAR-1624854 to Vali Memeti. In addition, funding from the Geological Society of America
Student Research grant program (2015, 2017), Lipman Research Awards, and USC Department of Earth
Sciences Graduate Student Research Fund to Katie Ardill supported fieldwork and lab analyses.
I would firstly like to thank Scott, my Ph.D. advisor, who has helped me grow into a scientist these past 7
years, ever since I first visited Yosemite, and has always encouraged me to see the big picture in all things
but especially when field mapping and thinking about arc processes. His generous advice and feedback has
been invaluable as I have navigated the joys and challenges of juggling multiple research projects, writing
papers and grants, doing remote fieldwork, mentoring students, and building a career in academia. I would
like to thank him for all the opportunities he has provided for me at USC. I thank Vali Memeti, who has
supported me throughout this entire journey. Vali has spent a lot of time helping me improve my research
skills and look at things in new ways, as well as teaching me the ins and outs of labwork and fieldwork. I
am very grateful to have learned all of this from her. I still can’t believe how many zircons we separated in
the space of a month!
My committee members John Platt, Mike Inkpen, and Cal Barnes have also been very supportive these past
few months of the dissertation, especially during this global pandemic. I thank them for their feedback and
encouragement. I would like to thank John for taking our microstructure class to Monarch Canyon and
showing us (including my mum) all of the constellations. I thank Cal for his mentorship and guidance over
the entire course of my Ph.D., and from whom I have learned so much about minerals, magmas, and
geochemistry. I have immensely enjoyed my times in the lab with Cal and Melanie in Lubbock and look
forward to going back soon!
My Ph.D. experience would not have been the same without the huge network of colleagues and friends
who I am fortunate to work with; Ana Martinez-Ardila, Pablo Alasino, Kevin Werts, Melanie Barnes,
Rosario Esposito, Dustin Williams, Melissa Chambers, and Louis Oppenheim. I want to especially thank
the grads in my research group: Wenrong Cao, Sean Hartman, Babsi Ratschbacher, Snir Attia, Abby
Wesley, and Cullen Scheland, and the Platt group: Alex Lusk, Will Schmidt, Tarryn Cawood, and Naomi
Rodgers, for their friendship and support through the ups and downs of grad student life. The Earth Sciences
department has been a great place to work, and I thank all the undergraduates, many of whom who came
out to the field with me, the grad students, and the staff, especially Cindy Waite who has always made sure
I kept making progress towards my degree! I would also like to thank Doug Hammond for his guidance in
iii
teaching as I TA’d his Mineralogy class, as well as Jonny Stanback, Kelly O’Rourke, Ariel Borsook, Leslie
Insixiengmay and many students from Durham University for their assistance during fieldwork and their
work on senior thesis projects on the Sierra Nevada arc.
Finally, I reserve a special thanks for my family. For my mum and dad who have worked so hard and made
sacrifices for me to get to this point today, and who always encouraged me to do my best. To my brother,
who spent an entire summer in the Sierras with me and somehow didn’t become a geologist at the end.
Most of all, I would like to thank Sean, whose unwavering love and support has made this journey an
unforgettable one. I can’t wait to see what is in store for our next chapter!
iv
Table of Contents
Acknowledgements ..................................................................................................................................... ii
List of Tables.............................................................................................................................................. vi
List of Figures ........................................................................................................................................... vii
Abstract ....................................................................................................................................................... x
Chapter 1: Introduction .............................................................................................................................. 1
Chapter 2: Spatiotemporal magmatic focusing in upper-mid crustal plutons of the Sierra Nevada arc .... 9
Abstract ............................................................................................................................. 9
Introduction ..................................................................................................................... 10
Background ..................................................................................................................... 11
Arc-wide migration patterns ........................................................................................... 14
Spatiotemporal patterns in the central Sierra Nevada ..................................................... 17
Discussion ....................................................................................................................... 24
Proposed models for magmatic focusing ........................................................................ 27
Magmatic focusing at all crustal levels? ......................................................................... 29
Conclusions ..................................................................................................................... 29
Acknowledgements ......................................................................................................... 31
Chapter 3: Spatially and temporally dynamic continental arc activity in the Mesozoic central Sierra
Nevada: Implications for arc lithosphere evolution................................................................32
Abstract ........................................................................................................................... 32
Introduction ..................................................................................................................... 33
Background ..................................................................................................................... 34
Methods ........................................................................................................................... 43
Results ............................................................................................................................. 45
Discussion ....................................................................................................................... 65
Conclusions ..................................................................................................................... 77
Acknowledgements ......................................................................................................... 77
Chapter 4: Reconstructing the physical and chemical development of a pluton-porphyry complex in a
.....................tectonically re-organized arc crustal section, Tioga Pass, Sierra Nevada ............................ 78
v
Abstract ........................................................................................................................... 78
Introduction ..................................................................................................................... 79
Previous work ................................................................................................................. 80
Methods ........................................................................................................................... 84
Results ............................................................................................................................. 86
Discussion ..................................................................................................................... 105
Conclusions ................................................................................................................... 115
Acknowledgements ....................................................................................................... 115
Chapter 5: Schlieren-bound magmatic structures record crystal flow-sorting in dynamic upper-crustal
....................magma-mush chambers ....................................................................................................... 117
Abstract ......................................................................................................................... 117
Introduction ................................................................................................................... 118
Background ................................................................................................................... 120
Methods ......................................................................................................................... 129
Results ........................................................................................................................... 138
Discussion ..................................................................................................................... 155
Conclusions ................................................................................................................... 171
Acknowledgements ....................................................................................................... 172
Chapter 6: Conclusions .......................................................................................................................... 173
References ............................................................................................................................................... 181
Appendices .............................................................................................................................................. 212
Appendix A: Co-authored manuscripts ........................................................................ 212
Appendix B: Supplementary data to Chapter 2 ............................................................ 267
Appendix C: Supplementary data to Chapter 3 ............................................................ 283
Appendix D: Supplementary data to Chapter 4 ........................................................... 311
Appendix E: Supplementary data to Chapter 5 ............................................................ 330
vi
List of Tables
Table 4.1 Summary of LA-ICP-MS U-Pb zircon data for Tioga Pass samples ........................................ 88
Table 4.2 Whole rock major oxide, trace element and isotope compositions of the Tioga Pass intrusive
....................complex ............................................................................................................................... 103
Table 5.1 Modal mineralogy of representative schlieren and associated components ............................ 136
Table 5.2 Representative whole-rock compositions of schlieren and associated components ............... 137
Table 5.3 Magma viscosity summary for bulk-rock and calculated melt compositions ......................... 156
vii
List of Figures
Figure 1.1 Schematic diagram summarizing multi-scale dynamic processes ............................................. 3
Figure 2.1 Summary of magmatic focusing in volcanic fields and the central Sierra Nevada (CSN) ...... 12
Figure 2.2 Graph of pluton age vs. distance across the CSN .................................................................... 15
Figure 2.3 Map of spatiotemporal magmatic focusing in the CSN ........................................................... 16
Figure 2.4 Whole rock trace element and isotopic ratios vs. age of CSN samples ................................... 18
Figure 2.5 Graph of calculated volumes (km
3
) vs. pluton age (Ma) in the CSN focusing region ............ 20
Figure 2.6 Graph of SiO 2 variation with time within CSN plutons .......................................................... 22
Figure 2.7 Schematic diagram summarizing transcrustal and emplacement level models for magma
...................focusing .................................................................................................................................. 30
Figure 3.1 Cartoon summarizing the Paleozoic-Mesozoic evolution of the SW Cordilleran margin ....... 35
Figure 3.2 Geologic map of the CSN ........................................................................................................ 37
Figure 3.3 Geologic map of the broader central Sierra Nevada region and sample locations .................. 40
Figure 3.4 Graphs summarizing sample geochemical attributes within the compilation .......................... 44
Figure 3.5 Boxplots showing sample geochemistry divided by SiO 2 and crustal level ............................ 47
Figure 3.6 Plots showing trace element ratios vs. west-east position ....................................................... 50
Figure 3.7 Plots showing Dy and Eu anomalies vs. west-east position .................................................... 51
Figure 3.8 Plots showing Ba/La and Th/La ratios vs. west-east position ................................................. 52
Figure 3.9 Plots showing isotope ratios vs. west-east position ................................................................. 53
Figure 3.10 Plots showing trace element ratios vs. sample age ................................................................ 55
Figure 3.11 Plots showing Dy and Eu anomalies vs. sample age ............................................................. 56
Figure 3.12 Plots showing Ba/La and Th/La ratios vs. sample age .......................................................... 57
Figure 3.13 Plots showing isotope ratios vs. sample age .......................................................................... 58
Figure 3.14 Compiled K/Ar and
40
Ar/
39
Ar biotite and amphibole ages vs. west-east position ................. 62
viii
Figure 3.15 Plots comparing isotopic systems to major magma reservoirs .............................................. 64
Figure 3.16 Schematic diagram of the Cretaceous CSN dynamic crustal section .................................... 72
Figure 3.17 Block diagram comparing the Triassic-Jurassic arcs to the Cretaceous arc .......................... 75
Figure 4.1 Simplified geologic map of the CSN, illlustrating plutonic, hypabyssal, and volcanic fields . 81
Figure 4.2 Geologic map of the Tioga Pass area (1:10,000 scale) ............................................................ 82
Figure 4.3 Cross-sections of the Tioga Pass area ...................................................................................... 87
Figure 4.4 Schematic stratigraphic columns of units in the Tioga Pass area ............................................ 89
Figure 4.5 Normalized age probability plots for Paleozoic-Jurassic sediments and Mesozoic volcanic
....................rocks ...................................................................................................................................... 90
Figure 4.6 Cross section between the porphyry and quartz monzodiorite units ...................................... 93
Figure 4.7 Field photos and photomicrographs of Tioga Pass samples .................................................... 94
Figure 4.8 U-Pb zircon age distributions for Cretaceous intrusive units at Tioga Pass ............................ 98
Figure 4.9 Elemental geochemistry data plots for Tioga Pass samples .................................................. 101
Figure 4.10 Isotopic data from Tioga Pass with regional constraints ..................................................... 106
Figure 4.11 Cartoon of the Tioga Pass magmatic complex at 100 Ma ................................................... 107
Figure 4.12 Cartoon summarizing different plutonic-hypabyssal-volcanic models ............................... 110
Figure 5.1 Diagram summarizing formation mechanisms for schlieren and magmatic structures ......... 122
Figure 5.2 Map synthesis of magmatic structures in the Tuolumne Intrusive Complex (TIC) ............... 124
Figure 5.3 Field photos of planar schlieren and trough structures .......................................................... 132
Figure 5.4 Field photos of magmatic tubes ............................................................................................. 133
Figure 5.5 Field photos of plumes ........................................................................................................... 134
Figure 5.6 Map summarizing field data from the Glen Aulin area (NW quadrant) ................................ 140
Figure 5.7 Map summarizing field data from the Young Lakes area (NE quadrant) .............................. 143
Figure 5.8 Maps summarizing field data from the Tenaya Peak and Lyell Canyon areas (SE and SW
....................quadrants) ............................................................................................................................ 144
ix
Figure 5.9 Grid maps showing complex field relationships within schlieren-bound structures ............. 148
Figure 5.10 Major and trace element data from schlieren and associated components .......................... 152
Figure 5.11 Rare Earth Element patterns and isotope ratios in schlieren and associated components ... 153
Figure 5.12 Cartoon summary of multi-scale dynamic processes operating in mobile magma mushes . 159
x
Abstract
Continental arc systems, situated above subduction zones, are sites of continental crust formation and
recycling, volatile and sediment cycling, natural hazards, and economic ore deposits. Although subduction
is thought to be a continuous process, the behavior of arcs is known to be non-steady-state from magmatic,
deformation, and climate perspectives. This non-steady-state nature of arcs leads to uncertainty surrounding
the evolution of magma plumbing systems from the mantle to the surface.
This study incorporated field, structural, geochronologic, and geochemical observations of arc behavior
across several spatial and temporal scales to investigate the processes driving non-steady-state arc
magmatism within the Sierra Nevada continental arc section and evaluate the effects of these processes on
the arc lithosphere. At the arc-section scale, this involved characterizing regional-scale, inward magma
focusing of calderas and upper-crustal plutons within the context of an arc system that was also migrating,
dramatically thickening, and experiencing a magmatic flare-up. These dynamic processes, with distinct
spatial and temporal signals, led to increased magma chamber sizes and magma hybridization in the upper-
crust. Shallowly emplaced hypabyssal intrusions provide further information on the physical and chemical
relationship between volcanic and upper-crustal plutonic domains. They bear evidence for crystal
accumulation, magma ascent, and rapid cooling in magma feeder zones. Thus, they are useful to reconstruct
volcanic fields in ancient, exhumed arc sections like the Sierra Nevada. Within intrusive complexes,
compositionally-defined magmatic structures were found to represent records of spatially and temporally
heterogeneous magmatic flow in crystal-rich mushes and record the rheological evolution of dynamic
magma mushes from millimeter to kilometer scales.
Results of this research demonstrate that across wide spatial and temporal scales, during magma
chamber growth, ascent and eruption, the Sierran arc system was highly dynamic. The Late Cretaceous arc,
in particular, was thermally and mechanically primed to enable transcrustal magmatic differentiation.
Isotopic geochemical tracers in particular suggest that the mantle plays a critical role in controlling these
behaviors. The implications of this study are that dynamic, non-steady-state behavior may be the most
efficient way to generate mountains and continental crust by transforming the physical, thermal, and
chemical properties of the lithosphere over a relatively short duration (10’s myr) in transcrustal magmatic
plumbing systems.
1
Chapter 1: Introduction
Continental arc systems are among the most dynamic geologic settings on Earth, continually evolving
in space and time. Volcanic structures are the source of several natural hazards that affect human life and
property, release gases that can modify long-term climate, and are sites of economic ore deposits. However,
the volcanic products of arc magmatism are volumetrically minor compared to deeper, intrusive regions of
the arc, where 20-30 times more magma is trapped beneath the surface than is ever erupted (Paterson and
Ducea, 2015; Ward et al., 2017). Thus, the arc overall represents a significant transfer of heat and mass
from the Earth’s interior towards the surface. In addition, subduction zones are the primary locations of
continental crust formation and recycling (e.g., Stern and Scholl, 2010), as well as a crucial mechanism for
volatile cycling that supports life on our planet (Wallace et al., 2015; Zellmer et al., 2015). Continental arc
magmatism often (but not always) results in the formation of high-elevation mountain ranges, which
supplies and distributes sediments across the margin, and influences global weathering and climate (Lee et
al., 2015).
Although subduction of the oceanic slab beneath the continent is thought to be a continuous process,
many aspects of continental arc systems are known to be non-steady-state. This behavior is summarized by
the term arc tempos (Paterson and Ducea, 2015). Early studies recognized pulses, or episodes, of high-
volume magmatism (flare-ups) separated by periods of lower-volume magmatism (lulls) over 10’s of
millions of years (Armstrong, 1988; Armstrong and Ward, 1993; Bateman, 1992). This pattern has since
been identified widely in continental arc settings across geologic time and space using U-Pb in zircon ages
and geologic map area and volume calculations (Ducea, 2001; Gehrels et al., 2009; Ji et al., 2009; Paterson
and Ducea, 2015; Kirsch et al., 2016; Schwartz et al., 2017; Cao et al., 2017; Rapela et al., 2018; Martinez-
Ardila et al., 2019). Studies also established the occurrence of episodic deformation (DeCelles et al., 2009;
Cao et al., 2015), and the impacts of episodic arc activity on long-term climate (e.g., Lee and Lackey, 2015;
Ratschbacher et al., 2019). It remains unclear if flare-ups and lulls are a unique feature of continental arcs;
continental arc flare-up magma addition rates are comparable to oceanic arc magma addition rates,
suggesting that flare-ups might represent ‘regular’ arc conditions, and that lull periods are unusual (Jicha
and Jagoutz, 2015; Ratschbacher et al., 2019). However, lull periods, by their definition as low-volume
events, remain challenging to study.
Several driving mechanisms for arc flare-ups have been proposed, ranging from internal mechanisms
such as increased crustal input (Ducea and Barton, 2007; DeCelles et al., 2009) or increased mantle input
(Alasino et al., 2016; Martinez-Ardila et al., 2019; Attia et al., 2020), or driven by external processes or
2
changes in the subducting plate, such as a slab tear (Hughes and Mahood, 2008; Decker et al., 2017). In
general, upper-plate processes have been considered dominant in Cordilleran arc systems (e.g., Kirsch et
al., 2016). How flare-ups and lulls are generated has several implications for arc systems, including where
and how compositional diversity is generated in an arc (compare Annen et al., 2006 to Marsh, 2004), the
thermal budgets and thermal profiles of arcs (Barton and Hanson, 1989; de Silva and Gosnold, 2007;
Karakas et al., 2017), when and how dense lithospheric roots form (Saleeby et al., 2003), as well as arc-
wide mass balance calculations, including estimates of new continental crust formed in arcs (e.g., Paterson
et al., 2014; Ducea et al., 2015). If flare-ups are largely mantle-driven, they represent punctuated periods
of voluminous new crust formation, however in a crustal-driven flare-up, recycling of crustal material is
dominant.
Additional ways in which arcs are dynamic include: (i) arc migration, where the magmatic front sweeps
inboard or outboard of the continental edge (Coney and Reynolds, 1977; Kay et al., 2005; Karlstrom et al.,
2014), (ii) magma focusing, where the locus of magmatism migrates towards a central point (Smith, 1979;
Hildreth, 1981; Grunder et al., 2008; de Silva and Gosnold, 2007; Ardill et al., 2018), and (iii) crustal
thickening and thinning, driven by both magmatic and tectonic processes (Profeta et al., 2015; Cao et al.,
2016) (Fig. 1.1). As with flare-ups, the causes of each of these arc behaviors are debated and interpretations
vary widely (discussed in Chapters 2 and 3). In summary, contrasting models for each of the arc behaviors
listed above places different constraints on how magma plumbing systems evolve physically and
chemically, from the mantle to the surface. Between proposed models, there is little overlap; this influences
our interpretations of how these processes may interact with each other, and whether feedbacks can exist
between them.
Dynamic, episodic behavior is certainly not limited to the arc-wide, or subduction-zone scales. Within
the upper-crust (<10-15 km depth), incremental magma emplacement and volcanic eruptions represent
examples of dynamic behavior that in some cases is also non-steady-state (e.g., Miller and Paterson, 2001;
Grunder et al., 2008; Memeti et al., 2010; Coint et al., 2013a,b; Lipman and Bachmann, 2015; Paterson et
al. 2016). Much of the debate surrounding the dynamics of upper-crustal magma plumbing systems
concerns the physical and chemical connection between plutonic and volcanic rocks, and how magma
chambers evolve (specifically the size, longevity, and mobility of magma chambers), which are discussed
below:
Since seminal studies by Hutton (1788) and Lyell (1838), volcanic and plutonic rocks have been
considered spatially and temporally associated. Subsequent research has increasingly focused on
developing a framework to genetically relate volcanic and plutonic rocks by magmatic processes (e.g.,
Lipman, 1984; Bachmann et al., 2007; Glazner et al., 2015). End-member models propose plutons are either
3
Figure 1.1 Schematic diagram summarizing the dynamic processes explored in this study. The central Sierra
Nevada represents an upper-crustal arc section up to ~10 km depth. Some arc-scale features are shown to provide a
broader context for CSN magmatism. Boxes are used to show different scales of dynamic processes explored in
subsequent chapters. (A) Transcrustal dynamic arc processes expressed in the Cretaceous CSN. (B) Summary of
different volcanic-hypabyssal-plutonic connections. (3) Magma chamber evolution recorded in compositionally-
defined magmatic structures. References: [1] Miller et al. (2009); [2] Cao et al. (2016); [3] Walker et al. (2015); [4]
Brown (2007).
4
“crystal graveyards”, cumulate rocks that have lost melt to feed volcanic eruptions (e.g. Bachmann and
Bergantz, 2004; Gelman et al., 2014; Lee and Morton, 2015; Deering et al., 2016; Barnes et al., 2020), or
that plutons are “failed eruptions” and share no physical or chemical history with volcanic deposits (Glazner
et al., 2015). In the shallow crust, where the cumulate signature is subtle, geochemical tracers and mineral
studies have been critical in identifying and estimating the amount of crystal accumulation in plutonic rocks
to test these hypotheses (Deering and Bachmann, 2010; Schaen et al., 2017; Barnes et al., 2016, 2020;
Werts et al., 2020). Additional challenges in understanding the volcanic-plutonic relationship include the
limited exposures of multiple crustal levels, as well as the composite, overprinted records of both volcanic-
plutonic fields and deformation histories in both modern and ancient continental arcs (e.g., Klemetti et al.,
2014; Watts et al., 2016). Hypabyssal intrusive complexes are important features in the upper-crust as they
are exposed in both volcanic- and plutonic-dominated environments (e.g., Kistler and Swanson, 1981;
Colgan et al., 2018) and thus can provide a ‘snapshot’ of magmatic processes such as crystallization, ascent
and eruption, discussed in Chapter 4 (see also Zimmerer and McIntosh, 2013).
To scale down even further, within incrementally-assembled intrusive complexes, major advances have
been made on structural, geochronologic, geochemical, theoretical and computational fronts to reconstruct
magmatic histories (e.g., reviews by Miller et al., 2007; Paterson et al., 2018; Bachmann and Huber, 2016;
Sparks et al., 2019). The Tuolumne Intrusive Complex (TIC) represents one well-studied example where a
range of analytical tools have been applied, and debate on the emplacement and subsequent magmatic
history of the complex continues (e.g., Memeti et al., 2014; Paterson et al., 2016; Bartley et al., 2018).
Studies of the TIC have (i) documented protracted crystallization histories spanning 10
5
-10
7
yr. (e.g.,
Coleman et al., 2004; Memeti et al., 2010, 2014); (ii) defined various types of internal contact relationships
(Bateman, 1992; Memeti et al., 2010, 2014; Paterson et al., 2016); (iii) identified fractionation and mixing
processes, crystal accumulation, and crystallization conditions using bulk rock and mineral geochemistry
(Bateman and Chappell, 1979; Kistler et al., 1986; Bateman, 1992; Gray et al., 2008; Memeti et al., 2014;
Barnes et al., 2016; Werts et al., 2020); (iv) modeled thermal histories (Paterson et al., 2011), (v) interpreted
strain histories using magmatic fabrics (Žák et al., 2007; Paterson and Ardill, 2019), (vi) provided evidence
of magmatic erosion and recycling (Paterson et al., 2016); (vii) and examined the formation of
compositionally defined magmatic structures (e.g., Reid, 1993; Solgadi and Sawyer, 2008; Paterson, 2009;
Hodge et al., 2012; Bartley et al., 2018). The resulting interpretations of these datasets vary. On one hand,
the TIC has been proposed to consist of sheets with ephemeral mush systems that have little interaction
with each other during or after emplacement, reflecting a relatively static upper-crustal history (e.g.,
Coleman et al., 2012; Bartley et al., 2018). In contrast, other studies have proposed that the TIC formed
several long-lived (i.e., 0.5 – 2 m.y.) magma chambers, regions of interconnected magma mush that
experienced dynamic convection and return flow (e.g., Bateman, 1992; Burgess and Miller, 2008; Solgadi
5
and Sawyer, 2008; Memeti et al., 2010; Paterson et al., 2011, 2016). Each of these models has distinct
implications for all of the other arc processes discussed so far, including where in the crust compositional
diversity is attained, the thermal structure of the crust and the response of the surrounding host rock, the
volumes of magma that are available to erupt, and the rheology of the magma.
This study aims to investigate the patterns of dynamic arc activity across a range of scales to better
understand their causes and consider their contributions to the overall arc framework. From arc-wide to
pluton-scale these patterns include: (1) the lithospheric-scale record of magmatism during dynamic arc
activity, and the evolution of magma sources; (2) the links at the scale of the magma plumbing system
between upper-crustal magma storage and eruption; (3) the magma-chamber record of the hypersolidus
histories recorded in compositionally defined magmatic structures. Each study offers a different perspective
on arc systems, and explores different dynamic, non-steady-state aspects of arc behavior.
The Mesozoic Sierra Nevada arc section is the focus of this research; it is a classic example of a
continental arc, built across a transitional oceanic to continental margin (e.g., Kistler and Peterman, 1973).
The Sierra Nevada arc represents a natural laboratory for investigating physical and chemical aspects of arc
magmatism, owing to a rich body of existing research with which to investigate age, volume, and
geochemical patterns from the arc-scale to the intrusive-complex scale (and beyond to the mineral-scale!).
The approach taken here combines a compilation of previous work with new field mapping, structural data,
geochronology, and geochemistry. This study leverages the previous work that several dynamic processes
occurred in the Sierra Nevada arc during its lifespan, reducing the effects of variable tectonic setting,
basement type, and crustal thickness that can factor into interpretations when comparing across multiple
arcs.
To explore these research topics, the following questions were addressed:
(1) Transcrustal, arc-scale processes: What are the spatial and temporal signals of dynamic processes (e.g.,
flare-ups, migration, focusing and crustal thickening) operating in the Sierra Nevada? At what rates do these
processes occur? How do dynamic arc behaviors influence the evolution of the lithosphere? Do they
effectively ‘prime’ arcs to form large, upper-crustal magma chambers? How is magma focusing manifest
in the upper crust, from plutonic fields (6-10 km depth) to volcanic fields?
(2) Upper-crustal processes: What are the physical and chemical connections between upper-crustal
magma storage regions and volcanic rocks? In what ways do ancient, exhumed arc sections (where volcanic
exposures are limited) offer a new perspective? Which aspects of magma ascent, differentiation, storage,
and eruption are preserved in hypabyssal intrusions?
6
(3) Internal, magma chamber processes: Which processes are dominant in creating compositional and
structural diversity in magma storage regions? What were the magmatic conditions and magma rheology
during the formation of compositionally-defined magmatic structures? How can we constrain the size,
longevity, and mobility of magma mushes in now-exhumed plutons? What are the implications of structural
diversity in plutons for their emplacement history?
These research questions are discussed in more detail in the following chapters:
Chapter 2: In this study, Cretaceous magmatic activity in the central Sierra Nevada (CSN) was investigated
from the perspective of spatial and temporal evolution, using a compilation of U-Pb zircon ages, bulk-rock
element and isotope compositions, and area and volume measurements. Late Cretaceous spatial and
temporal patterns are distinct from Early Cretaceous arc magmatism, and characteristic of regional magma
focusing. This phenomenon resulted in compositional hybridization across a ~4,000 km
2
field of >102-85
Ma plutons, with the nested 1,100 km
2
Tuolumne Intrusive Complex (TIC) formed at the center of this
focusing. Additionally, several pendants in this area preserve volcanic rocks with evidence for explosive
eruptions of the same age. Volumetrically minor hypabyssal intrusions (emplaced between 0-6 km depth)
match plutonic and volcanic trends, strengthening the volcanic-plutonic connection, and establishing
magma focusing as a multi-level, possibly transcrustal, dynamic process. The CSN focusing center is the
first view of the deeper, plutonic regions of a magma focusing center; a comparison with known volcanic
focusing examples all demonstrate that focusing result in the generation of large silicic batholiths (observed
geophysically or inferred) and ignimbrite eruptions during arc-wide flare-up periods. I compiled data from
the literature and collected new data, including new U-Pb in zircon ages with V.M. I wrote the first draft of
the paper, and drafted figures; all authors (K.A, S.P, and V.M) were involved in editing of the manuscript
for publication. This chapter is published in Earth and Planetary Science Letters (doi:
10.1016/j.epsl.2018.06.023).
Chapter 3: Large structural, geochronologic, and geochemical datasets from the CSN demonstrate that the
Mesozoic Sierra Nevada arc experienced three magmatic-tectonic flare-ups, with the most voluminous
flare-up occurring in the Cretaceous. During the Cretaceous flare-up the arc also migrated eastward and
locally magmatically focused, while the crust dramatically thickened, culminating in a ~70 km thick crustal
column. Magmatic flare-ups have distinct bulk-rock isotopic signals through time, in contrast to previous
findings using entire-arc datasets (e.g., DeCelles et al., 2009; Paterson and Ducea, 2015). During the
Cretaceous, some of the isotopic variation can be attributed to eastward arc migration, as magmas traversed
through and incorporated different crust and mantle material, a signature which is also seen in trace-element
ratios that suggest an increasing role for amphibole and feldspar fractionation with eastward position (and
time). The restricted range of Sr and Nd whole-rock isotopes in the Late Cretaceous focusing zone, along
7
with field evidence, suggests that magmas that localized at the center of magma focusing (the Tuolumne
Intrusive Complex) were not only voluminous, but also hydrous, evolved, and readily able to mix. These
arc behaviors thus appear to have effectively primed the arc thermally and mechanically, culminating in
mid- to upper-crustal mixing, assimilation, storage, and homogenization zones, large, dynamic magma
chambers, and potentially modifying the physical and chemical connections between plutons and
volcanoes. Non-steady-state behavior in the Sierra Nevada resulted in mountain-building and voluminous
(new) continental crust formation by transforming the physical, thermal, and chemical properties of the
lithosphere over a relatively short duration (10’s m.y.). I wrote the first draft of the paper and drafted figures.
S.A drafted Figures 2 and 3, and H. Garcia Carmen. drafted Figure 17. This chapter will be submitted to
Geochemistry, Geophysics, Geosystems.
Chapter 4: At Tioga Pass, in the eastern-central Sierra Nevada, a proposed Triassic caldera (Schweickert
and Lahren, 1999) is re-interpreted as a Late Cretaceous hypabyssal complex, bridging plutonic and
volcanic magmatic fields. New field mapping, U-Pb zircon ages, and whole-rock element and isotopic
geochemical data from Tioga Pass were collected to reconstruct magmatic and tectonic histories from
Triassic to Late Cretaceous time. Both the Tioga Lake quartz monzonite and intrusive dacite-rhyolite
porphyry are ca. 100 Ma and represent co-magmatic parts of a punched laccolith (steep-sided and flat-
topped, with vertical walls discordant to host rock structure; Corry 1988) that intruded blocks of faulted
and tilted Triassic and Jurassic volcanic and sedimentary strata. The Tioga Pass system shares
compositional and structural affinity with other Late Cretaceous intrusive and volcanic rocks in the east-
central Sierra Nevada region and is one of several hypabyssal intrusions found across a ~50-kilometer-wide
belt in the central Sierra Nevada arc section. These intrusions are key structural markers and may represent
feeders to volcanic eruptions or stalled late melts from plutons. All authors were involved in fieldwork. I
collected age and geochemical data (with the assistance of A. Angulo, J. Ayers, and K. Werts) and wrote
the first draft of the paper. All authors (K.A, V.M and S.P) were involved in editing the manuscript for
publication. This chapter has been accepted for publication in Lithosphere and is currently in press.
Chapter 5: This study examines compositionally defined magmatic structures in the TIC from field,
structural and geochemical perspectives to investigate the evolution and dynamics of an upper crustal
magma storage region at scales ranging from individual structures to the pluton-wide dimensions. In
summary, results indicate that schlieren-bound structures formed by physical flow-sorting in an active
crystal-mush environment, triggered by local flow instabilities. Dense minerals were selectively
accumulated, to varying degrees, along both steeply- and gently-dipping boundaries. Schlieren-bound
structures may form highly intricate and disorganized patterns at the outcrop scale, reflecting the complexity
of crystal-melt and crystal-crystal interactions in a hydrogranular medium (e.g., Bergantz et al., 2017).
8
Within mappable domains, the clustering of structures reflects the spatial and temporal heterogeneity of the
magma mush, where fluidized (more mobile crystal mush) regions promote structure formation and
preservation, and less dynamic zones are devoid of schlieren-bound structures. At the regional scale,
schlieren-bound structures show weak to moderate alignment with contacts and outward younging
directions that we propose are the result of internal return flow and convection of the mush. These patterns
can be used to estimate the maximum and minimum sizes of active magma chambers. All authors were
involved in fieldwork. I compiled data, wrote the first draft of the paper and drafted figures. J.S, J.K, and
S.C compiled field data and drafted figures. K.A, S.P and P.A were involved in editing and revision to
prepare the manuscript for publication. This chapter has been accepted for publication in Frontiers in Earth
Science: Petrology and is currently in press.
Chapter 6 concludes with a discussion of the broad impacts of this study and integrates findings from
pluton to arc-wide scales.
During my Ph.D. I have been involved in additional studies as a co-author. I have collaborated with
colleagues on topics that have broadened my understanding of arc systems. These articles are found in
Appendix (A). They include: a review of mesoscopic magmatic structures in plutons and their significance
(Paterson et al., 2018); investigating a zone of spectacularly magmatically deformed schlieren, driven by
magma avalanching within a solidification front (Alasino et al., 2019); and presenting a new way to quantify
the amount of crystal accumulation in plutonic rocks using hornblende compositions (Barnes et al., 2020).
Appendix (B) includes the supplementary data to Chapter 2, Appendix (C) includes the supplementary data
to Chapter 3, including a compilation of whole-rock geochemistry that is referred to in chapters 2, 4, and 5.
Appendix (D) includes the supplementary data to Chapter 4, and Appendix (E) includes the supplementary
data to Chapter 5.
9
Chapter 2: Spatiotemporal magmatic focusing in upper-mid crustal
plutons of the Sierra Nevada arc
This chapter is published in Earth and Planetary Science Letters
Ardill, K., Paterson, S. and Memeti, V., 2018. Spatiotemporal magmatic focusing in upper-mid crustal
plutons of the Sierra Nevada arc. Earth and Planetary Science Letters, 498, pp.88-100
(doi: 10.1016/j.epsl.2018.06.023)
Abstract
In the upper crustal sections of arcs, spatiotemporal focusing of magmatism involves the inward
migration and younging of many volcanic centers, as well as homogenization of erupted magma, over
spatial scales of ~10
2
to 10
5
km
2
and temporal scales of 10
5
-10
7
yr. Magma focusing occurs during, and may
be a consequence of, magmatic flare-ups in arc systems. The occurrence of magmatic focusing is well
documented in volcanic systems but has been largely inferred for the underlying plutonic footprints. At the
plutonic level, focusing predicts similar inward migration and younging of plutonic bodies and the growth
of larger, long-lived magma chambers in which magma compositions shift towards “monotonous” biotite-
hornblende granodiorite that potentially feed large volcanic eruptions. In addition, focusing predicts
increased magma devolatilization, the formation of ore deposits, and increased rates of emplacement-
related regional deformation. In the central Sierra Nevada, CA, we find: (1) a pattern of spatial and temporal
focusing in a field of Cretaceous plutons between ca. 102 and 85 Ma, with the Tuolumne Intrusive Complex
(TIC) involved in and occurring at the center of this focusing; (2) metamorphic pendants in this same area
preserve volcanic rocks between 102 and 95 Ma involved in focusing; (3) a number of coeval subvolcanic
porphyry intrusions in this area that are potential links between these volcanic and plutonic fields. Plutons
show an inward younging, increase in the size of final plutonic complexes, and trends towards hybrid
compositions, with restriction in the range of
87
Sr/
86
Sr i and εNd isotope ratios. Thus, plutonic focusing
contributes to the construction and longevity of large, central magma chambers that increase the potential
for large-scale fractionation and thorough mixing, and provide a source for large volcanic eruptions.
Although focusing is likely a transcrustal process, this study documents the continuation of magma focusing
from surface to upper-mid crustal levels. We speculate that other Late Cretaceous intrusive complexes in
the Sierra Nevada, California (e.g. Sonora, John Muir, and Whitney intrusive suites) may all reflect focused
systems.
10
1. Introduction
Arc systems are non-steady-state in time and space. At the arc-segment scale (10
3
-<10
4
km length),
magmatism and tectonism wax and wane during flare-ups and lulls (Armstrong and Ward, 1993; Paterson
and Ducea, 2015; Cao et al., 2017), and arc fronts migrate laterally relative to the continental edge (e.g.
Coney and Reynolds, 1977). Concurrent with arc-wide episodicity and migration, arc magmatism
demonstrates additional spatiotemporal trends, such as magmatic focusing (Smith, 1979; Hildreth, 1981;
Lipman, 2007, and references therein; de Silva and Gosnold, 2007).
Ultimately, magma focusing may be a transcrustal process involving partial melting of large mantle to
lower-crustal melt source regions to produce volcanic eruptions at discrete point centers (e.g. Brown, 2007).
However, in this study we examine focusing at surface to upper-crustal levels. At the surface, magmatic
focusing is described as the evolution of volcanic centers from initially scattered positions towards a central
area (e.g. de Silva et al., 2006; Grunder et al., 2008; Lipman, 2007). In the volcanic record, magmatic
focusing has resulted in voluminous ignimbrite eruptions (10
2
-<10
4
km
3
) of “monotonous intermediate”
composition, ore formation, and increased regional deformation (de Silva et al., 2006; Grunder et al., 2008;
Lipman, 2007). Scales of volcanic focusing range from 10
2
to 10
5
km
2
, (or 10 to 10
3
km length scale) over
durations of 10
5
-10
7
years (e.g. Bacon and Lanphere, 2006; Lipman, 2007; Grunder et al., 2008; Longo et
al., 2010; Kern et al., 2016). Despite a broad range in spatial scale, long-lived, focused volcanic fields share
(to a first order) a common history of temporal and chemical evolution. A popular interpretation of these
trends is that the amalgamation of volcanic centers records the construction and evolution of a transcrustal
magma plumbing system, which makes clear predictions, and has important implications for deeper crustal
levels, including the potential connections between volcanic and upper-mid crustal plutonic rocks (e.g.
Bachmann et al., 2007; de Silva and Gosnold, 2007; Kern et al., 2016; Best et al., 2016).
Consequences of magmatic focusing for the underlying magma plumbing include the increased size and
longevity of magma chambers, leading to the increased potential for magma fractionation, mixing and
compositional hybridization (Lipman, 2007; Grunder et al., 2008). Although many studies have examined
spatiotemporal and chemical trends on the scale of single, zoned plutons (e.g. Bateman, 1992, and
references therein; Lackey et al., 2012), the nature of regional magmatic focusing in the upper-mid crust
involving multiple plutons is not well understood.
In this study, we document that magmatic activity in the central Sierra Nevada (CSN) is well-
characterized by patterns of regional spatiotemporal focusing and compositional hybridization across a
~4,000 km
2
field of >102-85 Ma plutons, with the Tuolumne Intrusive Complex (TIC) formed at the center
of this focusing. Additionally, several pendants in this area preserve volcanic rocks with evidence for
11
explosive eruptions of the same age. Volumetrically minor hypabyssal intrusions (emplaced between 0-6
km depth) match plutonic and volcanic trends, strengthening the connection between different levels of the
magma plumbing system. Thus, the CSN provides the opportunity to characterize deeper levels of magma
focusing and investigate the causes of these spatiotemporal and geochemical trends in the mid-upper crust
that result in the generation of large silicic batholiths and ignimbrite eruptions during arc-wide flare-ups.
2. Background
Arcs at the largest scale (10
3
-<10
4
km length) represent areas of focused magmatism, defined by narrow
belts of magmatism with discrete volcano spacing at convergent plate boundaries (Bremond d'Ars et al.,
1995). At smaller scale (10
1
-10
2
km length), nested, zoned intrusions represent centers of focusing (e.g.
Bateman, 1992). Mass balance calculations predict that these ranges of scales are linked, since enormous
mantle-lower crustal source reservoirs undergoing partial melting are needed to produce the observed
quantity of upper crustal plutonic and erupted volcanic rocks (Brown, 2007; Lake and Farmer, 2015). This
implies that magma is broadly focused vertically through the crustal column, from wide mantle melting
zones to discrete volcanic centers. Magmatism is also focused in time, resulting in episodes of arc flare-ups
separated by lulls (e.g. Armstrong and Ward, 1993).
2.1 Magmatic focusing in volcanic fields
During flare-ups, smaller scale temporal and spatial focusing occurs as best documented, to date, in
volcanic systems. Over the lifespan of several long-lived volcanic fields (10
5
-10
7
yr), the position of
volcanic centers transitions from initially scattered towards a focused, central area (Fig. 2.1A). Examples
from South America include the Altiplano-Puna Volcanic Complex (e.g. de Silva et al., 2006), the
Aucanquilcha Volcanic Cluster (e.g. Grunder et al., 2008), and the Yanacocha Volcanic Field (e.g. Longo
et al., 2010). North American examples include the San Juan magmatic locus (e.g. Lipman, 2007) and the
southern Great Basin ignimbrite province (e.g. Best et al., 2016). Synchronous with magmatic focusing,
several other factors evolve with time. Geochronologic data in conjunction with petrologic trends and
volume estimates for these systems highlight common attributes in eruptive history and compositional
evolution, discussed further below, that appear to operate across a range of spatial scales (~10
2
to 10
5
km
2
;
Fig. 2.1B):
On average, centrally-located volcanic centers erupt considerably higher volumes of magma (10
2
-<10
4
km
3
; Fig 2.1B) than their scattered predecessors, defining a phase of peak volcanism (Lipman, 2007; de
Silva and Gosnold, 2007; Grunder et al., 2008; Fig. 2.1B). Early, low-volume volcanism characterized by
compositionally diverse, crudely bimodal lavas and tuffs (basaltic-andesite to rhyodacite) shifts towards
12
Figure 2.1: Characteristics of magmatic focusing in long-lived volcanic fields compared to the central Sierra
Nevada (CSN) plutonic field. (A) Spatiotemporal map patterns within volcanic systems and the CSN across a broad
range of spatial scale (10²-10⁵ km²). All maps are drawn to the same scale. Lighter colors indicate younger ages.
SJML= San Juan Magmatic Locus. Thick dashed line outlines calderas (in blue) and red areas indicate resurgent
plutonic material associated with them. The extent of Fish Canyon Tuff is marked by the dashed pattern. Modified
from Lipman (2007). AVC= Aucanquilcha Volcanic Cluster. Purple areas indicate dated volcanic deposits (lavas
and ignimbrite deposits). Modified from Grunder et al. (2008). YVF=Yanacocha Volcanic Field. Orange areas
indicate dated volcanic deposits. Modified from Longo et al. (2010). CSN map modified from Huber et al. (1989).
(B) Time-normalized plots showing cumulative volumes of magma erupted for the YVF, AVC and SJML volcanic
systems, and cumulative plutonic volumes in the CSN. Volumes of CSN plutons shown in Figure 5 and included in
the supplementary materials. Circles show main magmatic events. Stars show hydrothermal events linked to ore
formation. For the SJML, the filled plot shows ignimbrite eruption volume and the curved line shows erupted lava.
The dashed line represents the onset of dominantly hornblende-biotite dacite (a.k.a., monotonous intermediate).
Adapted from Grunder et al. (2008). Data sources: Riciputi et al. (1995); Lipman (2007); Grunder et al. (2008);
Longo et al. (2010); Walker et al. (2010); Lipman and Bachmann (2015).
13
compositionally restricted eruptions of amphibole-biotite dacite at the onset of peak volcanism (e.g.
Grunder et al., 2008; Lipman, 2007; dashed line in Fig. 2.1B).
The largest eruptions, often of “monotonous intermediate” type, are phenocryst-rich and uniform in bulk
composition (between 63-71 wt.% SiO 2; Hildreth, 1981; Bachmann et al., 2007). They appear largely
chemically un-zoned at the kilometer scale but are texturally heterogeneous at the centimeter-millimeter
scale (e.g. Bachmann et al., 2002), indicating that the tuffs are crystal mixtures (Bachmann et al., 2002;
Kern et al., 2016). In some cases, evidence for the resorption and recrystallization of older zircons shows
that magmas homogenized in the subvolcanic mush before eruption (Walker et al., 2010; Lipman and
Bachmann, 2015; Kaiser et al., 2016). In other cases, volcanic products appear to carry a complex crystal
cargo resulting from open-system behavior at multiple crustal depths (e.g. Chambefort et al., 2013; Walker
et al., 2013; Kern et al., 2016). Mineral recycling of zircon antecrysts and rare xenocrysts may be more
prominent in low-volume, early eruptions (Lipman and Bachmann 2015; Walker et al., 2010). Isotopic
hybridization is most extensive in the magma chambers that feed large ignimbrite eruptions (e.g. Riciputi
et al., 1995). The formation of shallow hydrothermal systems and the generation of economic mineral
deposits predominantly postdates the main phase of ignimbrite volcanism and defines the waning stage of
volcanism (Longo et al., 2010; Lipman, 2007; stars in Fig. 2.1B).
Magmatic focusing inferred in long-lived volcanic fields is often interpreted as the surface expression
of the growth and maturation of a vertically extensive magma plumbing system (e.g. de Silva and Gosnold,
2007; Lipman, 2007; Grunder et al., 2008); in some cases, the remnants of this plumbing have been
geophysically imaged, revealing a voluminous batholith, or active magma-mush at depth (e.g. Lipman
2007; Ward et al., 2017). The spatiotemporal and geochemical evolution described above for volcanic
systems predicts several patterns in the underlying magma plumbing (Bachmann et al., 2007; Grunder et
al., 2008; Kern et al., 2016). These patterns include: (1) The younging of small, initially scattered plutons
towards a large, central batholith; (2) an inward increase in magma addition, resulting in larger magma
chambers and final intrusive complexes; (3) the transition from compositionally diverse magmas to
hybridized, ‘monotonous’ bodies which are uniform in appearance and composition at the hand sample
scale, though they may contain multiple, mixed crystal populations; (4) the inward transition towards
isotopically homogeneous magmas; (5) an inward increase/concentration of volatiles and late formation of
ore deposits. Plutons of the central Sierra Nevada are examined in regard to these postulated patterns.
2.2 Sierra Nevada Batholith
The Sierra Nevada arc is the type locality for subduction-driven, calc-alkaline magmatism, built across
a transitional oceanic to continental margin (Kistler, 1990). Magmatism is episodic, with Late Triassic,
14
Middle Jurassic and Late Cretaceous magmatic flare-ups (Paterson and Ducea, 2015, and references
therein). During the Late Cretaceous flare-up, the largest of the magmatic episodes (e.g. Paterson and
Ducea, 2015), the crust thickened (e.g. Profeta et al., 2015), and the locus of magmatism migrated eastward
(e.g. Chen and Moore, 1982). These arc-wide spatiotemporal and chemical trends as recognized in the
Sierra Nevada batholith (SNB) by earlier studies (Moore, 1959; Chen and Moore, 1982; Bateman, 1992;
Profeta et al., 2015) are briefly summarized below, before describing local spatiotemporal patterns within
the CSN:
3. Arc-wide migration patterns
To define arc migration patterns that represent the spatial expression of the Cretaceous arc flare-up, an
ENE-WSW transect was drawn perpendicular to the long-axis of large intrusive bodies and the Sierran
range between 37.5° and 38.13° latitude (Fig. 2.2). The same approach was taken as previously
demonstrated by Chen and Moore (1982) between latitudes 36° and 37°, and Cecil et al. (2012) between
latitudes 38.5° and 39.5°. The western boundary of the Cretaceous SNB is defined by 140-130 Ma mafic
cumulates, diorite, and tonalite sampled in drill cores from the Great Valley (Saleeby, 2007). The eastern
boundary of the transect is the Owens Valley. New and published Cretaceous U-Pb zircon ages for the CSN
were compiled and plotted by sample location (references included in Supplementary Tables S1 and S2 and
Supplementary Fig. S1 in Appendix C). Where samples did not lie on the transect line, the sample location
was projected along a range-parallel, transect-perpendicular line.
Between ca. 140 Ma to 85 Ma, most plutonic samples form an eastward younging trend, with a best-fit
migration rate of 2.6 km/myr. This is consistent with migration rates estimated north and south of the study
area, between 2.7 km/myr and 4.8 km/myr (Chen and Moore, 1982; Cecil et al. 2012; Fig. 2.2). Exceptions
include the Iron Mountain area and eastern parts of the Fine Gold Intrusive Suite (FGIS in Figs. 2.2, 2.3),
where younger plutons surround much older bodies, and in the region around the TIC, the focus of this
study (light grey box in Fig. 2.2; Fig. 2.3A).
West-east variations in pluton composition across the SNB have been interpreted to result from lateral
compositional changes of the pre-Mesozoic basement into which the granitoids intruded (e.g. Moore, 1959;
Bateman, 1992). Compiled geochemical data of Cretaceous plutons illustrates eastward trends consistent
with those reported south of the study area by Bateman (1992) for all Mesozoic intrusions. In summary,
these patterns are as follows (described from west to east): Cretaceous plutons show a large spread in SiO
2
contents (46-77 wt.%), with no lateral variation (diagrams included in Supplementary Fig. S2). K 2O on the
other hand, increases, and together with decreasing Al 2O 3 and CaO, has been attributed to a modal increase
in the proportion of alkali feldspar: plagioclase ratio (Moore, 1959; Bateman, 1992). Other major and trace
15
Figure 2.2: Graph of pluton age vs. distance across the Sierra Nevada Batholith, between latitudes 37.5°N and
38.13°N. Transect line marked by red line in inset map (map adapted from Jennings, 1977). In inset map, light grey
represents metamorphic host rock pendants, and dark grey represents plutonic material. Dark grey box in lower left
corner defines ages of granitoids beneath the Great Valley Sequence, sampled in drill cores (Saleeby, 2007). U-Pb
zircon ages from plutons are shown as points. Eastward migration rate approximated by wide-dashed line, ~2.6
km/myr (trend 2; this study). Trend 1 (purple dot-dashed line) is the ~2.7 km/myr average migration rate by Chen
and Moore (1982) between 36°N-37°N. Trend 3 (green dotted line) is the ~4.8 km/myr average migration rate by
Cecil et al. (2012) between 38.5°N-39.5°N. Light grey box in upper right corner defines the area of magmatic
focusing based on the deviation of points from the approximate best fit line. Red ring circles plutons from the Iron
Mountain area that also deviate from the migration trend. Blue ring circles samples of the Bass Lake Tonalite (Fine
Gold Intrusive Suite) from Bateman (1992) and Lackey et al. (2012). Explanation in text. Data sources for
geochronology compilation included in the supplementary materials.
16
Figure 2.3: Spatiotemporal magmatic focusing in the central Sierra Nevada, between 37.5° and 38.13°. (A) Spatiotemporal map patterns within
the central Sierra Nevada (CSN). Map modified from Huber et al. (1989) and Paterson et al. (2016). Plutons, hypabyssal bodies and volcanic
packages are colored by age. Lighter colors represent younger ages (see key). The uncolored area represents earlier Mesozoic volcanic/plutonic
rocks, Cretaceous plutons and volcanic rocks not involved in magmatic focusing, e.g. Fine Gold Intrusive Suite (FGIS), El Capitan (EC), Iron Mtn
pendant (IM). Units of the Tuolumne Intrusive Complex are labelled as follows: Kuna Crest Lobe (KCL); Equigranular Half Dome (EHD);
Porphyritic Half Dome (PHD); Cathedral Peak (CP); Johnson Peak Porphyry (JPP).
87
Sr/
86
Sr 0.706 line drafted from Kistler (1990) and Lackey et
al. (2012). Stars indicate the location of U-Pb zircon ages (various methods: see key in Fig. 2.2 for explanation). (B) Interpreted isochrons within
the CSN showing the shrinking area of magmatism through time. Magmatic focusing rates estimated from map pattern as shown by dashed lines.
Data sources for geochronology compilation included in Supplementary Tables S1, S2 and Figure S1.
17
elements have weak trends or lack trends (Fig. 2.4). Sr/Y increases from ~25 to 85 and using the method
of Profeta et al. (2015) records arc-wide crustal thickening from 30 km to >75 km at rates of ~1 km/myr
(Fig. 2.4D), and La/Yb suggests comparable behavior (not shown). Whole rock Nd, Sr and Pb isotopic
ratios of plutons trend towards more evolved signatures (e.g. Kistler, 1990; Fig. 2.4E, F). These isotopic
patterns are interpreted to record the migration of magmatism across laterally varying basement terranes
(Kistler, 1990). However, a decrease in whole rock and zircon δ18O from ~7.5‰ to 5.5 ‰ towards the
central and eastern parts of the transect requires a significant mantle contribution (De Paolo, 1981; Lackey
et al., 2008), and thus suggests a more complicated relationship.
In comparison to the spatial, temporal and geochemical arc-wide patterns associated with arc migration
and the Cretaceous magmatic-tectonic flare-up, additional patterns are now discernable at a finer spatial
and temporal scale. Below, we present the spatial (~10
3
km
2
), temporal (10
7
yr) and geochemical patterns
in plutonic, hypabyssal, and volcanic units of the CSN. We then evaluate these records from a perspective
of focused volcanic systems and discuss the importance of crustal-scale effects versus upper-crustal
processes operating in focusing systems.
4. Spatiotemporal patterns in the central Sierra Nevada
4.1 The plutonic record
At 102 Ma, a deviation of pluton ages away from the projected arc migration trajectory is noticeable in
the CSN. Several plutons, 100-94 Ma, are younger in the west and older in the east than predicted by arc-
wide migration rates (Fig. 2.2). Plutons <94 Ma are increasingly spatially restricted through time, and do
not fit the arc-wide eastward migration pattern. Thus, the inward-younging zone of 102-85 Ma plutons
defines the area of spatial focusing (grey shaded area in Fig. 2.2).
In the CSN, plutonic, hypabyssal, and volcanic rocks are broadly age-zoned, with the Tuolumne
Intrusive Complex (TIC) recording the innermost inward and northeast pluton younging (Fig. 2.3A;
Bateman, 1992; Memeti et al., 2014). The spatiotemporal pattern of inward younging is most clearly
demonstrated in the ca. 102-85 Ma plutonic record, and weakly defined for volcanic and hypabyssal levels
(Fig. 2.3A). Several small, ca. 98-93 Ma plutonic bodies exposed east of the TIC demonstrate a younging
direction opposite to the eastward arc migration patterns. The center of the focused region is the NE corner
of the TIC where the youngest U-Pb zircon ages are reported from evolved, peraluminous granites of the
Cathedral Peak (Paterson et al., 2016; Fig. 2.3B). Since the inward focusing is asymmetric, the rate of
inward migration from margins to the NE TIC varies between ~0.7-4.2 km/myr (Fig. 2.3B), which is
comparable to arc-wide migration rates of ~2.7 km/myr. Rates of inward younging in the CSN are highest
in the west, suggesting that arc-wide migration contributes to faster inward-younging. Where inward-
18
Figure 2.4: Whole rock trace element and isotopic ratios vs. age between 37.5° and 38.13° latitude. Circles
represent plutonic samples, triangles represent volcanic samples, squares represent hypabyssal samples. The x-axis
represents the eastward distance along the migration-parallel transect (method outlined in section 3). Colors
represent sample age. Age bins are in 5 Myr intervals, except for the period 102-100 Ma, which has been expanded
to show the start of focusing at ~102 Ma. Where sample age fell on the boundary of two age bins, the sample was
moved to the older age bin. Best-fit line shown as a black dashed line and 95% confidence intervals of the fit are
represented by the grey field. Grey vertical dashed line represents the approximate western boundary of inward
spatial focusing. Data sources listed in Supplementary Tables S3 and S4. (A) Rb/Sr vs. eastward distance; (B) Zr/Hf
vs. eastward distance; (C) Dy/Yb vs. eastward distance; (D) Sr/Y vs. eastward distance. Data are compositionally
filtered using method of Profeta et al. (2015); (E) 87Sr/86Sri vs, eastward distance; (F) ƐNd vs. eastward distance. In
(E) and (F) arrow represents the lateral change in basement type proposed by Kistler (1990).
19
younging is opposite to eastward arc migration, focusing is slower ~1 km/myr, and represents a minimum
rate.
Long and short axes of exposed CSN plutons are measured in map view and combined with likely
maximum and minimum vertical extents to calculate volumes of intruding plutons through time (techniques
discussed in Karlstrom et al., 2017). The largest uncertainties in the volume calculations are found in: 1)
estimating the vertical extent of plutons using relief, barometry, and the orientation of pluton contacts; and
2) recognizing the amount of plutonic material removed by subsequent intrusions (reconstructed pluton
volumes). To address 1), we assigned plutons a “best-fit vertical extent” from the Karlstrom et al. (2017)
pluton size database to represent a conservative interpretation of pluton volume.
To address 2) in the nested TIC, we included reconstructed volumes following Paterson et al. (2016).
For 102-95 Ma plutons, available geochronology and map data preclude an accurate estimate of
reconstructed volumes. The dataset of CSN pluton size measurements is included in Supplementary Table
S5.
A 102-95 Ma ring of plutons with calculated volumes between 3-1,000 km
3
preceded the intrusion of
the centrally located 95-85 Ma, ~11,000 km
3
TIC (total present-day volume; Fig. 2.5). Small volume
plutons and magmatic sheets were emplaced throughout focusing (square symbols in Fig. 2.5), yet the
maximum volumes of magma bodies emplaced through time increased by one order of magnitude during
the emplacement of the TIC (circles and hexagon symbols in Fig. 2.5). As described above, maximum
volumes of older intrusions (>95 Ma) may be underestimated due to subsequent intrusion, removal or
recycling during TIC construction. However, older intrusions in the study area could not have been more
voluminous than the TIC without some part having been preserved. 102-95 Ma plutons that have well-
constrained areal sizes (i.e., are not re-intruded) are one to two orders of magnitude smaller than the TIC.
A synthesis of whole-rock geochemical data from the study area (references in Supplementary Tables
S3 and S4) is used to track temporal trends in magma composition as plutons migrate inwards and pluton
volumes increase. Trends are partly decoupled to the spatial pattern of inward focusing. In any single pluton
there are several magmatic signals that the whole-rock chemistry may record: 1) arc-wide flare-up and
migration processes; 2) processes driving inward spatial focusing; and 3) intra-chamber fractionation,
mixing or contamination. Thus, complexity in regional-scale compositional trends through time is expected.
To address this problem, SiO
2 is used as a first-order differentiation indicator. To minimize the effects
of sampling bias and link chemistry to spatial pluton size patterns, SiO 2 is estimated (in 10 wt.% bins) by
area and pluton age (pluton age binned in 5 myr increments; see Fig. 2.3A). Pluton area is calculated, rather
20
Figure 2.5: A graph of calculated volume (km³) vs. pluton age (Ma) in the CSN focusing region. Colors are used to
differentiate calculated volumes from each TIC unit. KC=Kuna Crest, HD = Half Dome, CP=Cathedral Peak,
JP=Johnson Peak Porphyry. Colored fields display estimates of magma chamber sizes in the TIC given in Memeti et
al. (2010a) and Paterson et al. (2016). The calculation of magma chamber sizes in Paterson et al. (2016) involves
restoration of pluton volumes that were removed and/or recycled into younger units (Open symbols; Paterson et al.,
2016). Pluton volumes are calculated using the method outlined in Karlstrom et al. (2017), where initially map areas
of plutons are measured, then multiplied by estimated thickness (i.e., vertical extent of the pluton). Constraints for
vertical extent include: exposed relief, the dip of pluton contacts, Al-in hornblende barometry, regional geology and
geophysical data where available. Following estimation of minimum and maximum volumes, a “best fit” vertical
extent was calculated and plotted. Volume calculations are made at various spatial scales, from measurements of
individual sheets (3rd order measurement) to individual mapped magma bodies (2nd order measurement) to volumes
of entire units (the sum of genetically related mapped magma bodies; 1st order measurement). Each type of volume
measurement is ranked by order and designated a different symbol. The dashed line indicates the maximum volume
of magma bodies for a given time.
21
than volume in this case, as the compositional variation beneath the exposed surface of the intrusive
complexes is unknown.
Mafic intrusions <55 wt. % SiO 2 are rare and of low volume (<10 km
2
; Fig. 2.6) and largely absent
during emplacement of the TIC (e.g. Paterson et al., 2016). Between 95-85 Ma the area of 65-75 wt.% SiO 2
in plutons increases, with a corresponding increase in abundance of >75 wt.% SiO 2 plutonic material, while
55-65 wt.% SiO 2 compositions are not found after 90 Ma. This suggests increasingly more intermediate to
felsic magmas were emplaced in the CSN through time (Fig. 2.6). The compositional uniformity of inner
TIC units (Half Dome granodiorite and Cathedral Peak biotite-hornblende granite) at the outcrop- to map-
scale supports these trends in the data.
Minor and trace elements generally lack clear temporal trends. Rb/Sr varies little through time and
median values are consistently <0.5 up to the start of inward focusing, except in high-silica rocks, such as
aplite dikes and leucogranites, which have Rb/Sr values >2 (Fig. 2.4A). Between 102-85 Ma, Zr/Hf
decreases from 102-85 Ma, indicating more extensive fractionation of zircon through time (Fig. 2.4B).
Maximum values of Dy/Yb also decrease, while minimum values increase (Fig. 2.4C). Generally, element
ratios are more scattered between 102-85 Ma than in earlier intrusions. Spatial focusing of plutons between
102-85 Ma correlates with increasingly radiogenic
87
Sr/
86
Sr i (median from 0.7059 to 0.7065; Fig. 2.4E) and
less radiogenic εNd (median from -2.2 to -6.2). The spread in isotopic values in
87
Sr/
86
Sr i and εNd is more
restricted through time (Fig. 2.4F).
87
Sr/
86
Sr i and εNd within the focusing zone converge to values that are
less evolved than predicted by west to east migration alone (arrow in Fig. 2.4E and F).
Previous studies have demonstrated that from map to mineral scale the TIC preserves evidence for
magma fractionation, mixing, and recycling, resulting in compositional hybrid zones and truncations of
contacts between TIC units, in addition to outcrop-scale structures (Paterson et al., 2016 and references
therein). Within a single sample from some TIC domains, multiple populations of certain minerals are found
(e.g. Memeti et al. 2014; Barnes et al. 2016). For example, zircon antecrysts are rare in the Kuna Crest outer
unit but increase in abundance within the inner units of the Half Dome and Cathedral Peak (Miller et al.,
2007; Paterson et al., 2016). Rarely, zircon xenocrysts are found within the TIC with ages of nearby host
rock units (Memeti et al., 2010a). These observations document at least local recycling of older plutonic
material (<95 Ma) in the TIC (Paterson et al., 2016). It remains unclear how much recycling occurred in
the surrounding 102-95 Ma plutons.
22
Figure 2.6: SiO 2 variation with time. The range of SiO 2 (binned in 10 wt. % intervals) is estimated for each pluton
using the geochemical samples with map locations and field observations (e.g. the Cathedral Peak unit is
approximately 85% SiO 2 between 65-75 wt.% and 15% >75 wt.%). Each subarea within each pluton, defined by
SiO 2 contents, is assigned a map area (km²) using measurements from geological maps, tables included in
Supplementary Table S5. Plutons are next binned according to age, in 5 Myr intervals (except for the first bin which
is a 2 Myr interval and represent the start of spatiotemporal focusing). Gabbros <55 wt.% SiO 2 are not found in 90-
85 Ma CSN plutons.
23
Qualitative indicators of volatile evolution between 102-85 Ma include the increasing abundance
through time of hydrous mineral assemblages and pseudomorphic reactions, as well as pegmatitic vugs,
miarolitic cavities, and dikes, and the record of late δ
18
O isotopic exchange (e.g. Bateman, 1992; Memeti
et al., 2014; Lackey et al., 2008). However, more detailed studies are needed before this record can be
linked to the spatiotemporal evolution described here.
4.2 The volcanic and hypabyssal record
Cretaceous volcanic rocks in the CSN are relatively scarce in the CSN due to erosion of the uppermost
arc section (6-10 km removed; Ague and Brimhall, 1988); however, volcanic sections are preserved in the
Saddlebag Lake, Ritter Range, and Piute pendants of the CSN, contemporaneous with nearby plutons and
hypabyssal intrusions (e.g. Fiske and Tobisch, 1994; Memeti et al., 2010b; Fig 2.3A). Shallow-level (0-6
km) hypabyssal intrusions represent a transitional zone between upper crustal plutons and erupted materials
and may provide additional information to investigate plutonic-volcanic spatiotemporal trends, particularly
in the absence of a complete volcanic record.
The fairly-intact 100-98 Ma Minarets and Merced Peak calderas in the Ritter Range pendant, 600 km
2
each in size, are an example of the scale of explosive volcanism in the CSN (Fig. 2.3A; Fiske and Tobisch,
1994). The size of the complexes, and the ~2.3 km thick section of caldera fill (tuff and collapse breccia),
preserved within the Minarets caldera suggests ignimbrite activity was voluminous by ca. 100 Ma (e.g.
Fiske and Tobisch 1994; Tomek et al., 2017). No direct evidence has yet been found for volcanism
associated with the construction of the 95-85 Ma TIC and other, contemporaneous intrusive suites within
the wider Sierra Nevada batholith, likely owing to exhumation of the arc to 6-10 km depths. However,
bentonite beds in the western interior (Nevada/Utah), relicts of 88-90 Ma (K-Ar in biotite) ash-fall deposits,
are interpreted to have been sourced from volcanic centers in the Sierran arc (Elder, 1988) and further
illustrates the scale of the eruptive activity that continued into the Late Cretaceous (Lipman, 2007).
Hypabyssal intrusions are proportionally smaller than both plutonic and volcanic exposures (at a 1:30:5
ratio), with measured areas in the CSN <30 km
2
, and little change through time. However, there are
indications that they may also provide an indirect record of volcanism. The largest hypabyssal intrusion,
the ca. 88 Ma Johnson Peak porphyry (JPP; Fig. 2.3A), within the TIC, has been interpreted to have at least
partly erupted (Bateman, 1992; Titus et al., 2005). Cretaceous volcanic packages host mineral deposits (e.g.
Bateman, 1992), however a full temporal record is limited by the ‘preservation filter’ that indicates most
deposits formed above 6 km depth have been eroded (Barton, 1996).
Cretaceous volcanic deposits range from andesite to rhyolite in composition and have SiO 2 ranging
between 52-78 wt.%. Values of Zr/Hf, Dy/Yb and Sr/Y generally overlap with plutonic samples of the same
24
age with an increase in the number of evolved volcanic samples (Fig. 2.4). Rb/Sr is notably increased in
volcanic samples relative to plutons. Hypabyssal samples plot closely with volcanic compositions in the
CSN. SiO 2 content of hypabyssal intrusions are predominantly in the intermediate-felsic range. Radiogenic
isotope data (
87
Sr/
86
Sr i and εNd) from volcanic deposits and hypabyssal intrusions overlap with plutonic
values (Fig. 2.4E, F). One exception is the Johnson Peak Porphyry, which shows enriched Sr i compared
with all other samples in the CSN (Fig. 2.4E).
5. Discussion
5.1 Spatiotemporal evolution in the upper crust (6-10 km)
Local spatiotemporal patterns in 102-85 Ma CSN plutons developed during the arc-wide Late
Cretaceous flare-up, eastward arc migration, and crustal thickening. While arc-wide migration and local,
inward-focusing occurred at similar rates, between 1-4 km/myr, their spatial extent and patterns through
time are distinct. For example, in the eastern CSN, the regional inward-younging of plutons is directly
opposite to eastward arc-wide migration. The Late Cretaceous arc (120-85 Ma) evolved to an increasingly
dextral-transpressive tectonic regime, which argues against an extensional process creating the broadly
symmetrical spatial patterns in the CSN plutons and supports a magmatic origin for the spatiotemporal
evolution (Fig. 2.3A; Cao et al., 2016).
The local spatial migration of magmatism in the CSN is coincident with an increase in the volume of
magma emplaced in the upper crust. Calculated volumes of intrusive bodies combined with field
observations, geochronology, bulk and mineral geochemistry, and thermal modeling, where available, are
used to estimate the extent and duration of resulting paleo-magma chambers, defined as interconnected
crystal-melt mixtures (e.g. Paterson et al., 2016). Minimum magma chamber volumes consisted of small
ephemeral pulses of magma that froze quickly, forming sheets or compositionally distinct batches ~1 km
3
in size (Memeti et al., 2010a; Paterson et al., 2016). Maximum chamber volumes have been estimated for
the TIC in previous studies, and range in volume from ~100-4,000 km
3
, with hypersolidus durations of <0.5
myr to >1 myr (Fig. 2.5; Memeti et al., 2010a; Paterson et al., 2016). Assuming an ephemeral magma
chamber existed in >95 Ma plutons, with a maximum size equal to the inferred volume of the pluton, the
maximum volume of estimated magma chambers, and final intrusive complexes, increased through time
(fields in Fig. 2.5).
A consequence of larger, long-lived magma chambers is the increased potential for fractionation and
mixing processes (e.g. Hildreth, 1981; Memeti et al., 2010a, 2014; Barnes et al., 2016). In the CSN, there
is evidence that bulk-rock magma compositions (to a first order) become increasingly monotonous (i.e.,
compositionally uniform at the outcrop- to map-scale, but not necessarily homogeneous at the mineral
25
scale) within larger magma chambers through time. Granitoids with SiO 2 in the 65-75 wt.% range are an
order of magnitude more abundant than all other compositions by 90 Ma. The outcrop-scale uniformity in
composition observed across the inner TIC units additionally attests to a hybridization process that acts on
the magma plumbing system across 10’s km and over millions of years. Alternatively, this pattern could be
attributed to hybridization of the source or processes as magma ascends (Hildreth and Moorbath, 1988;
Coleman et al., 2004).
Furthermore, the convergence of Sr and Nd isotopes to more restricted values by 85 Ma is permissive
of the hypothesis that magma compositions homogenize in larger magma chambers, particularly since arc-
scale, west-to-east migration predicts more evolved compositions through time than is seen in the regional-
scale CSN dataset (Fig. 2.4E, F). The change in isotopic signature between 130-102 Ma and 102-85 Ma
may be attributed to the development of a mature MASH zone below the emplacement level (e.g. Hildreth
and Moorbath, 1988), or homogenization in the crustal column due to established magma ascent pathways,
rather than a change in lithospheric basement type that can explain arc-migration signatures (e.g. Kistler,
1990). The variability in isotope compositions between coeval intrusive complexes requires multiple
magma sources. The relatively restricted range in Sr and Nd isotopes by 85 Ma may result from assimilation
of the surrounding older intrusive material, which is isotopically closer to the rising magma than older
metamorphic arc framework units (e.g. Kistler, 1990; Lackey et al., 2008). Whether this compositional
signature is attained primarily at the emplacement level, during ascent, or at depth in the source region (e.g.
Coleman et al., 2004) remains an important open question.
Minor and trace element patterns in CSN plutons are complex and non-unique, representing the
integration of arc-wide migration and local spatiotemporal patterns (Fig. 2.4). Element trends are sensitive
to local magma chamber processes, such as fractionating assemblages, and/or recording evolving arc
boundary conditions (e.g. crustal thickness, assimilation of lithospheric basement, slab properties). It is
likely that the overall chemical signature results from a combination of these components that we are unable
to resolve fully at the regional scale at this time. The volcanic and hypabyssal record in the CSN is one way
to investigate internal magmatic changes through time and across the vertical dimension. For example,
elevated Rb/Sr values from hypabyssal and volcanic rocks within the CSN suggest that they underwent
additional fractionation in their upper crustal magma chambers compared to the record of contemporaneous,
deeper plutons (Fig. 2.4A).
Some TIC units contain multiple crystal populations, suggesting that at least locally, magmas are
complex crystal mixtures where crystals have been sourced from older intrusions, wall rocks, or deeper
levels in the magma plumbing system (Memeti et al., 2014; Barnes et al., 2016; Paterson et al., 2016).
26
Voluminous, explosive volcanism contemporaneous with plutonism is documented by the Minarets and
Merced Peak calderas in the CSN and the far-reaching extent of Sierran volcanic ash preserved in bentonite
beds in the Great Basin (Elder, 1988). It is also suggested in partially-vented leucogranites in the TIC
(Bateman, 1992; Titus et al., 2005). Additional studies are needed to track the spatiotemporal history of
magmatic volatiles and ore deposits.
5.2 Comparing focusing at plutonic and volcanic crustal levels
Focused volcanic fields and the CSN plutonic system share a common spatiotemporal magmatic history
of inward younging (10
5
-10
7
yr), concurrent with an increase in the volumes of magma (10
2
-10
5
km
3
)
emplaced at depth or erupted. Areal footprints of focusing can vary by over an order of magnitude,
indicating that inward-focusing is scale-invariant. The volume estimates from the volcanic record are
instantaneous snapshots of the volume of eruptible magma at depth (i.e., minimum magma chamber sizes),
which during the largest eruptions, are of comparable size to maximum TIC magma chamber estimates
(Lipman, 2007; Paterson et al., 2016). One difference is that volcanic systems demonstrate a sharp increase
in the volume erupted during ‘peak volcanism’, while the CSN has a steadier rate of increasing volumes
emplaced over ~10
7
yr, indicating that the upper crustal CSN represents a magma ‘staging area’ from which
eruptions periodically occur (e.g. Tierney et al., 2016; Fig. 2.1B). The peak of volcanic activity coincides
with the formation of compositionally uniform, monotonous intermediate type magmas (Fig. 2.1B). The
TIC also appears increasingly compositionally uniform (at the outcrop- to map-scale) to a first order
estimate (Fig. 2.6). Each of these monotonous units in both volcanic and plutonic systems (e.g. the Fish
Canyon tuff, the Cathedral Peak granodiorite) contains thermally- and compositionally-buffered mineral
assemblages (Lipman and Bachmann, 2015; Kaiser et al., 2016), and multiple mineral populations within
single hand samples in both volcanic and plutonic systems (Walker et al., 2013; Memeti et al., 2014;
Paterson et al., 2016), suggesting the importance of increased magma mixing. We consider the differences
between volcanic and plutonic systems, such as isotopic and minor/trace element patterns, a second-order
feature, which in part may reflect variations between arc settings.
Given the characteristics of magmatic focusing systems outlined above, the CSN system is closely
comparable to the Andean systems such as the Altiplano-Puna Volcanic Complex (APVC; de Silva et al.,
2006), or the Aucanquilcha Volcanic Cluster (AVC; Grunder et al., 2008). Grunder et al. (2008) noted
several commonalities between the magmatic histories of the AVC and the TIC, such as similar temporal
and spatial scales of focusing (~10 myr, ~1,000 km
2
). We propose based on our investigation, that the TIC
is the central portion of a spatially larger (~4,000 km
2
) and longer-lived (~18 myr) plutonic focusing system
in the CSN. Volumes of 100-2,000 km
3
for single eruptions in the APVC are commensurate with calculated
pluton volumes in the CSN. In addition, within the APVC, the total erupted volume of ignimbrites (>15,000
27
km
3
; Best et al., 2016) is comparable to the estimated volume of plutonic material removed from the
Cretaceous CSN upper crustal system (>12,000 km
3
; Paterson et al., 2016), either by recycling material or
by upward ascent and eruption. Thus, the CSN upper crustal system could feasibly have led to eruptions at
the scale of those documented in the APVC. Both Andean and Sierran systems dominantly produce
intermediate to felsic compositions, yet the Andes system appears to have a larger crustal component in the
generation of magmas than interpreted for the CSN (de Silva et al., 2006; Lackey et al., 2008). Within the
largest dacite/granodiorite magma pulses, crystal mixtures are common (Walker et al., 2010; Kern et al.,
2016; Paterson et al., 2016). Both the AVC and the CSN systems document changes in the volatile content
during spatiotemporal focusing, which in part led to ore deposit formation. At the arc-wide scale, both
systems experienced 10
7
yr arc flare-ups and magmatism was coincident with continuous crustal thickening
(e.g. de Silva et al., 2006). In the case of the APVC, arc migration largely preceded magmatism (de Silva
et al., 2006).
6. Proposed models for magmatic focusing
In testing physical and petrologic models for regional magmatic focusing, it is necessary to separate the
driving mechanisms of arc-wide flare ups and arc migration. This paper does not discuss the causes of this
arc-scale episodicity (see DeCelles et al., 2009; Paterson and Ducea, 2015). However, since focusing occurs
during arc flare-up periods (in both volcanic examples and the CSN), we recognize that there may be a link;
the association of magmatic flare-ups and focusing raises the possibility that arc-wide flare-ups in magma
addition create thermal and rheologic feedbacks in the crustal column, driving the organization of the
magma plumbing system (e.g. de Silva et al., 2006, 2015; Bachmann et al., 2007).
Time-space-volume patterns recorded in volcanic systems have previously been attributed to thermal
and mechanical maturation throughout the arc crustal column, which progressively enhances the growth of
larger magma bodies in the upper crust (e.g. Smith, 1979; de Silva and Gosnold, 2007; Karlstrom et al.,
2017; Karakas et al., 2017). Continued magma addition to the crust over 10
5
-10
6
yr thermally ‘primes’ the
crust (de Silva et al., 2006; Karakas et al., 2017). Physical models propose that successive magma pulses
take advantage of preheated passageways, facilitating the ascent of magma to higher levels with time while
also spatially concentrating it (e.g. Hildreth, 1981; Marsh, 1982). Petrologic models invoke the
simultaneous development of a mature MASH zone in the lower crust, which progressively hybridizes
magmas before they ascend to the upper crust (e.g. Hildreth and Moorbath, 1988).
Once magmas arrive at the upper crustal emplacement level, with increasing ambient temperatures
through time, host rocks may deform in the ductile regime (e.g. Jellinek and DePaolo, 2003; Karlstrom et
al., 2017). Physical models propose larger volumes of magma to accumulate in central, thermally weakened
28
zones, the size of which may be further enhanced by magma lensing, re-orienting the magma pathway by
magma stress fields (Whitehead and Helfrich, 1991; Karlstrom et al., 2009) until volumes are attained
where overpressure is readily dissipated throughout the magma chamber and additional growth is
accommodated through continued thermal-rheological feedbacks (>100 km
3
; Gregg et al., 2013; de Silva
and Gregg, 2014). Large volume magma ascent and emplacement is partly accommodated by downward
flow of host rocks returning fertile crustal material towards the MASH zone, which may modify the isotopic
signature of subsequent magma pulses and drives petrologic evolution (Hildreth and Moorbath; 1988;
Paterson and Farris, 2008; Cao et al., 2016). Thus, over time, the entire crustal column is re-worked
physically and chemically.
The spatiotemporal patterns documented in the CSN are broadly compatible with the upper-crustal
processes described above. Our study of plutonic spatiotemporal focusing has additional implications for,
and refinements to, transcrustal focusing models. Firstly, our results provide evidence supporting the view
that spatiotemporal focusing is a multi-level process, occurring at both volcanic and upper-crustal plutonic
levels. Our interpretation of the CSN system is that it represents an upper-crustal magma staging area, from
which volcanic focusing mimics the inward focusing of intrusive magma bodies. We infer this in the CSN
from the subtle spatiotemporal record of volcanic and hypabyssal rocks contemporaneous with plutonism.
This in turn, strengthens the physical connection between volcanic and plutonic levels. Secondly, we note
that emplacement-level mechanisms are necessary to drive the organization of magma chambers into larger,
centralized intrusive complexes from initially smaller, scattered plutons, since both are found at exposed
levels in the CSN. Nesting of younging intrusive units in the TIC supports either a rheological softening
(Jellinek and DePaolo, 2003) or a stress field, pluton capture mechanism (Karlstrom et al., 2009; 2017).
Evaluating the contribution of emplacement-level driving forces, both mechanical and thermal, is an
important future step. Upper-crustal emplacement-level models are difficult to test in the CSN, since
repeated intrusion of plutons typically removed most of the host rock record. However, strain analyses
around plutons combined with field relationships at pluton-host rock contacts (e.g. Cao et al., 2016) suggest
that pluton emplacement did not significantly displace host rocks laterally. Therefore, the dominant
response of the host rocks must have been to move material vertically through the crustal column (Paterson
and Farris, 2008; Cao et al., 2016). Nadin et al. (2016) proposed that the thermal history of the Cretaceous
Sierra Nevada batholith closely tracked eastward arc migration, which in turn determined the deformation
history and the ductile-brittle response of host rocks at the emplacement level. Our initial observations
suggest that the Cretaceous thermal history of the CSN also tracks the west-east arc migration and local
magmatic focusing.
29
7. Magmatic focusing at all crustal levels?
Continental arc systems are vertically stratified based on our understanding of exposed arc crustal
sections (e.g. Miller et al., 2009). Commonalities among crustal sections include: (1) pluton: host rock ratios
increase with depth; (2) strain and deformation increase with depth; (3) magmatism is progressively less
focused with depth, favoring emplacement as irregular bodies and sheets, progressing to a chaotic,
interconnected network of melt channels (Saleeby, 1990; Brown, 2007; Miller et al., 2009). Plutons in the
upper crust in contrast are highly organized into discrete bodies (Saleeby, 1990; Brown, 2007; Walker et
al., 2015; Fig. 2.7). This implies that there is a systematic transcrustal-scale vertical organization of the
magma plumbing system within continental arc sections that finer-scale processes, such as magma focusing,
may be just one aspect of. Geophysical data from the mid-crustal Altiplano Puna Magma Body highlights
the importance of magma focusing in active magmatic systems, where it has been proposed that discrete
magma bodies have coalesced into a single, central ~200 km wide magma reservoir (Kern et al., 2016;
Ward et al., 2017). This central region also corresponds to the area of voluminous volcanism (Kern et al.,
2016). Combining depth-dependent characteristics of arc crustal sections with emplacement-level
mechanisms, we view the CSN section as part of a linked, transcrustal magma transport system. We posit
that magma focusing occurs at all crustal levels.
The above discussion of arc systems suggests that there should be other examples of magmatic focusing
in the SNB. Promising targets to explore include local areas with ages that do not follow arc-wide
spatiotemporal patterns, for example at Iron Mountain (Fig. 2.2). SNB-wide U-Pb zircon age contour maps
suggest that the Fine Gold Intrusive Suite, the Mt. Whitney and Mitchell Intrusive Suites and the Lake
Isabella region of the southern Sierra Nevada all deviate from the eastward migration pattern (Fig. 2.2;
Lackey et al., 2012; Nadin et al., 2016). Speculatively, all four of the batholithic intrusive complexes that
define the eastern Sierra Crest (Sonora, John Muir, Whitney, in addition to the TIC) may represent the
centers of focused systems.
8. Conclusions
1) 102-85 Ma plutons in the CSN provide the opportunity to explore the first-order characteristics of
magma focusing in the upper crust (~6-10 km) and compare them to well-characterized examples
in long-lived volcanic fields. The CSN shows spatial-temporal patterns closely resembling those
seen in Andean and Cordilleran focused volcanic systems.
2) While the volcanic record is incomplete in the exhumed CSN arc section, widespread hypabyssal
bodies contemporaneous with deeper plutons and locally preserved volcanic packages, including
two caldera systems involved in magma focusing, provide the opportunity to explore the depth-
dependency of this phenomenon.
30
Figure 2.7: Schematic diagram of an arc crustal column, summarizing transcrustal and emplacement level models
for magma focusing. Volumes of known magmatic focusing systems are given for sense of scale. Brackets refer to
references in text. [1] YVF=Yanacocha Volcanic Field; Longo et al. (2010); [2] AVC=Aucanquilcha Volcanic
Cluster; Grunder et al. (2008); [3] SJML=San Juan magmatic locus; Lipman (2007); [4] APVC=Altiplano-Puna
Volcanic Complex; de Silva et al. (2006); [5] APMB=Altiplano Puna Magma Body; Ward et al. (2017); [6] Saleeby
(1990); [7] Walker et al. (2015); [8] Miller et al. (2009), describing upper crustal versus lower crustal magmatic
conditions. Porous-flow melt channels are depicted in white in the lower-crust level as proposed by Brown (2007).
Inset box at the upper-crust level summarizes magma-host rock interactions that likely drive focusing at the
emplacement level (e.g. Jellinek and DePaolo, 2003; Karlstrom et al., 2009; Cao et al., 2016). Vertical arrow
represents an increasing geothermal gradient resulting from protracted magma addition, enhancing upper-crustal
magma emplacement as described by de Silva and Gosnold (2007). The schematic cartoon to the right represents an
interpretive view of magmatic focusing in the CSN between 102-85 Ma, from 0-10 km depth. Colors of plutons are
loosely based on Figure 3 to show the age progression. Grey plutons and volcanoes represent inactive features of the
magma plumbing system.
31
3) During magma focusing in the CSN, larger magma chambers led to increased potential for larger-
scale magma processes such as fractionation and mixing. Emplacing large volumes of magma in
the upper crust increases the removal and some recycling of adjacent host rock/magma and thus
contributes to crustal re-working and differentiation. Thermal and mechanical mechanisms driving
focused magmatism likely operate throughout the crustal column, which requires additional testing
in the CSN, but models should include an upper-crustal component to drive initially unfocused
intrusive magmatism towards a highly organized configuration.
4) Eastward arc migration, crustal thickening and magma focusing in the CSN occurred at 1-4 km/myr
rates, increasing the likelihood that these processes created thermal and mechanical feedbacks
within the crustal column.
9. Acknowledgements
We acknowledge support from National Science Foundation grants EAR 1624847, EAR 1019636, and
EAR 1624854, a Geological Society of America Graduate Student Research Grant (2017), and the
University of Southern California (USC) Department of Earth Sciences Graduate Student Research Fund.
We thank researchers at the Arizona LaserChron Center for their support in obtaining U-Pb zircon ages,
and John Ayers for assistance with zircon separation at California State University, Fullerton and Yosemite
National Park Rangers for their support in the field. Snir Attia, Wenrong Cao, Sean Hartman, Barbara
Ratschbacher, and Sean Zalunardo are thanked for discussions and comments on the manuscript.
Discussions with Anita Grunder and Olivier Bachmann helped build a better understanding of magma
focusing in the volcanic record. We thank reviewers Shan de Silva and Calvin Miller for their thoughtful
and constructive comments that improved this paper, as well as An Yin for editorial handling.
32
Chapter 3: Spatially and temporally dynamic continental arc
activity in the Mesozoic central Sierra Nevada: Implications for arc
lithosphere evolution
This paper is in preparation for submission to Geochemistry, Geophysics, Geosystems:
Ardill, K.E., Paterson, S.R., Memeti, V., and Attia, S., (in prep). Spatially and temporally dynamic
continental arc activity in the Mesozoic central Sierra Nevada: Implications for arc lithosphere evolution.
Abstract
Although subduction is thought to be a continuous process, arc system behavior is known to be non-
steady-state, from magmatic, deformation, and climatic perspectives. This episodic and dynamic behavior
of arcs leads to uncertainty surrounding the spatial and temporal evolution of magma plumbing systems
from the mantle to the surface.
This study integrated field, geochronologic, and geochemical datasets from the Sierra Nevada arc
section to resolve the physical and chemical signals of arc behaviors across several spatial and temporal
scales, from arc-wide flare-ups, migration, and crustal thickening to regional magma focusing. These
spatiotemporal processes were considered within a Mesozoic-scale framework of spatial and temporal
patterns to explore the effects of dynamic processes on the lithosphere. Magmatic flare-ups each have
distinct bulk-rock isotopic signals, which during Cretaceous magmatism is closely coupled to eastward arc
migration, as magmas traversed through and incorporated different crust and mantle material. Additionally,
during this time, dramatic magmatic and tectonic thickening recorded by strain data (Cao et al., 2016) and
geochemical proxies doubled the thickness of the arc crust, and magmatism focused towards a central zone.
The restricted range of Sr and Nd whole-rock isotopes in the Late Cretaceous focusing zone, together with
field, geochemical, structural, and theoretical datasets, supports interpretations that magmas intruded at the
center of focusing, forming the Tuolumne Intrusive Complex, were interacting and mixing within large
(>1000 km
2
), long-lived (>0.5-1 m.y.), magma chambers.
These dynamic behaviors effectively primed the arc, which culminated in lower, mid, and upper-crustal
zones of magma melting (or mixing for upper-crustal magmas), assimilation, storage, and homogenization,
or MASH zones of Hildreth and Moorbath (1988). In comparison to volcanic systems, this priming process
may have also modified the physical and chemical connections between plutons and volcanoes (Lipman,
2007; de Silva et al., 2006). Non-steady-state behavior in the Sierra Nevada resulted in mountain-building
33
and voluminous continental crust formation by transforming the physical, thermal, and chemical properties
of the lithosphere over a relatively short duration (10’s m.y.).
1. Introduction
The importance of spatial versus temporal controls on the evolution of the lithosphere is a far-reaching
and unresolved question in arc systems, both oceanic and continental (e.g., Hildreth and Moorbath, 1988;
Bremond d’Ars et al., 1995; Cecil et al., 2018; Till et al., 2019). One aspect of this broad topic concerns
how spatial position in the arc (inboard vs. outboard), or the temporal evolution of the arc (immature vs.
mature), impact the types of sources, and their relative proportions, that generate compositionally diverse
arc magmas, whether that is depleted mantle, enriched mantle lithosphere, overlying crust, subducted
sediments, or fluids. In the Sierra Nevada arc section, the nature of the magma source region has been a
long-standing question, with early studies recognizing the significance of across-strike changes in arc
basement in controlling magmatic compositions (Moore, 1959; Kistler and Peterman, 1973, 1978),
followed by studies highlighting the importance of the depleted mantle component, and the enriched mantle
component, respectively (DePaolo, 1981; Coleman et al., 1992). With the recognition of arc tempos (e.g.,
episodic, or non-steady state magmatism, deformation, and climate), the question of spatial and temporal
magma evolution and estimates of magma sources, becomes increasingly complex (Armstrong and Ward,
1993; Bateman, 1992; Paterson and Ducea, 2015). The arc is no longer a ‘static’ reference frame, but is
highly dynamic, evolving in space and time.
Arc episodicity operates at multiple spatial and temporal scales, from arc-wide flare-ups and lulls,
migration, and crustal thickening or thinning, to regional magma focusing, and includes local magma
emplacement and eruption. The mechanisms and conditions that drive each of these processes are
controversial, and the extent to which these processes interact, or create feedback loops, is not well
understood (e.g., de Silva et al., 2006, 2015). Several studies have documented the transformative effect of
flare-ups, migration, focusing and crustal thickening on the arc lithosphere, including developing
voluminous batholiths that contain long-lived upper-crustal magma chambers and produce high-volume
eruptions, forming thick lithospheric roots, and potentially shutting off arc magmatism entirely (Lipman,
2007; de Silva and Gosnold, 2007; Grunder et al., 2008; Karlstrom et al., 2014; Cao et al., 2016, 2017;
Schwartz et al., 2017). As each of these processes occurred at some point in the Sierra Nevada arc history
(and sometimes repeatedly), this study will examine the geochemical record of dynamic, episodic
processes, the effects of these processes on the lithosphere, and whether they influence or modify our view
of spatial and temporal controls on arc magma sources.
34
Large structural, geochronologic, and geochemical datasets from the central Sierra Nevada (CSN)
demonstrate that the Mesozoic Sierra Nevada arc experienced three magmatic-tectonic flare-ups, with the
most voluminous flare-up occurring in the Cretaceous. During the Cretaceous flare-up the arc also migrated
eastward and magmatism focused, while the crust dramatically thickened, culminating in a ~70 km thick
crustal column. In contrast to previous findings using entire-arc datasets, each magmatic flare-up has a
different bulk-rock isotopic signal (e.g., DeCelles et al., 2009; Paterson and Ducea, 2015). Closely linked
spatial and temporal isotopic variation can be attributed to eastward arc migration, as magmas traversed
through and incorporated different crust and mantle material. This signal is also seen in trace-element ratios
that suggest an increasing role for amphibole and feldspar fractionation with eastward position (and time).
The restricted range of Sr and Nd whole-rock isotopes in the Late Cretaceous focusing zone, along with a
range of field evidence, mineral studies, and thermal models, suggests that magmas that localized at the
center of magma focusing (forming the Tuolumne Intrusive Complex) were not only voluminous, but also
hydrous, evolved, and readily able to mix.
The Cretaceous arc represents a sharp contrast to earlier arc magmatism in the Jurassic and Triassic.
Thus, we interpret that dynamic arc behaviors effectively primed the arc thermally and mechanically,
culminating in mid- to upper-crustal mixing, assimilation, storage, and homogenization zones, and large,
dynamic magma chambers. The implication is that this also potentially modified the physical and chemical
connections between upper-crustal reservoirs and volcanoes, which remains a target for future work. This
study finds that the mantle source contribution is dominant, the lithospheric mantle source is visible in
whole-rock datasets, and that there is a smaller proportion of compositional modulation by the crust. Non-
steady-state behavior in the Sierra Nevada resulted in mountain-building and voluminous (juvenile)
continental crust formation by transforming the physical, thermal, and chemical properties of the
lithosphere over a relatively short duration (10’s m.y.).
2. Background
2.1 Regional geology: lateral variations in Sierra Nevada pre-arc basement
Continental margin lithosphere was tectonically juxtaposed with accreted oceanic basement during the
mid-Late Paleozoic (Attia et al., 2017) (Fig. 3.1a-b). This formed a heterogeneous collage of displaced
continental lithosphere, accreted oceanic lithosphere, and arcs. These blocks had distinct lithologies and
isotopic characteristics (e.g., Kistler and Peterman, 1973). As a result, the Cordilleran magmatic arc
inherited significant along- and across-strike lithospheric variations (Fig. 3.1c). The earliest record of arc
magmatism is preserved in ca. 275 Ma plutons south of the Sierra Nevada, in the El Paso terrane (not
35
Figure 3.1: Paleozoic-Mesozoic evolution of the SW Cordilleran margin. (A) During the Neoproterozoic to Late
Paleozoic, western US strata formed in a passive margin environment, with thinning continental lithosphere
outboard, and transitioning to oceanic lithosphere. (B) Tectonic reshuffling of the passive margin led to truncation of
the margin, displacement of blocks, and formation of sub-vertical fault boundaries. (C) During the Mesozoic, from
ca. 250-85 Ma, arc magmatism in the Sierra Nevada stitched together crustal blocks with varying characteristics. (C)
Modified from Chapman et al. (2017).
36
shown), before arc magmatism migrated northwards into the Sierra Nevada, and southwards, towards the
Peninsular Ranges batholith (Cecil et al., 2019).
The pre-arc crustal materials in the Sierra Nevada form three broad groups, including: (1) allochthonous
Neoproterozoic- lower Paleozoic strata with interpreted western Laurentian sources; (2) Neoproterozoic-
Permian strata parautochthonous to southwest Laurentia; and (3) middle-upper Paleozoic deposits related
to the McCloud arc (Attia et al., 2017, and references therein). NW-SE striking, arc-parallel boundaries
between crustal blocks are steeply dipping in the present-day. From west to east these crustal-scale
boundaries include: the Foothills Suture; the Bench Canyon Shear Zone; and the Eastern Sierra Break
(Saleeby et al., 1986, 1989; Schweickert and Lahren, 1990; Stevens et al., 1997; Memeti et al., 2010) (Fig.
3.2). They bracket domains with distinct isotopic values and/or trends (e.g., Kistler and Peterman, 1973,
1978). From west to east, these domains are the Western Metamorphic Belt (domain 1); the Snow Lake
Block (domain 2); the Eastern Sierra Block (domain 3); and the Owens Valley Block (domain 4) (Fig. 3.2).
Lateral variations in basement composition occur from east to west across these domains (Doe and
Delevaux, 1973; Kistler and Peterman, 1973; DePaolo, 1981; Chen and Tilton, 1991; Ducea, 2001; Lackey
et al., 2005, 2008). Domains 1 and 3 have been interpreted to record a component of oceanic- to deep-
marine passive margin affinity basement (crust and mantle lithosphere), while domains 2 and 4 are
interpreted to record a component of shallow-marine passive margin affinity basement (e.g., Kistler and
Peterman, 1973, 1978; Kistler, 1990; DePaolo, 1981; Lackey et al., 2008; Memeti et al., 2010; Attia et al.
2017; Ardill et al., in press).
2.2 Dynamic arc processes
Studies are increasingly recognizing dynamic aspects of arc activity across a range of spatial and
temporal scales. These include: (1) arc-wide variations in magma addition rate, migration of the volcanic
front, and crustal thickening or thinning occurring at 10
2
-10
3
km scales, and 10
7
yr timescales (e.g.,
Karlstrom et al., 2014; de Silva et al., 2015; Chapman et al., 2017; Cao et al., 2017; Schwartz et al., 2017;
Cecil et al., 2018); (2) regional magma focusing phenomena at 10
1
-10
2
km scales, and 10
5
-10
7
yr timescales
(de Silva et al., 2006; Grunder et al., 2008; Lipman, 2007; Ardill et al., 2018); and (3) local intrusive-
complex scale dynamic processes at 10
-4
-10
2
km scales, and <10
6
yr timescales (e.g., Paterson et al., 2016;
Bergantz et al., 2017). Below we describe arc- to regional-scale behaviors that are the focus of this study
(groups (1) and (2) above):
Episodic magma addition to the arc alternates between periods of high magma addition rates (arc flare-
ups) and periods of low magma addition rates (lulls). Volume calculations using maps, and zircon ages with
time in plutonic, volcanic and sedimentary rocks, all record this pattern (Bateman, 1992; Ducea, 2001;
37
Figure 3.2: Central Sierra Nevada geologic map. Key crustal boundaries are labelled (Foothills Suture; Bench Canyon Shear Zone),
along with the names of different domains (Western Metamorphic Belt; Snow Lake Block; Eastern Sierra Block). The locations of
known geographic markers and pendants are included for reference. Triassic, Jurassic and Cretaceous plutons and Cretaceous
volcanic rocks are highlighted and discussed in the text. Drafted by S.Attia.
38
Paterson and Ducea, 2015). The cause(s) of arc flare-ups, including the transition between flare-up and lull
states, remains unresolved. The main competing models invoke either external, lower-plate episodic
processes, controlled by subducting plate parameters, or internal, upper-plate episodic processes, driven by
mantle wedge or overlying crustal column feedbacks (e.g., Paterson and Ducea, 2015; Cecil et al., 2018;
Martinez-Ardila et al., 2019). Kirsch et al. (2016) did not find any statistical connection between lower
plate parameters (e.g., slab age, subduction rate) and arc flare-ups in Cordilleran arc systems (see also
Ducea and Barton, 2010). In contrast, in the Fiordland arc section, ridge subduction and/or slab tear were
proposed as possible triggers of episodic mantle melting (Decker et al., 2017; Schwartz et al., 2017),
illustrating the importance of considering models on a case-by-case basis. Upper-plate episodicity is further
divided into mantle-driven flare-up models (e.g., Martinez-Ardila et al., 2019), or crustal-driven flare-up
models (Ducea and Barton, 2007; DeCelles et al., 2009).
During arc migration, the magmatic front sweeps inboard towards the continental interior, or retreats
towards the trench; sometimes this occurs repeatedly during an arc’s lifespan (e.g.; Kay et al., 2005; Morton
et al., 2014). The causes of arc migration are unclear, and a range of models have been proposed to explain
this behavior. One popular model involves changes in the dip of the subducting slab (e.g., Coney and
Reynolds, 1977). Additional models propose migration of the trench by subduction erosion or accretion
(e.g., Clift et al., 2005), or migration of the mantle-melting zone by crustal thickening of the upper plate
(Chin et al., 2012; Karlstrom et al., 2014). Numerical models highlight the importance of relative coupling
between the subducting and overriding plates in controlling arc migration directions and rates (e.g., O’
Driscoll et al., 2009; Holt et al., 2015a), as well as the importance of mantle wedge and upper mantle
properties and dynamics (Skinner and Clayton, 2011; Schellart et al., 2017).
Regional-scale magma focusing is the spatial localization of magma from scattered, broad regions
towards a central location, where magmas then amalgamate to form larger bodies (either intrusive or
volcanic complexes). Magma focusing is, in part, an emplacement-level process, as both unfocused and
focused magmatism is present in the CSN and in volcanic settings (e.g., Southern Rocky Mountain Volcanic
Field; Lipman, 2007). Possible emplacement-level mechanisms for magma focusing include: (1) ascent
through previously established thermal pathways (Marsh, 1982); (2) thermal/ductile weakening of host-
rocks promoting growth of magma bodies over eruption (Jellinek and DePaolo, 2003); and (3) the growth
of pluton stress-fields enabling larger plutons to capture and focus rising magmas (e.g., Karlstrom et al.,
2009). The occurrence of magma focusing in both plutonic and volcanic environments demonstrates that it
is a process operating at multiple crustal levels, and potentially throughout the crustal column, where several
other processes (e.g., thermal gradients) may influence magma focusing (e.g., de Silva and Gosnold, 2007;
Ardill et al., 2018).
39
Crustal thickening extends the length of magma pathways from the source to the surface (Karlstrom et
al., 2014). As such, it can produce a petrologic signal (Profeta et al., 2015). Both tectonic shortening and
magmatic addition contribute to crustal thickening, thus bridging the magmatic arc behaviors described
above with deformation processes (Cao et al., 2016).
2.2.1 Dynamic arc processes in the Sierra Nevada
Arc flare-ups: The Sierra Nevada arc experienced three magmatic-tectonic flare-ups, each lasting 20-30
m.y. in the Triassic, Jurassic, and Cretaceous (Armstrong and Ward, 1993; Ducea, 2001; DeCelles et al.,
2009; Paterson and Ducea, 2015; Cao et al., 2015). The flare-ups occurred between ~230-205 Ma, ~170-
150 Ma, and 120-85 Ma, with a ~5 m.y. uncertainty on the start and end dates, and are expressed in plutonic,
volcanic, and sedimentary records (Attia et al. 2020) (Fig. 3.3a). During each flare-up, magma addition
rates increased 100-1,000 times over lull periods (Paterson and Ducea, 2015). For the Cretaceous arc,
magma addition rates are estimated at 0.85 km
3
/km
2
/Ma, approximately 2-5 times more voluminous than
the Jurassic and Triassic flare-ups (Ratschbacher et al., 2019). The mantle contribution is estimated between
0.56-0.82 km
3
/km
2
/Ma during this period (Ratschbacher et al., 2019). Previous work in the Sierra Nevada
has interpreted εNd i ‘pull downs’ during flare-ups as a record of recycling lithospheric material (Ducea and
Barton, 2007; DeCelles et al., 2009). In contrast, Attia et al. (2020) have shown that each Sierran flare-up
is dominated by juvenile mantle input (70-90%). The implication of high magma addition rates during flare-
ups is that they can form larger plutons (and larger magma chambers) and lead to larger volcanic eruptions
(see also de Silva et al., 2015). The Cretaceous arc flare-up also coincided with arc migration, magma
focusing, and crustal thickening, described below.
Arc migration: The Cretaceous arc migrated from west to east at rates between ~2.6-4.8 km/Myr (Chen and
Moore, 1982; Cecil et al., 2012). In the central Sierra Nevada (CSN), the migration rate calculated from
comparing sample age and eastward position is 2.7 km/Myr (Ardill et al., 2018). The implication of arc
migration in the CSN is that the locus of magmatism passed through laterally variable pre-arc lithospheric
basement (section 2.1). Cretaceous arc migration has been associated with decreasing εNd i and increasing
87
Sr/
86
Sr i with decreasing age and with eastward position (Kistler, 1990; Ardill et al., 2018). Chapman and
Ducea (2019) associated this arc migration signal with increased recycling of hydrous mantle lithosphere
into arc magmas towards the continental interior.
Magma focusing: Between ca. 102 Ma and 85 Ma, upper-crustal magmatism in the central Sierra Nevada
spatially focused (Ardill et al., 2018). Magmatism transitioned from the emplacement of several disparate
plutons towards the emplacement of a central, nested intrusion, the Tuolumne Intrusive Complex (TIC).
The TIC has previously been associated with magma focusing phenomena (Grunder et al., 2008). During
40
Figure 3.3: Central Sierra Nevada geologic map and sample summary. (A) Histogram showing kernel density
estimates for intrusive rock ages, volcanic rock ages, and Mesozoic detrital zircon ages from the CSN. Hatched
areas indicate arc flare-up periods, with the change in pattern representing uncertainty on the timing of each flare-up.
Adapted from Figure 1 of Attia et al. (2020). (B) Map showing the locations of samples compiled in this study.
Drafted by S. Attia.
41
this time, magma volumes emplaced increased, and magmas formed hydrous, hybrid compositions. The
implication of upper-crustal magma focusing in the CSN is that it promoted the formation of large, long-
lived magma chambers by localizing the magma supply (Ardill et al., 2018). Magma focusing thus has
consequences for aiding magma mixing and driving meso-scale homogenization processes, the potential
for large volcanic eruptions of ‘monotonous intermediates’, and also for the formation of ore deposits at
hypabyssal and volcanic levels (Lipman, 2007; Grunder et al., 2008; Longo et al., 2010). In the CSN,
magma focusing has been associated with a restriction in the range of
87
Sr/
86
Sr i and εNd i isotopes, and
increased dispersion in some trace element indicators, such as Rb/Sr (Ardill et al., 2018).
Crustal thickening and thinning: Initial estimates on crustal thickness in the Sierra Nevada come from field
relationships. For example, a relatively thin arc crust in the Jurassic is supported by field observations of
marine volcanic and sedimentary packages, compared to largely sub-aerial deposits in the Early Triassic
and Cretaceous arcs. Secondly, Late Cretaceous, ~100 Ma plutons in the southern Sierra Nevada extend to
approximately 40 km depth (Al in hornblende barometry; Pickett and Saleeby, 1993), below which a ~30-
80 km thick lower crustal and mantle lithosphere root is interpreted from xenolith barometry (Saleeby et
al., 2003; Chin et al., 2012). Additional, quantitative estimates of Mesozoic crustal thickness in the Sierra
Nevada are made using strain measurements, isostatic models, and whole-rock geochemical proxies. Strain
measurements by Cao et al. (2016) estimated 65% crustal shortening during the Mesozoic, which, together
with magmatic addition, resulted in a ~80 km thick crust and a ~30 km thick lithospheric root by the Late
Cretaceous. Cao and Paterson (2016) found similar results from a 1D mass-balance and isostasy model.
Profeta et al. (2015) calibrated Sr/Y and La/Yb bulk-rock ratios against modern arc crustal thicknesses to
estimate paleo-arc thicknesses in the Sierra Nevada, with results ranging from ~35 km in the Triassic and
Jurassic to >70 km by the Late Cretaceous (depth to petrologic moho). The implications of dramatic crustal
thickening in the Cretaceous Sierra Nevada are that it generated a high elevation mountain range, promoted
high erosion rates, and could have contributed to arc migration or even arc cessation (Chin et al., 2012;
Karlstrom et al., 2014; Cao et al., 2016).
2.3 Links between dynamic processes and arc maturity
Some existing criteria in the literature for evaluating the thermal, mechanical, and chemical maturity of
an arc include: (1) high geothermal gradients, long-lived, transcrustal magma plumbing systems, larger
volumes of upper crustal magmatism (de Silva and Gosnold, 2007; Karakas et al., 2017; 2019); (2)
evolving, or upward-migrating, brittle-ductile transition zone; increasing the ratio of magmas emplaced vs.
erupted (Jellinek and DePaolo, 2003; Karlstrom et al., 2009); (3) forming a thick crust and a thick, high-
density lithospheric root (e.g., Saleeby et al., 2003); and (4) magma processing in one, or several, MASH
(melting, assimilation, storage, and homogenization) zones (Hildreth and Moorbath, 1988; de Silva et al.,
42
2006; Walker et al., 2015). Below we explore how these arc attributes (and thus arc maturity in general)
are currently thought to be influenced by dynamic arc processes:
Arc flare-ups: de Silva and Gosnold (2007) proposed links between arc systems in ignimbrite flare-up
periods and thermal-mechanical maturation of the arc crustal column. In summary, increased mantle-
derived magma addition rates provides the heat to modify the geothermal gradient (and brittle-ductile
transition), facilitating magma ascent and accumulation at increasingly shallower levels. The link between
magma addition rates and geothermal gradients in the crust was subsequently explored by Karakas et al.
(2017, 2019) in thermal models of the Ivrea crustal section, which demonstrated thermal maturity of the
system (achieved after 4 m.y. of magmatism) promoted transcrustal magmatism and upper-crustal magma
chambers lasting 10
5
-10
6
yr. This study stressed the combination of the rate and duration of magma addition
was crucial to achieving thermal maturity. Variations in crust-mantle interaction, tracked by εHf zrc ratios
were interpreted to be closely associated with the thermal signature of this crustal section (Storck et al.,
2020).
Arc migration: As the arc migrates through different crust and mantle lithosphere, it may encounter more
(or less) melt-fertile materials. The heterogeneity in source materials can lead to evolved magma
compositions but is highly dependent on the nature of the source (e.g., Chapman et al., 2017). Chapman
and Ducea (2019) explored this idea in the Sierra Nevada arc to explain evolved εNd i ratios during arc
migration. It is linked to the observation made by Chin et al. (2014), that ~100 Ma mantle-wedge refertilized
xenoliths from <90-105 km depth were one signal of a mature arc system with thick crust. If arc migration
is linked to a process like crustal thickening, or due to slab flattening, then it can also be expected to have
thermal as well as chemical effects on the overlying magma plumbing system (e.g., Dumitru et al., 1991).
Magma focusing: The magma focusing process may facilitate both the formation and organization of
discrete magma reservoirs at multiple crustal levels (Miller et al., 2009; Ardill et al., 2018). Furthermore,
the sustained addition of magma in a localized zone could generate ‘thermal weakening’ effects of the
surrounding wall rocks, allow magma reservoirs to grow, and drive the brittle-ductile transition to shallower
levels (Jellenik and DePaolo, 2003). Focusing examples all occur during flare-ups, which is permissive that
the two processes may be linked (e.g., Grunder et al., 2008; de Silva et al., 2006; Ardill et al., 2018).
Crustal thickening: Crustal thickening by tectonic and magmatic addition is known to generate a thick,
high-density, lithospheric mantle root (Saleeby et al., 2003; Cao et al., 2016). This may refrigerate the arc
crust; however, this relative cooling effect is likely countered by the (much larger) heating effects associated
with mantle-derived magma addition and radioactive decay in the continental crust column. A thick crust
and mantle lithosphere provides a longer pathway for magmas to differentiate towards evolved
43
compositions and provides a renewed supply of crustal materials to be incorporated into MASH zones via
downward flow (Paterson and Farris, 2008; Cao et al., 2016).
Continental arcs are potentially more likely to reach these ‘prime’ conditions than island arcs, due to
their longer lifespans, and the involvement of multiple magma sources, including pre-existing crustal
materials (Ducea et al., 2015) (see also Ribeiro et al., 2020 for Mariana island arc maturation example for
comparison). In summary, the consequences of forming a mature magma plumbing system are that magmas
can be efficiently processed in the crust and mantle. Thermal and mechanical feedbacks can promote ascent
and accumulation of greater volumes of evolved magma at higher crustal levels, and culminate in the
formation of large, long-lived magma reservoirs and volcanic eruptions (de Silva and Gosnold, 2007;
Paterson et al., 2011; Lipman and Bachmann, 2015).
To explore the role of dynamic (i.e., temporally and spatially variable) processes in the Sierra Nevada
arc, spatial patterns are examined initially, in particular, patterns that persist for the duration of arc
magmatism. Then, patterns that are temporally controlled, and those that relate to dynamic processes, are
compared and summarized.
3. Methods
3.1 Data filtering, binning, and statistics
Plutonic, volcanic, and hypabyssal samples were filtered by age and location. Samples between 250-80
Ma were included, located between latitudes 36.85° and 38.4°, and longitudes -121° and 118°, across a
~120 km arc-perpendicular transect (Fig. 3.3b, 3.4a). Xenoliths and sedimentary rocks were excluded.
Geochemical data were binned into 10 m.y. age intervals, in order to account for errors in the sample age
estimates on the order of 0.5-5 m.y. (method dependent). Outliers in the data were defined as points
extending beyond the whiskers on boxplots (1
st
or 3
rd
quartile, +/- IQR*1.5). Outliers are included on the
plot, although some are excluded from the plots if they extended beyond reasonable y-axes values (details
listed in Figure captions). A cubic spline interpolation was applied to the dataset, with a lambda
(smoothness indicator) of 0.05. The cubic spline function interpolates between areas of high and low data
density (e.g., flare-ups and lulls). Samples were divided into two groups by SiO 2 content, determined by
the median SiO 2 value of 65 wt.%. Samples <65 wt.% SiO 2 are informally termed ‘mafic-intermediate’ and
samples >65 wt.% SiO 2 are referred to as ‘intermediate-felsic’.
Data sources from the compilation include: Ague and Brimhall (1988); Alasino et al. (2019); Ardill et
al. (2018; in press); Attia (2020); Attia et al. (2020); Barbarin et al. (1989); Barth et al. (2011, 2012);
Bateman and Chappell, (1979); Bateman et al. (1988); Bateman, (1992); Burgess and Miller (2008); Cao
44
Figure 3.4: Summary of sample attributes. (A) Plot showing location vs. binned sample age for all samples in the
compilation. Locations (plotted in the NAD27 long-lat co-ordinate system) have been reprojected into an arc-
perpendicular orientation, to examine across-arc variation. Sample age is binned into 10 m.y. intervals to account for
wide variation in the uncertainty of rock ages. Flare-up periods shown in Figure 3a are highlighted to show that the
concentration of samples is coincident with flare-ups. (B) Histogram of the number of samples vs. binned age (10
m.y. interval), divided into plutonic (pink, n=1284), hypabyssal (orange, n=114), and volcanic (green, n=405)
categories. Curves represent probability density estimates. Grey fields represent flare-up periods, as shown in Figure
3a. There are no Jurassic hypabyssal samples in the compilation, and two Triassic samples. (C) Histogram of the
number of samples vs. SiO 2 contents (2.5 wt.% SiO 2 intervals). The median value is shown with a grey dashed line.
Colors as in Figure 4b. Plutonic (n=1137), hypabyssal (n=114), and volcanic (n=389).
45
(2015); Chen and Tilton (1991); Coleman et al. (2012); DePaolo (1981); Economos et al. (2008); Frost
(1987); Gray et al. (2008); Greene (1995); Kistler and Peterman (1973); Kistler and Fleck, (1994); Kistler
et al. (1986); Lackey et al. (2005, 2008, 2012); Loetterle (2004); Lowe, (1995); Macais (1996); Masi et al.
(1981); Memeti (2009); Paterson et al. (2008); Peck and Van Kooten (1983); Putirka et al. (2014): Ratajeski
et al. (2001); Reid et al. (1993); Scheland (2019); Snow (2007); Solgadi and Sawyer (2008); Stanback
(2018); Tobisch et al. (1991); Truschel (1996); Williams (2018); Wooden et al. (1999); Žák et al. (2009);
and unpublished data.
The thermochronology compilation includes
40
Ar/
39
Ar and K-Ar biotite and hornblende ages, collected
from: Cao et al. (2015); Dodge and Calk (1986); Evernden and Kistler (1970); Hillhouse and Gromme
(2011); Kistler and Dodge (1966); Matzel et al. (2005, 2006); McNulty (1995); Memeti et al. (2010); Nadin
et al. (2016); Robinson and Kistler (1986); Sharp et al. (2000); and unpublished data.
3.2 Location of west-east geologic boundaries (domains 1-4)
Major geologic boundaries were located on geologic maps (Huber et al., 1989; Wagner et al., 1991;
Bateman, 1988) and studies (Saleeby et al., 1986, 1989; Schweickert and Lahren 1990; Stevens et al., 1997;
Memeti et al., 2010). The location of these boundaries in the CSN was then projected onto an arc-
perpendicular transect. These boundaries divide domains 1-4 in subsequent figures (Fig. 3.2b). From west
to east these boundaries include the Foothills Suture (arc perpendicular easting -119.9); the Bench Canyon
Shear Zone (arc perpendicular easting -119.37) and the Eastern Sierra Break (arc perpendicular easting -
119.14).
4. Results
4.1. Regional compilation overview
The regional compilation consists of 1816 samples with bulk-rock element data; a subset of these
samples (n= 139-446) also include bulk-rock isotopic data (Rb-Sr, Sm-Nd, and U-Pb systems) (Fig. 3.3b).
Multiple crustal levels are sampled in the compilation, from surface metavolcanic deposits (herein volcanic
samples) to ~10 km depth plutons. Plutonic samples make up the largest population of the dataset (71%),
with a smaller group of volcanic samples (22%) and a minor population of hypabyssal samples (6%).
In Figure 3.4a, where sample age is plotted against west-east position, the Triassic and Jurassic arcs are
laterally extensive, with wide spatial distribution in each age bin. In contrast, the Cretaceous arc shows the
eastward younging pattern. The sample-age distribution of samples in the compilation mimics the arc flare-
up pattern defined by zircon ages (Figs. 3.3a, 3.4a-b). This pattern is evident at both plutonic and volcanic
levels and it may also be true at the hypabyssal level; however hypabyssal data is extremely limited for
46
Jurassic and Triassic flare-ups. As expected, based on present-day exposures, the Cretaceous record is the
most well-expressed flare-up in the compilation (Fig. 3.4b).
Geochemically, the compilation is dominated by felsic compositions. The median SiO 2 contents for all
samples is 65 wt.% (Fig. 3.4c). The sample distribution for plutonic rocks is strongly skewed towards
granodiorite and granite compositions. Hypabyssal rocks are skewed towards felsic compositions and share
a similar SiO 2 range to the most felsic volcanic rocks (>70 wt.% SiO 2). Volcanic rocks are weakly bimodal,
with broad peaks at ~55 wt.% SiO 2 and ~72 wt.% SiO 2 (Fig. 3.4c).
Some trace element ratios examined in section 4.2 below show trends with SiO 2. These are summarized
in Figure 3.5. In summary, Rb/Sr, La/Yb, and Th/La have higher median values in intermediate-felsic
samples containing >65 wt.% SiO 2 than in mafic-intermediate samples (<65 wt.% SiO 2) as well as show a
greater range in values (Figs. 3.5a, d, i). This likely reflects the incompatibility of Rb, La, and Th in the
melt during differentiation. In contrast, Dy/Yb, and by extension Dy/Dy* [Dy N/La N
4/13
*Yb N
9/13
], both
decrease from intermediate-mafic samples to samples with higher SiO 2 contents (Figs. 3.5c, f). This appears
to reflect the compatibility of Dy during differentiation, possibly from amphibole fractionation (Davidson
et al., 2007, 2012). Sr/Y, Zr/Hf, Ba/La and Eu/Eu* [Eu N/√Sm N*Gd N] show no significant difference
between mafic-intermediate samples and intermediate-felsic samples (Figs. 3.5b, e, g, h).
As with SiO 2, there are variations in the trace element ratios with respect to crustal level sampled
(plutonic, hypabyssal, or volcanic). For example, median values of Rb/Sr and Ba/La increase from plutonic
to volcanic levels, in both mafic-intermediate and intermediate-felsic groups (Figs. 3.5a, h). This is also
apparent for Zr/Hf but there is less variation in the median here (Fig. 3.5b). Increasing values could reflect
upper-crustal differentiation and ascent of incompatible element-enriched melts to shallower levels. Median
values of Dy/Yb, La/Yb, and Sr/Y decrease from plutonic to volcanic levels (Figs. 3.5c, d, e). This could
reflect a signal of amphibole/garnet fractionation, feldspar fractionation, or accumulation of these minerals
in upper-crustal plutons. Th/La also decreases from plutonic to volcanic levels (Fig. 3.5i). The smaller
dataset for hypabyssal samples generally falls between the plutonic and volcanic median, although there
are several exceptions, shown in Figure 3.5.
4.2 Spatial geochemical patterns
Compositional variations from west to east across the arc are considered for the Triassic, Jurassic, and
Cretaceous arcs, in order to: (1) explore the significance of pre-arc basement contributions through time;
and (2) test which arc attributes persist throughout its lifetime.
47
Figure 3.5: Trace and Rare Earth element analysis of the compilation, by SiO 2 group and by crustal level (plutonic,
hypabyssal, volcanic). Rare Earth elements are normalized to Anders and Grevesse (1989). Tukey boxplots are
described in section 3.1. Violin plots are included to show the distribution of the data. Outliers are shown as points.
Filled symbols = mafic-intermediate group; Open symbols = intermediate felsic group. Symbols as in Figure 3b. (A)
Rb/Sr; (B) Zr/Hf; (C) Dy/Yb; (D) La/Yb; (E) Sr/Y; (F) Dy/Dy* (Davidson et al., 2012).
48
Figure 3.5 (cont.) Trace and Rare Earth element analysis of the compilation, by SiO 2 group and by crustal level
(plutonic, hypabyssal, volcanic). (G) Eu/Eu*; (H) Ba/La; (I) Th/La.
49
4.2.1. Major, minor, and trace elements
West-east variations in major element compositions were first documented by Moore (1959) and
supported by several subsequent studies of Mesozoic Sierra Nevada rocks (Chen and Moore, 1982;
Bateman, 1992; Ardill et al. 2018). From west to east, SiO 2 is invariant, K 2O increases, and Al 2O 3 and CaO
decrease, all attributed to the increase in alkali feldspar: plagioclase ratio (Moore 1959; Bateman, 1992).
Other major elements lack discernable west-east trends (Ardill et al., 2018).
Trace element indicators of differentiation show variable west-east patterns (Fig. 3.6a-e). In the case of
Rb/Sr, the spread of values is minimal in Jurassic and Cretaceous rocks from the western metamorphic belt
(domain 1), compared to eastward sections for all Mesozoic rocks (Fig. 3.6a). Ba/Sr (not shown) depicts
the same relationship, except that it sharply decreases at the boundary of domains 3 and 4. Additional
patterns, from west to east include: (1) Zr/Hf decreasing in Cretaceous samples, but this trend is not
mirrored in Jurassic and Triassic samples, which are largely invariant (Fig 3.6b); (2) Dy/Yb decreasing in
each of the arc datasets, although the Cretaceous arc shows the most pronounced change (Fig. 3.6c); (3)
La/Yb increasing from west to east in each of the arc datasets, with Cretaceous samples showing the largest
range (Fig. 3.6d); (4) Sr/Y increasing, matching the pattern of La/Yb (Fig. 3.6e). Sr/Y samples are filtered
by SiO
2 and MgO according to Profeta et al. (2015), in order to use their calibration for crustal thickness
(subsequent section). Both Eu/Eu* and Dy/Dy* decrease from west to east (Figs. 3.7a-b). This pattern is
observed in Triassic, Jurassic and Cretaceous arc datasets.
In oceanic arc basalts, the ratio of Ba/La has been interpreted to represent the contribution of aqueous
fluids to arc magmas, and Th/La has been interpreted to trace sedimentary inputs (Johnson and Plank, 1999;
Plank, 2005; Brounce et al., 2014). If fractional crystallization is the dominant process generating
compositional diversity in magmas, then this ratio may be extrapolated to apply to felsic granitoid rocks.
However, open-system processes such as magma recharge and mixing, and assimilation in the crustal
column will modify Ba/La and Th/La ratios. Thus these ratios will not record a lower-plate source signal.
Additionally, in upper-crustal plutons, Ba may behave compatibly, during crystallization of K-feldspar
and/or biotite, and Th during zircon and/or allanite crystallization (GERM database). West-east variations
in these ratios show that where Ba/La is high, Th/La is low, and vice versa, as Th is likely behaving
incompatibly (Figs. 3.8a-b). Notably, the inflection points of the spline function on these west-east trends
lines up well with both isotopic and geologic breaks (see below; Figs. 3.6-3.9). Furthermore, the west-east
patterns of Ba/La and Th/La match the west-east patterns defined by Pb and Sr isotopes, indicating that
they may be coupled and tracing the same magma source component (Figs. 3.8-3.9, see below).
50
Figure 3.6: Trace element ratios plotted against re-projected west-east location. Domains are labelled 1-4
corresponding to section 2.1 and Figure 3.2. Shading highlights out-of-sequence sections that have been tectonically
displaced. Colors and symbols as in Figure 3.3b. Dashed line represents cubic spline interpolation, summarized in
section 3.1. (A) Rb/Sr; (B) Zr/Hf; (C) Dy/Yb; (D) La/Yb; (E) Sr/Y, filtered by SiO 2 (55-68 wt.%) and MgO (<4
wt.%), following methods outlined in Profeta et al. (2015).
51
Figure 3.7: Dysprosium and Europium anomalies plotted against re-projected west-east location. Refer to Figure 3.6
for explanation of colors, symbols and lines. (A) Eu/Eu*; (B) Dy/Dy*.
52
Figure 3.8: Trace element ratios Ba/La and Th/La plotted against re-projected west-east location. Refer to Figure
3.6 for explanation of colors, symbols and lines. (A) Ba/La; (B) Th/La.
53
Figure 3.9: Initial Sr (
87
Sr/
86
Sr i), Nd (
143
Nd/
144
Nd i converted to εNd i), and Pb (
206
Pb/
204
Pb i,
207
Pb/
204
Pb i,
208
Pb/
204
Pb i)
isotopic ratios plotted against re-projected west-east location. Refer to Figure 3.6 for explanation of colors, symbols
and lines. On the left side of each plot, constraints on magma sources are plotted. Arrows indicate 100 Ma xenolith
compositions from the Sierra Nevada from Mukhopadhyay and Manton (1994); Ducea (1998) and Ducea and
Saleeby (1998). Stars represent the isotopic compositions of lower crustal rocks from the Southern Sierra Nevada
(Pickett and Saleeby, 1993). Patterned fields represent possible Pb crust and mantle reservoirs from Zartman (1974)
and Chen and Tilton (1991). Note the significant overlap where patterns intersect (e.g., cross-hatched pattern). (A)
87
Sr/
86
Sr i; (B) εNd i; (C)
206
Pb/
204
Pb i; (D)
207
Pb/
204
Pb i; (E)
208
Pb/
204
Pb i.
54
4.2.2 Isotopes
Bulk-rock isotopic signals of plutonic, hypabyssal, and volcanic rocks vary from west to east across the
arc section (Fig. 3.9). This is a well-known spatial pattern that has been attributed to lateral variations in
arc basement rocks from oceanic to continental lithosphere types (e.g., DePaolo, 1981; Chen and Tilton,
1991; Kistler, 1990; Lackey et al., 2008) (Fig 3.1).
This pattern is well-defined for
87
Sr/
86
Sr i and εNd i ratios of plutonic, hypabyssal and volcanic rocks;
Juvenile isotope signals are found in the Western Metamorphic Belt, and evolved isotopic compositions
occur in the Owens Valley, although it is not a uniform, or linear, transition (Figs. 3.9a-b).
206
Pb/
204
Pb,
207
Pb/
204
Pb, and
208
Pb/
204
Pb oscillate between high and low values from west to east (Figs. 3.9c-e).
Comparison of this dataset to Cretaceous lower crust and mantle xenoliths by Mukhopadhyay and Manton
(1994) and Ducea and Saleeby (1998) shows that, although sampling of xenoliths is spatially unconstrained
in terms of basement type, the values of samples from the lower crust and mantle lithosphere span almost
the entire range of values we see in the upper-crustal section (Ducea, 2001; Fig. 3.9). Furthermore, there is
no observed difference between the isotopic signal of volcanic, hypabyssal or plutonic rocks in a given
domain (Fig. 3.9).
Inflections in the slope of the median values of each isotopic system (best fit lines in Fig. 3.9) correspond
to the prominent structural features described in the background section (Saleeby et al., 1986; Kistler, 1990).
Domains 1 and 3 are characterized by low
87
Sr/
86
Sr i and high εNd i, as well as increasing Pb isotope ratios.
In contrast, domains 2 and 4 are characterized by high
87
Sr/
86
Sr i and lower εNd i and decreasing Pb isotope
ratios. Due to large overlap in the composition of Pb isotopic reservoirs defined by Zartman (1974), it is
not possible to uniquely resolve the contribution of passive margin sediments to the magmatic rocks at the
arc scale (Fig. 3.9c-e).Our compilation further defines the location and extent of the pre-arc framework
rocks and supports previous work that this spatial pattern persists throughout the arc’s lifetime, from the
Triassic to the Cretaceous in domains 1, 3, and 4. Domain 2 contains only Cretaceous magmatic rocks;
therefore it is not possible to evaluate this model there.
4.3 Temporal geochemical patterns
Due to the episodic nature of flare-ups and lulls, the data are concentrated during flare-up periods (Figs.
3.4, 3.10-3.13). The cubic spline function (red dashed line) interpolates between areas of low data density
(lulls). Trends for each trace-element or isotope ratio are compared across two timescales: (1) Mesozoic-
scale, from arc initiation to cessation; and (2) Intra-flare-up scale, from start of flare-ups to end of flare-
ups, defined by U-Pb zircon age distribution in Figure 3.3b (Attia et al., 2020).
55
Figure 3.10: Trace element ratios plotted against sample age, binned in 10 m.y. intervals. Grey shaded areas show
flare-up periods as documented by Attia et al. (2020) (Fig. 3.3a). Red dashed line shows the cubic spline
interpolation. Filled symbols have SiO 2 <65 wt.%, and open symbols have SiO 2 >65 wt.%. Symbol key as in Figures
3.3 and 3.9. (A) Rb/Sr; (B) Zr/Hf; (C) Dy/Yb; (D) La/Yb; (E) Sr/Y, filtered by SiO 2 (55-68 wt.%) and MgO (<4
wt.%), following methods outlined in Profeta et al. (2015).
56
Figure 3.11: Dysprosium and Europium anomalies plotted against sample age. Refer to Figure 3.10 for explanation
of colors, symbols and lines. (A) Eu/Eu*; (B) Dy/Dy*.
57
Figure 3.12: Trace element ratios Ba/La and Th/La plotted against sample age. Refer to Figure 3.10 for explanation
of colors, symbols and lines. (A) Ba/La; (B) Th/La.
58
Figure 3.13: Initial Sr (
87
Sr/
86
Sr i), Nd (
143
Nd/
144
Nd i converted to εNd i), and Pb (
206
Pb/
204
Pb i,
207
Pb/
204
Pb i,
208
Pb/
204
Pb i) isotopic ratios plotted against sample age. Refer to Figure 3.10 for explanation of symbols and lines.
Samples are color-coded by spatial location in the arc, grouped into domains 1-4, see legend. (A)
87
Sr/
86
Sr i; (B)
εNd i; (C)
206
Pb/
204
Pb i; (D)
207
Pb/
204
Pb i; (E)
208
Pb/
204
Pb i.
59
4.3.1 Major, minor, and trace elements
Major elements lack discernable trends through time, at both the Mesozoic-scale and intra-flare-up scale.
Rb/Sr lacks a Mesozoic-scale trend, instead showing variations within individual flare-up periods. Median
Rb/Sr (boxplots in Fig. 3.10a) initially increases during the Triassic and Cretaceous flare-ups and then
decreases towards the later stages of flare-up magmatism (Fig. 3.10a). In contrast, the Jurassic flare-up
records only decreasing Rb/Sr median values. The range in Rb/Sr also decreases in the final ~10 Myr of
each flare-up. This trend is also observed in Ba/Sr (not shown). Zr/Hf from the Triassic to Cretaceous
remains approximately invariant, however, the trends within each flare-up are distinct (Fig. 3.10b). Zr/Hf
increases during the Triassic flare-up, increases then decreases during the Jurassic flare-up, and decreases
during the Cretaceous flare-up. Dy/Yb shows a Mesozoic-scale trend, increasing consistently from the
Triassic through to the Cretaceous (Fig. 3.10c), and intra-flare-up trends appear to be secondary compared
to the overall increasing trend. La/Yb and Sr/Y during the Triassic and Jurassic flare-ups remains at low
values until ~100 Ma when they both dramatically increase (Figs. 3.10d-e). Higher Sr/Y or La/Yb values
indicates a greater depth of magma generation, and this has been leveraged to estimate paleo-crustal-
thicknesses in arcs (Chapman et al., 2015). Using the calibration of Profeta et al. (2015), this corresponds
to an arc crustal thickness of ~30-40 km (Fig. 3.10e). La/Yb and Sr/Y increase during the Cretaceous flare-
up, to corresponding thicknesses of >70 km. Dy/Dy* does not show a Mesozoic-scale trend. It appears to
decrease during the Triassic flare-up, but the 200-210 Ma bin contains unusually high values (Fig. 3.11a).
The Jurassic flare-up is largely invariant, and Dy/Dy* then decreases during the Cretaceous flare-up.
Eu/Eu* also lacks a Mesozoic-scale trend and may weakly decrease during flare-up periods (Fig. 3.11b).
Ba/La decreases overall from the Triassic through to the Cretaceous, each flare-up period records
decreasing Ba/La (Fig. 3.12a). In contrast, Th/La increases during the start of flare-ups, before decreasing,
creating cyclic Th/La peaks and troughs (Fig. 3.12b). The most significant changes occur during the
Cretaceous flare-up
4.3.2 Isotopes
At the scale of the entire arc-perpendicular transect,
87
Sr/
86
Sr i and εNd i ratios show distinct trends for
each flare-up through time and show a coupled relationship to each other (Figs. 3.13a-b). Data for the
Triassic flare-up are limited and represent the starting point to compare subsequent isotopic changes.
Overall, it is isotopically most similar to the Cretaceous arc. The Jurassic arc is the most juvenile, with low
87
Sr/
86
Sr i and high εNd i values (Fig. 3.13a-b). During the Cretaceous flare-up
87
Sr/
86
Sr i and εNd i trend to
more evolved values (Fig. 3.13a-b). Median Pb isotope values for the Jurassic and Cretaceous flare-ups
weakly trend towards mantle values (mantle defined by Chen and Tilton, 1991) (Fig. 3.13c-e). This is
60
consistent with
87
Sr/
86
Sr i and εNd i trends for the Jurassic arc, but is decoupled from
87
Sr/
86
Sr i and εNd i trends
for the Cretaceous arc.
4.3.3 Within-domain temporal patterns
In Supplementary Figure 1 (Appendix C), trace elements are color-coded by domain, in order to discern
patterns through time within each domain. They largely overlap and follow the same trends through time.
Because of this, only trends that represent exceptions to the ‘arc average’ are included below. The isotopes
also show a strong spatial variation (see section 4.3.2) (Fig. 3.13). Their temporal signal within-domains is
most clearly observed in
87
Sr/
86
Sr i and εNd i , and largely absent for Pb isotopes. For example, in the
Cretaceous flare-up, domains 1, 2, and 3 completely overlap in Pb composition. Below the general
characteristics of each domain and temporal trends are described:
Domain 1, Western Metamorphic Belt: Magmatism in domain 1 is the most compositionally distinct of the
four domains, with the lowest Rb/Sr, Sr/Y, and Th/La values and flat, MORB-like Rare Earth Element
(REE) patterns, which result in low La/Yb. Eu/Eu* and Dy/Dy* show the highest values compared to the
other domains. This domain includes the most juvenile isotope compositions, as whole-rock
87
Sr/
86
Sr i and
εNd i isotopes, spanning in age from <210 Ma to >120 Ma, are distinct from contemporaneous (within a 10
m.y. bin) magmas sampled in domain 3. Trends that are distinct to domain 1 magmas include the isotopic
shift during the mid-to-end of the Jurassic flare-up (ca. 150-140 Ma) towards more evolved
87
Sr/
86
Sr i, εNd i
, and
208
Pb/
204
Pb i values. At this time, there is a corresponding shift in trace element compositions, as Dy/Yb
and Ba/La start to increase.
Domain 2, Snow Lake Block: Domain 2 magmatic history is limited to the Cretaceous flare-up period, and
matches temporal trends described earlier. Briefly summarized, Rb/Sr, La/Yb, Sr/Y all increase, while
Zr/Hf and Dy/Yb decrease. Eu/Eu* becomes more restricted in range, and Dy/Dy* decreases. Ba/La
increases in the Early Cretaceous, then sharply decreases. Th/La increases, as does Zr/Y. Cretaceous
87
Sr/
86
Sr i and εNd i isotopes trend towards evolved values.
206
Pb/
204
Pb and
208
Pb/
204
Pb decrease during the
Cretaceous, towards mantle values, which indicates that they are decoupled from the Sr and Nd system.
Domain 3, Eastern Sierra Block: Domain 3 temporal history is well-summarized by the ‘arc average’ trends
described earlier (section 4.3.1). Ba/La is particularly elevated compared to the other domains, and La/Yb
is weakly elevated compared to domains 1 and 2. Notably, by ca. 110 Ma, the domain 3 trends are
indistinguishable from domain 2 trends.
87
Sr/
86
Sr i isotopes are relatively juvenile from the Triassic to the
Jurassic and εNd i isotopes are restricted to the most juvenile values (of the range found in this domain) in
the Jurassic, mimicking the trend of domain 1.
208
Pb/
204
Pb increases from the Triassic to the Jurassic,
decoupled from
87
Sr/
86
Sr i and εNd i behavior. Both
87
Sr/
86
Sr i and εNd i trend to more evolved compositions
61
during the Cretaceous, where they overlap with domain 2. Pb isotopes overlap with domain 2 in range but
show no temporal trend.
Domain 4, Owens Valley Block: In domain 4, samples show the lowest Eu/Eu* and Dy/Dy*, and slightly
elevated La/Yb and Sr/Y. Limited data from domain 4 show that from the Triassic through to the
Cretaceous, Rb/Sr, Zr/Hf, La/Yb and Sr/Y all decrease, opposite to the main, arc average trends. Zr/Y is
distinct from all other domains in that it also decreases through time. There is also limited isotopic data
from domain 4, but generally shows the most evolved
87
Sr/
86
Sr i isotopes, increasing throughout each flare-
up. In the Triassic and Cretaceous arcs, the isotopic composition overlaps with domain 3 isotopes. Pb
isotopes, restricted to the Triassic and Jurassic arcs, show no pattern.
4.4 Comparison to the zircon isotope record
Lackey et al. (2008) demonstrated that oxygen isotopes in zircon were sensitive to the spatial position
of the sample, defining mantle zones, which match with the geologic/bulk-rock isotopic boundaries
described above. Domain 1 requires input of hydrothermally altered oceanic materials to give highest
oxygen isotope ratios, while domain 3 has strongest mantle influence, producing the lowest oxygen isotope
values (Lackey et al., 2008). During the Cretaceous, where data are most heavily concentrated, there are
temporal trends towards mantle values, as the arc migrates from west to east. Paterson et al. (2014) Figure
6-15 illustrates whole-rock oxygen isotope space-time patterns (and recalculated zircon isotopes), which
extends the record back to the Jurassic and Triassic, showing that within each spatial region there are
temporal trends. Eastern Sierra samples trend towards juvenile values through time, whereas western Sierra
samples show increasing whole-rock oxygen isotopes from the Jurassic to Cretaceous (no Triassic data).
Attia et al. (2020) demonstrate that erupted zircons spanning all three arc flare-ups record primarily juvenile
mantle input between 70-90% of the overall composition, showing trends towards higher εHf
i values with
each progressive flare-up. Within a flare-up the isotopic range narrowed towards juvenile compositions.
4.5 Thermochronology
At the western and eastern extents of the arc, hornblende and biotite cooling ages are generally Jurassic
or older (Fig. 3.14). The central region of the arc is dominated by Cretaceous hornblende and biotite ages,
< 100 Ma (Fig. 3.14). Ages between 115-130 Ma were not found, resulting in a data gap.
The difference in zircon age and hornblende-biotite age can be used to evaluate thermal histories. In the
western portion of the arc, hornblende-biotite ages are either produced from cooling of the pluton (in cases
where they overlap or closely post-date the zircon age) or represent thermal resetting by the Jurassic
magmatism (age difference > 10 m.y.). In the central portion of the arc, there is a wide variation in zircon
62
Figure 3.14: K/Ar and
40
Ar/
39
Ar biotite and amphibole ages vs. re-projected west-east position. Samples are color-coded by the age
difference in the zircon U-Pb age and K/Ar-
40
Ar/
39
Ar biotite/amphibole ages. Cooling ages in the western arc are not dramatically reset by
Early Cretaceous magmatism, due to lower magmatic volumes. In the eastern part of the arc, some Triassic plutons show reset ages to 100-
80 Ma. Data sources listed in section 3.1.
63
ages, from Triassic through to Cretaceous, yet all hornblende and biotite ages are < 100 Ma. This reflects a
broad (approximately homogeneous) region of thermal resetting, as a result of Cretaceous magmatism. The
eastern region of the arc is similar to the western region, in that cooling ages are generally Jurassic-Triassic
in age. Samples which have Triassic crystallization (zircon) ages and Jurassic cooling ages, have been reset
by Jurassic magmatism.
4.6 Summary
The compilation of Mesozoic arc rocks illustrates several types of spatial geochemical patterns. These
patterns appear superimposed on the compositional differences between volcanic, hypabyssal, and plutonic
rocks of both mafic and felsic compositions (i.e., the spatial and temporal patterns are not controlled by
these attributes). This dataset supports interpretations from previous work that there are NW-SE striking,
sharp isotopic boundaries between distinct crustal domains, and that different domains have distinct
compositional characteristics (Kistler and Peterman, 1973; Kistler, 1990; Lackey et al., 2008; Memeti et
al., 2010; Attia et al. 2017). Importantly, several trace element ratios can also detect these geologic
boundaries, expressed as inflections in the cubic spline interpolation (e.g., La/Yb, Dy/Yb, Rb/Sr, Ba/La,
Th/La, Zr/Y). In addition, some trace element west-east patterns match certain isotope patterns (e.g., Th/La
and Sr
i isotopes, and Ba/La and Pb isotopes). This suggests similar behavior of these elements and isotopes,
and the possibility that they record the same magma source. Minerals that appear important when
considering spatial trends include amphibole and feldspar, as both Dy/Dy* and Eu/Eu* indicate increased
fractionation to the east.
Although a spatial trend across the arc is discernable for many of the trace element ratios studied here,
not all of the trace element ratios show a clear temporal trend at the Mesozoic timescale. For example, both
Dy/Dy* and Eu/Eu* lack temporal trends. More commonly, trace-elements and isotopes show distinct
temporal trends at the intra-flare-up scale (e.g., Rb/Sr, Zr/Hf, Th/La). Within domains, trace element
temporal patterns show considerable overlap, with few ‘distinct’ signals, but in contrast, the within-domain
isotopic trend is can be quite distinct. Temporal trends within samples from the same crustal domain can
span the entire Mesozoic (e.g., domain 3 trends) (Fig. 3.15a-d). This supports interpretations based on the
zircon record that there are changes in the magma sources through time, as well as space (e.g., Lackey et
al., 2008; Paterson et al., 2014; Attia et al., 2020). Importantly, bulk-rock isotopic trends are distinct for
each flare-up.
The Cretaceous arc is the time period in which trace element and isotope signals show the most dramatic
changes. This is particularly evident for the REE, including La/Yb, Dy/Dy* and Eu/Eu*, as well as for
87
Sr/
86
Sr i and εNd i isotope systems. It is also the only flare-up where we see decoupling between Sr and Nd
64
Figure 3.15: Comparing isotope systems to main magma reservoirs. Symbols and colors as in Figures 3.9 and 3.13.
(A) and (C): Initial Sr (
87
Sr/
86
Sr i) vs, Nd (
143
Nd/
144
Nd i converted to εNd i) isotopes, showing mantle reservoirs (DM,
HIMU, BSE, EMI, EMII), mid-ocean ridge basalt (MORB) and altered oceanic crust (AOC) labelled (Zindler and
Hart, 1986). Cretaceous Sierra Nevada peridotite and pyroxenite xenoliths (Ducea and Saleeby, 1998) are included
for comparison. (B) and (D):
208
Pb/
204
Pb vs.
206
Pb/
204
Pb showing the location of mantle reservoirs (DM, EMI, EMII,
HIMU) (Zindler and Hart, 1986). Western Cordilleran Pb isotope sources are drawn from Zartman (1974). Central
Sierra Nevada fields drawn from Chen and Tilton (1991) and pyroxenite xenolith field drawn from Ducea (1998).
65
isotope systems indicating evolved compositions, and the Pb isotope system, which indicates juvenile
compositions, matching interpretations of O in zircon and Hf in zircon (Lackey et al., 2008; Attia et al.,
2020). The Cretaceous arc in the CSN is defined by a regional thermal aureole, resetting biotite-hornblende
ages to <100 Ma (see also Barton, 1996). We posit that these distinct geochemical and thermal signals are,
in part, a result of the several dynamic arc processes operating during the Cretaceous arc (flare-ups,
migration, focusing, and crustal thickening) that are discussed in section 5 below.
5. Discussion
5.1 Deconvolving space and time in geochemical indicators
5.1.1 Spatial signal
Spatial isotopic patterns in the central Sierra Nevada have been interpreted to represent lateral changes
in the lithospheric architecture, inherited from the pre-arc basement (e.g., Kistler and Peterman, 1973;
DePaolo, 1981; Kistler, 1990). The overall trace element and isotopic characteristics of domains 1, 3, and
4 are distinguishable in Triassic, Jurassic, and Cretaceous samples (Figs. 3.6, 3.7, 3.8, 3.15). Thus, this
record remains relatively ‘static’ or ‘steady-state’ through time (see below for discussion of the temporal
signal). The components that could be changing in this model from west to east are the crust and underlying
mantle lithosphere of either continental or oceanic affinity.
Other subduction zone attributes that could cause compositional variation across the Sierra Nevada
include: (1) the depth of slab dehydration and mantle-wedge melting (k-h relationship), which has been
proposed to explain to increasing K 2O across the Sierra Nevada (Dickinson, 1975; Bateman and Dodge,
1970); (2) variation in the degree of partial melting, with lower degrees of partial melting away from the
trench (Stern et al., 1993); (3) mantle wedge source heterogeneity (Hochstaedter et al., 2000; Jacques et al.,
2013); and (4) slab fluid source heterogeneity as the slab progressively dehydrates with depth (Hochstaedter
et al., 2001). Importantly, most west-east trace element ratios and all isotope patterns are not displaying
linear increasing or decreasing eastward trends; instead they oscillate between high and low values and
show sharp inflections at known crustal boundaries. This supports the interpretations that the Sierra west-
east trace element and isotopic trends are primarily documenting changes in source materials (mantle and/or
lithosphere and/or fluid), rather than by gradients in subduction zone process (depth to melting, degree of
melting, progressive fluid dehydration).
Source materials derived from the mantle wedge are interpreted to be relatively constant across the arc
at any given time. This is likely because the mantle wedge is continually convecting by corner flow, and
because possible mantle indicators do not show simple arc-perpendicular trends (e.g., mantle: δ
18
O in
66
zircon, εNd i). As open-system processes are evident in forming the range in composition of Sierran magmas
(as above) then the use of Ba/La and Th/La are not appropriate for tracing slab fluids or sediments. Changes
in Ba/La in the analogous Andean arc have been interpreted to represent variation in crustal-mantle
interactions (Hildreth and Moorbath, 1988; Mamani et al., 2010), indicating that these ratios may track
changes in the deep crustal source type across the arc. In the case of these fluid-mobile elements, these
indicators match isotope trends of Sr i and Pb i, supporting this idea. Specifically, Ba increases in domains 1
and 3 within deep marine passive margin and oceanic materials and decreases within domains 2 and 4
within shallow marine/continental passive margin sediments. High Ba contents in the deep marine/oceanic
materials is documented in modern sediments (Church, 1979), metamorphosed cherts (Mottana, 1986), as
well as in deep marine Devonian metasediments of the Neoproterozoic-Paleozoic passive margin (Roberts
Mountain allochthon; Jewell and Stallard, 1991).
The crust and mantle lithospheric sources that vary laterally across the Sierra Nevada should become
depleted over time (as it is a finite source) as each subsequent arc flare-up built into the older one. Since
the overall domainal characteristics persist for all three flare-ups, it seems either these sources do not deplete
significantly (this reservoir is large), the source is ancient, thus does not evolve significantly (e.g., Coleman
and Glazner, 1997), the materials are replenished through time (via downward flow- works for crust only),
or the crust and mantle lithosphere sources represent only a small proportion of magmatic addition to the
arc (and this signal is retained via limited amounts of new lower crustal melting and/or magma recycling).
Cretaceous peridotite and pyroxenite xenolith compositions overlap almost entirely with Cretaceous upper
crustal samples, suggesting that much of the isotopic range can be explained by deep processing between
the mantle wedge, lithospheric mantle and variable interaction with the lowermost crust (Ducea and
Saleeby, 1998) (Figs. 3.9, 3.15). It hints that mantle reservoirs such as the depleted mantle wedge, or
enriched mantle lithosphere are dominant magma sources, and that crustal sources are secondary (Fig.
3.15). This is consistent with εHf
i mantle-crust mixing models from the Sierra Nevada (Attia et al., 2020),
and bulk-rock mixing models from the Sierra Nevada and other arc sections (DePaolo, 1981; Coleman et
al., 1992; Martinez-Ardila et al., 2019).
5.1.2 Amphibole and feldspar significance
Decreasing Dy/Yb and Dy/Dy* from west to east is a signal that persists for each of the three arc
sections, although is most pronounced for the Cretaceous arc. It suggests that as the arc moves inboard,
amphibole fractionation is more significant. A weaker, west-east feldspar signal, of decreasing Eu/Eu* is
also observed, suggesting that the importance of feldspar fractionation also increases inboard. This could
correspond to the development of a thicker, continental crustal column, including hydrated mantle
lithosphere that enables the formation of hydrous magmas in the overlying crust (see also Chin et al., 2012;
67
Chapman and Ducea, 2019). Hydrated mantle lithosphere has only been demonstrated for the Cretaceous
arc (Chin et al., 2012; Chapman and Ducea, 2019). The weak west-east trend of eastward increasing La/Yb
that suggests either: (1) the arc may thicken across-strike in the Triassic and Jurassic arcs, due to
continental-affinity basement having a proportion of continental mantle lithosphere (domains 2 and 4) that
is absent in oceanic-affinity basement (domains 1 and 3); or (2) increasing La/Yb is partly linked to
increased amphibole fractionation eastward (decreasing Dy/Yb). However, by the Cretaceous period,
La/Yb is increasing in all domains across the arc, while Dy/Yb shows strong spatial and temporal shifts to
decreasing values, suggesting they are now decoupled.
5.1.3 Temporal signal
Temporal patterns are apparent in the dataset between and within crustal domains, suggesting that there
is a record of temporal processes at the Mesozoic-scale and the intra-flare-up scale, superimposed on the
spatial signal. The domains each follow broadly the same temporal history, summarized in section 4.3.3
and Supplementary Figure 1. The patterns that are specific to each domain, from west to east, are discussed
below:
Domain 1, Western Metamorphic Belt: Domain 1 characteristics are consistent with the interpretations of
the western metamorphic belt having characteristics of fore-arc magmas, underpinned by oceanic
lithosphere and depleted mantle (Cady, 1975; Saleeby et al., 1989; Lackey et al., 2008; Saleeby, 2011).
Saleeby (2011) interpreted the Early Cretaceous isotopic signal also seen in southern portions of the arc to
represent geochemical maturation of the mantle lithosphere as arc magmatism continued. Lackey et al.
(2008) presented oxygen isotope data that requires involvement of hydrothermally altered crust in early
Cretaceous western Sierra magmas. Increasing Ba/La during this time could represent contributions from
assimilated oceanic crust, corresponding to increasing Sr
i isotope ratios. Increasing Dy/Yb, Sr/Y, and La/Yb
in the Early Cretaceous, is here interpreted to result from a thickening crust, and the involvement of garnet
in the source region.
Domain 2, Snow Lake Block: Although we only have information from the Cretaceous arc, the domain 2
trends are well sampled and document the largest changes within the arc’s history, from ca. 120-85 Ma.
Within domain 2, magma compositions trend towards more evolved values through time (Rb/Sr, Zr/Hf)
with a progressive signal of amphibole fractionation (decreasing Dy/Yb and Dy/Dy*). Sr/Y and La/Yb
significantly increase from ca. 120 Ma, during a period of flare-up magmatism, migration, focusing and
crustal thickening (see section 5.2 below). Thus domain 2 patterns are consistent with the continuation of a
domain 1 temporal signal. Whole-rock isotope decoupling between Sr and Nd, versus Pb and O zrc, is clear
in domain 2 and could have multiple sources: (1) Sr and Nd are tracking different source components, such
68
as the lower crust and mantle-lithosphere beneath the arc, while Pb and O track sedimentary/supracrustal
inputs; (2) Pb isotopes have been subsequently reset by fluids; or (3) Sr isotopes become decoupled from
Pb and O isotopes during alteration of crustal sources (Lackey et al., 2008).
Domain 3, Eastern Sierra Block: An earlier history is recorded in domain 3, which shows little change from
the Triassic to the Jurassic. The Triassic-Jurassic history is summarized by low La/Yb, Sr/Y, Dy/Yb and a
shift towards juvenile isotopic compositions. The Cretaceous element and isotopic history of domain 3 is
the same as domain 2 from ca. 110 Ma, indicating that any compositional distinctions between domains
had diminished by this time, and that an extensive, transcrustal magma plumbing system averaged out pre-
existing differences.
Domain 4, Owens Valley Block: This domain has the least amount of data but tends to show the most
evolved trace element patterns (low Dy/Dy*, decreasing Zr/Hf) and isotopic compositions (
87
Sr/
86
Sr i)
through time. The temporal trends within domain 4 are sometimes opposite to the arc average trends. This
is surprising, considering that the domains are consistently overlapping in composition, suggesting that this
could be an artifact of interpolating through the sparse dataset.
Overall, for the trace-elements studied, the difference in trends between domains through time appear
to be smaller than, or secondary to, the changes through time documented in all domains combined. In
contrast, the isotopes show that the isotopic characteristics of each domain is an important factor in forming
the isotopic temporal trends. In the Cretaceous, the domains 2 and 3 are indistinguishable, both in trace
element composition and isotopically. Next the signals that are likely linked to spatiotemporal processes
are examined:
5.2 Spatiotemporal dynamic arc processes
Processes that could modify magma trace element and isotopic compositions in space and time include:
(1) increasing the volume of mantle-derived magma added to the arc during flare-ups (e.g., Martinez-Ardila
et al., 2019; Attia et al., 2020); (2) changing the composition of the mantle-crustal end-members through
time due to arc migration; (3) pluton armoring, where younger magmatic pulses are shielded from
contamination by the crust, and potentially also hybridize due to magma focusing; and (4) tectonic
processes controlling magma generation, including root formation and delamination, due to magmatic and
tectonic crustal thickening.
5.2.1 Arc flare-ups
Element ratios that show patterns on flare-up timescales include Rb/Sr, Ba/Sr, Zr/Hf, Ba/La, and Th/La.
They could be tracking the changing magmatic conditions or processes between flare-ups and lulls.
69
Changing the crust:mantle ratio of arc magmas during flare-ups, by either increasing the volume of mantle
component, or the volume of the crustal component, could play a role in creating the recurring intra-flare-
up patterns.
In section 5.1.1, the mantle component (either enriched mantle lithosphere, or depleted mantle wedge,
or a combination) is considered the dominant source of Sierran magmatism, supported by εHf i in zircon
modeling throughout the Mesozoic (Attia et al., 2020). Thus, we consider how increasing the volume of
mantle-derived magma addition to the arc, with secondary modulation by the crust, could affect the upper-
crustal magma compositions sampled here. It is clear from both trace elements and isotopes that each flare-
up does not behave in the same, repeatable way, suggesting that a single mechanism does not seem
sufficient to explain all three flare-ups. Each flare-up varies in the volume of magma emplaced, with the
Cretaceous flare-up interpreted to be 2-8 times more voluminous than the previous two (Ratschbacher et
al., 2019).
Some trace element ratios suggest that the start of flare-ups is a period when incompatible elements
become enriched, relative to the end of flare-up periods (e.g., Rb/Sr, Ba/Sr, Zr/Hf, Th/La temporal signals).
This could result from higher magma addition rates, sourced from the mantle, heating the crustal column
and enabling greater degrees of fractionation. However, since differentiation within a single pluton can be
highly complicated and result from a myriad of local (and open-system) processes (e.g., Bateman and
Chappell, 1979; Gray et al., 2008; Memeti et al., 2014), it is not yet clear if a signal of a flare-up could be
discernable across several dozen intrusive complexes that each have complex trace-element histories.
Because Ba/La and Th/La show matching west-east patterns with Sr and Pb isotope systems
respectively, it is possible that they record changes in the lithology of pre-existing basement rocks (section
5.1.1.). In addition, the fluid and sedimentary inputs from the subducting slab and their variation during
flare-ups remain unconstrained and are not discernable in this thick continental arc section using these ratios
(e.g., Hildreth and Moorbath, 1988; Mamani et al., 2010).
5.2.2 Arc migration
Arc migration is not observed in the Triassic or Jurassic arcs in the CSN, based on the ages and spatial
position of exposed rocks (Fig. 3.4a). Thus migration does not explain spatial-temporal isotopic shifts in
the older arcs. Significant isotopic shifts during Cretaceous flare-up magmatism are associated with
eastward migration of the arc front, from ca. 140 Ma in the Great Valley to ca. 85 Ma in the east-central
Sierra Nevada (Chen and Moore, 1982; Saleeby, 2007; Ardill et al., 2018). The inflections in the isotopic
trends from west to east match up well with previously identified geologic and geochemically-defined
lithospheric boundaries (e.g., Kistler and Peterman, 1973; Saleeby et al., 1986; Kistler, 1990; Lackey et al.,
70
2008). Lateral changes in Dy/Yb and Eu/Eu* are most pronounced during the Cretaceous and have also
been attributed to the formation of hydrous arc magmas sourced from re-fertilized continental mantle
lithosphere during Cretaceous arc migration (Chapman and Ducea, 2019). However, this signal also occurs
(at a subdued level) during the Jurassic and Triassic, so it is not necessarily a good indicator of arc
migration.
5.2.3 Magma focusing
In order to explain the generally juvenile whole-rock isotope compositions during the Late Triassic to
Jurassic, there are a few possibilities: (1) the heat budget from the mantle-derived magmas was not large
enough to establish an extensive lower-crustal MASH zone, or incorporate substantial crustal or
lithospheric melting; (2) Jurassic magmatism was armored against the crust after the initial pulse of Triassic
magmatism due to spatially focusing rising magmas; (3) Jurassic and Late Triassic magmas are spatially
restricted to the western metamorphic belt or eastern-Sierra domains which are oceanic-affinity and thus
more juvenile (Kistler and Peterman, 1973).
The spatial location of Jurassic magmatism is part of the overall isotopic character of the flare-up period,
however, within one domain (domain 3), isotopes trend towards juvenile values from the Triassic to the
Jurassic, indicating that there must also be a temporal change. This is also supported by the observation that
Owen’s Valley Block plutons (domain 4), intruding the interpreted shallow-passive margin sequence
overlap in composition with the eastern Sierra Block plutons (domain 3), which intrude the deep marine
passive margin sequence. Both would suggest that process 3 is only part of the explanation. Studies estimate
that the flare-ups were volumetrically smaller than the Cretaceous (e.g., Paterson and Ducea, 2015),
suggesting a lower heat input to the arc overall, but also recognize that the Jurassic plutons and volcanic
rocks are spatially overlapping with the exposures of the Triassic arc. Thus it is not possible to determine
if the overall immaturity of the arcs (process 1) or pluton armoring via magma focusing (process 2) is
responsible for the juvenile signal. We have not identified regional magma focusing centers in the Jurassic
or Triassic arcs with this compilation; in the CSN this is partly a result of the Cretaceous arc overprinting
or removing the record of the earlier arcs.
However, it is a different situation in the Cretaceous flare-up, where
87
Sr/
86
Sr i and εNd i both trend
towards evolved values within individual domains. Pluton armoring is significant at the regional scale
within the CSN magma focusing zone between 102-85 Ma, where magmas progressively hybridize in larger
magma chambers (Memeti et al., 2010, 2014; Ardill et al., 2018). The element and isotopic compositions
of domains 2 and 3 are significantly overlapping by ca. 100-110 Ma, showing the same trends from the
initiation of focusing onwards. Particularly for the Sr and Nd isotope systems, but also suggested by
71
Dy/Dy*, and La/Yb, this indicates a switch from the spatial controls on composition (where crustal
boundaries are observed) to a signal where internal magmatic processes, such as mixing within several large
magma chambers are influential to forming the observed compositions across 10-20 m.y and across 4000
km
2
area. Since the period of magma focusing overlaps with the peak of the Cretaceous flare-up, it remains
unresolved whether magma focusing requires the high magma addition rates of flare-ups in order to
amalgamate several magma bodies.
5.2.4 Crustal thickening
Estimates of crustal thickness in the Sierra Nevada have been quantitatively estimated by strain-
measurements (Cao et al., 2016) and paleo-thicknesses to the Moho have been inferred using Sr/Y and
La/Yb ratios (Profeta et al., 2015). During the Triassic and Jurassic, crustal thicknesses were comparable
to average continental crust, ~30-40 km. Lower crustal thicknesses may be feasible, due to the onset of
marine deposition during the Late Triassic and Jurassic. Both Cao et al. (2016) and Profeta et al. (2015)
show substantial crustal thickening during the Cretaceous. In the CSN compilation, the western
metamorphic belt (domain 1) records the earliest period of crustal thickening, as recorded by La/Yb and
Sr/Y, starting at ca. 140-150 Ma. Using the Profeta et al. (2015) calibration, crustal thickness more than
doubled, from ~30 km at 140-150 Ma, to >70 km by 85 Ma. This temporal signal is orders of magnitude
larger than the slight eastward thickening documented between the different domains in the Jurassic and
Triassic.
A mohometer analysis using a comprehensive range of trace elements in this dataset yields broadly
comparable results to the Profeta et al. (2015) calibration for the Cretaceous arc but predicts slightly lower
crustal thickness values overall (P. Luffi, personal communication). It also identifies an Early Cretaceous
period of slower crustal thickening, followed by rapid crustal thickening starting ca. 100 Ma (P. Luffi,
personal communication).
5.3 Potential links or feedbacks between dynamic processes in the Sierra Nevada
Comparing the rates of different spatiotemporal processes indicates that they operated at ~1-10 km/m.y.
rates. Arc migration in the CSN occurred at 2.7 km/m.y. and magma focusing ranged between 1-4 km/m.y.,
depending on the direction of focusing; faster rates are found on the western margin of the focusing region,
where they may combine with eastward migration rates (Ardill et al., 2018). Crustal thickening is, on
average, estimated at 2 km/m.y., but this is likely an overestimate of Early Cretaceous thickening and an
underestimate after 100 Ma (section 5.2.4). Rates of magma addition during the peak of the flare up were
estimated by Ratschbacher et al. (2019) at 1.35 km
3
/km
2
/m.y. In a related process, Cao et al. (2016)
estimated the downward transfer of host rock between 1-10 km/m.y. over the duration of magmatic and
72
73
Figure 3.16 (previous page): Cretaceous Central Sierra Nevada dynamic crustal section, summarizing the
spatiotemporal patterns of flare-ups, migration, focusing and crustal thickening. (A) Schematic crustal column,
showing the change in magmatism between 140-85 Ma. The background is color coded to represent the different
types of basement across the arc (e.g., Fig. 3.1), where yellow represents continental affinity basement, and green
represents oceanic or transitional arc basement. The LHS of the column shows a ~30 km thick crust, that begins to
migrate, with moderate volume magma plumbing system, and the initiation of magmatic-tectonic crustal thickening.
The RHS of the diagram shows how magmatism has evolved, moving eastward, and increasing magma addition
rates. Crustal thickness has more than doubled. We interpret that focusing helped facilitate the formation of large
upper-crustal magma chambers, which fed by mid and lower crustal magma chambers. Inset box describes models
that have been proposed to explain how magma focusing can increase the size of magma chambers. Rates of
processes are shown and demonstrate overlapping values. Data sources linked to figure: [1] Cao et al. (2016); [2, 3]
Chin et al. (2012, 2014); [4] Coleman et al. (1992); [5] Kistler (1990); [6] Lackey et al. (2008); [7] Saleeby (2007);
[8] Ratschbacher et al. (2019); [9] Karlstrom et al. (2009). (B) Isotopic timeline, illustrating the decoupling between
Sr and Nd systems, relative to Pb, O and Hf systems, between 125-115 Ma, which is here interpreted as a flare-up
signal of increasing mantle-derived magma addition.
74
tectonic thickening. Therefore, these processes operated synchronously, and possibly set up feedback loops
that promoted the growth of the high-volume magma plumbing system (Fig 3.16a). Chapman and Ducea
(2019) proposed a link between flare-ups and arc migration; this association is apparent for the Cretaceous
arc but does not apply to the Triassic or Jurassic arcs. Similarly, each of the known magma focusing centers,
in volcanic and plutonic settings occurs during a flare-up (e.g., de Silva et al., 2006; Grunder et al., 2008;
Lipman, 2007). However, this requires further investigation.
Flare-ups and arc migration are generally considered to be driven by lower lithosphere processes, either
in the slab, mantle wedge, the overlying mantle lithosphere, or the coupling between all these components
(DeCelles et al., 2009; Kirsch et al., 2016; Karlstrom et al., 2014; Holt et al., 2015). The effects of these
processes are expressed at multiple crustal levels. In contrast, crustal thickening and magma focusing
operate within and throughout the crustal column (Fig. 3.16a).
5.4 Comparison between the Triassic, Jurassic, and Cretaceous arcs
The Triassic and Jurassic arcs had average continental crust thicknesses (30-40 km) and were laterally
extensive (Fig. 3.17). So far, although exposures of these arcs are limited in the CSN compared to the
Cretaceous, we have not found evidence for arc migration or magma focusing processes. The overall
juvenile isotopic signatures could reflect chemical, thermal, or mechanical processes in the crust that limited
the involvement of pre-arc basement rocks volumetrically. In general, they show the same west-east
geochemical patterns as Cretaceous rocks, (although we cannot test domain 2, the Snow Lake Block, due
to the lack of Triassic-Jurassic magmatic products exposed). Temporal trends are similar in the Triassic and
Jurassic and appear more ‘muted’ relative to the Cretaceous flare-up.
The Cretaceous arc is the period where we see the largest shifts in geochemical signals through time,
indicating that dynamic processes described above have the capacity to dramatically change arc lithosphere,
especially when operating synchronously. Each of the processes has potentially conflicting and contrasting
geochemical signals, which may be an additional explanation why some isotopes (e.g., Sr, Nd vs. Pb, O,
Hf systems) are decoupled (Fig. 3.16b). Sample characteristics are a composite of all of these processes,
and not only at the scales described above, but also involving processes at the local, magma reservoir scale
(e.g., Memeti et al., 2014). For example, arc migration, which is tracking the spatial expression of the arc
flare-up, modifies the magma isotope composition as different basement types are introduced into the melt
zone, and this could be exacerbated by extending the length of the lithospheric column (e.g., Sr and Nd
signal). Due to increased mantle melting in the flare-up, juvenile contributions during the Cretaceous flare-
up are also apparent in Pb, O and Hf systems, which is distinct from the ‘crustal’, or evolved, Sr and Nd
migration trend.
75
Figure 3.17: Block diagram highlighting the contrasts between the Triassic and Jurassic arcs (left column) and the
Cretaceous arc (right column) using the findings from this study. They are interpreted to represent immature and
mature systems, respectively. BDT=Brittle-Ductile Transition. Brown and hashed pattern shows the laterally
varying lithospheric basement. In the left hand column, purple colors indicate plutons in the magma plumbing
system, and the blue triangular regions depict vertical focusing of magmatism from broad lower crustal magma
zones towards discrete volcanic centers. In the right hand column, pink colors indicate plutons and green and yellow
show assimilation of crustal pieces. The dashed orange line indicates that the magmatism generates a broad thermal
aureole. Magmatism in blue and grey beneath the Moho depicts mantle lithosphere refertilization processes as
described by Chin et al. (2012). Plutonic percentages illustrated on the right hand side.
76
5.5 Formation of Cretaceous upper-crustal MASH zones
The 1,100 km
2
, nested Tuolumne Intrusive Complex (TIC) was emplaced at the center of the magma
focusing zone, between 95-85 Ma, during the latter stages of the arc flare-up, and during rapid crustal
thickening. Isotope trends indicate extensive hybridization of magmas with the TIC as magma focusing
continued, which is interpreted to result from the formation of larger magma chambers, regions of
interconnected melt and crystals (Figs. 3.16a, 3.17). Widely varying types of datasets have come to the
same conclusion; namely geochemical, structural, and computational approaches (Žák et al., 2007; Paterson
et al., 2011; Memeti et al., 2014; Paterson et al., 2016). This also appears to be a thermally-viable
interpretation, as the Cretaceous cooling ages cover the entire CSN region, extend well-beyond the mapped
boundaries of the TIC. The youngest biotite cooling ages are 82-80 Ma, in the youngest magmatic region,
the NE of the TIC. The interpretation that the upper- to mid-crust of the CSN was the site of extensive
MASH-like processing within a thermally-primed region, suggests that this occurred throughout the crustal
column, as part of a transcrustal magma plumbing system (e.g., Ward et al., 2017; Miller et al., 2009;
Cashman et al., 2017) (Fig. 13.6a, 3.17).
5.6 Implications of dynamic behaviors on the arc lithosphere
The arc behaviors studied here effectively ‘primed’ the arc, resulting in rapid batholith construction,
extending from the surface into deep magma source regions, that replaced 70-90% of the arc volume with
igneous material (Ratschbacher et al., 2019). Within the upper-crust, these processes contributed to the
growth and evolution of large (>100 km
2
), long lived (>0.5-1 m.y.) magma chambers (Memeti et al., 2014;
Paterson et al., 2016), that by the Late Cretaceous imparted a regional-scale thermal metamorphic signal
onto the surrounding crust (Kerrick, 1970; Barton, 1996). The rapid thickening of the crust formed a
mountain range which influenced weathering, erosion and sediment production (Cao et al., 2016). It may
have also had an effect on magmatic compositions by increasing the length of magma differentiation
pathways (Karlstrom et al., 2014). This substantial magmatic and tectonic evolution of the system may have
resulted in a geochemical switch where the importance of spatial, arc-wide signals is reduced, and the
importance of internal, magmatic differentiation processes increases.
Comparing the Sierran example to volcanic systems that display primed magmatic behaviors, the
process may have also modified the physical and chemical connections between plutons and volcanoes, in
a transition from low volume, generally more mafic and effusive eruptions, towards phases of caldera-
forming eruptions (Lipman, 2007; de Silva et al., 2006). In the Sierra Nevada, the Late Cretaceous Minarets
and Merced Peak calderas provide a glimpse into the scale and types of volcanism that were active during
77
arc-priming, and future work aims to test the physical connections between plutons and volcanoes at
different stages of the priming process by examining hypabyssal intrusions (e.g., Ardill et al., in press).
6. Conclusions
Distinguishing between the physical and chemical signals of crustal thickening, arc flare-ups, migration,
focusing and pluton differentiation remains challenging because the rock record represents a composite of
all these processes, as well as recording some of the effects of spatial position in the arc, and the evolution
of magma sources and magmatic conditions through time. However, some signals resulting from arc-
migration and magma focusing are discernable due to their distinct spatial or temporal patterns. Flare-up
magmatism remains the most elusive process to identify with the compiled dataset; overall our findings of
temporal isotopic trends within and between domains, that match trace element patterns, support previous
studies that mantle-derived magmas are the dominant source of flare-up magmas, and that the crustal
contribution is subordinate. Importantly, each flare-up has a different geochemical pattern, suggesting that
one simple, or end-member flare-up model is not sufficient to completely explain the Sierran magmatic
systems.
Results leave open the possibility that the dynamic processes operating during the Cretaceous could be
inter-connected as they operate at similar rates (1-10 km/Myr). This likely contributed to thermal,
mechanical, and chemical feedbacks in the crustal column that promoted voluminous, long-lived magma
storage in the upper crust. While some spatial arc attributes are pervasive throughout the arc’s lifetime (e.g.,
Kistler, 1990; Moore, 1959), the lithosphere is modified most substantially from the Late Jurassic to the
Late Cretaceous. Dynamic processes culminated in formation of a transcrustal, mature arc system in the
Cretaceous.
7. Acknowledgements
We acknowledge support from National Science Foundation grants EAR 1624847, EAR 1019636 to
Scott Paterson, and EAR 1624854 to Vali Memeti, a Geological Society of America Graduate Student
Research Grant (2017), and the University of Southern California (USC) Department of Earth Sciences
Graduate Student Research Fund. We thank Heidi Carmen-Garcia for assistance with drafting.
78
Chapter 4: Reconstructing the physical and chemical development
of a pluton-porphyry complex in a tectonically re-organized arc
crustal section, Tioga Pass, Sierra Nevada
This chapter is in press at Lithosphere
Ardill, K., Memeti, V., and Paterson S.R., in press. Reconstructing the physical and chemical development
of a pluton-porphyry complex in a tectonically re-organized arc crustal section, Tioga Pass, Sierra Nevada,
Lithosphere
Abstract
In ancient or partially eroded arc sections, a protracted history of tectonism and deformation makes
interpretation of local volcanic-plutonic relationships challenging. The fragmentary preservation of
volcanic rocks relative to the extensive plutonic record in upper-crustal arc sections also suggests that a
broader-scale approach that includes volcanic-hypabyssal-plutonic ‘fields’ is useful. In this context, studies
of hypabyssal intrusions emplaced at the intersection of volcanic and plutonic fields provide additional
physical and chemical constraints on shallow-level magmatic processes.
New mapping, U-Pb zircon geochronology, and geochemistry at Tioga Pass, in the central Sierra Nevada
arc section, document the physical and chemical evolution of the Tioga Pass hypabyssal complex; a ca. 100
Ma system that includes an intrusive dacite-rhyolite porphyry unit and co-magmatic Tioga Lake quartz
monzodiorite. We interpret these units as a Cretaceous subvolcanic magma feeder system intruding a
package of tectonically displaced Triassic and Jurassic volcanic and sedimentary rocks, rather than the
previous interpretation of a Triassic caldera. The Tioga Pass magmatic system is a well-exposed example
of a hypabyssal complex with meso- to micro-scale structures that are consistent with rapid cooling and
emplacement between 0–6 km depth and compositions suggestive of extensive fractionation of largely
mantle-derived magma. The Tioga Pass porphyry unit is one of many hypabyssal intrusions scattered along
a ~50-kilometer-wide belt of the east-central Sierra Nevada that are spatially associated with coeval
volcanic and plutonic rocks due to tectonic downward transfer of arc crust. They provide a valuable
perspective of shallow magmatic processes that may be used to test upper-crustal plutonic-volcanic links
in tectonically-reorganized arc sections.
79
1. Introduction
Studying the volcanic-plutonic connection in ancient and partially eroded arc sections presents some
unique challenges, particularly from a temporal and geochemical perspective. Exposed sections, often
biased towards deeper plutonic levels, may provide a somewhat time-averaged (yet longer-lived) record of
arc activity, while volcanic units are susceptible to the effects of hydrothermal alteration and metamorphism
(e.g., Hanson et al., 1993; Sorensen et al., 1998). Syn- to post-arc deformation and tectonism overprints the
magmatic record and can re-organize and displace local and regional stratigraphy, masking true field
relationships.
However, tectonic shuffling is also one of the primary reasons to study these crustal sections. Tectonism
often results in the exposure of multiple crustal levels, providing information on the physical and spatial
relationships of both plutons and contemporaneous volcanic rocks, and can additionally reveal the physical
structure (e.g., shape, size) of the subvolcanic magma plumbing system. Restoring local- to regional-scale
structures allows for inclusion and evaluation of magmatic-tectonic relationships through mapped spatial
and temporal associations of structures and magmatic rocks. Together, these factors suggest that ancient
arc crustal sections provide a different perspective to understanding plutonic-volcanic relationships
compared to modern systems.
In upper-crustal arc sections, hypabyssal intrusive complexes are important features as they are exposed
in both volcanic- and plutonic-dominated arc domains (e.g., Kistler and Swanson, 1981; Colgan et al.,
2018). They represent a physical connection between these environments, which is key to reconciling
existing plutonic-volcanic observations and models (e.g., Lipman, 1984; Bachmann et al., 2007; Glazner et
al., 2015; Bachmann and Huber, 2016). As hypabyssal complexes are emplaced at shallow levels, they have
the potential to provide a snapshot of magmatic processes that relate to both plutons and volcanic rocks,
such as crystal accumulation, melt extraction, and eruption (e.g., Zimmerer and McIntosh, 2013).
In the Sierra Nevada arc section, several studies have drawn attention to the Mesozoic volcanic-plutonic
connection, either by regional-scale comparisons between mid- to upper-crustal plutons and volcanic
stratigraphic sections (e.g., Bateman, 1992; Barth et al., 2011, 2012, 2018; Sisson and Moore, 2013;
Klemetti et al., 2014; Greene et al., 2017), or by the study of local volcanic features, such as calderas (e.g.,
Fiske and Tobisch, 1994; Lowe, 1995; Schweickert and Lahren, 1999). More information on shallow levels
of the plumbing system, accessed through the hypabyssal record, improves our understanding of the
physical development and evolution of upper-crustal magma plumbing systems and has the potential to link
observations made across wide spatial scales, from a single intrusion to a regional-scale magmatic ‘field’.
80
At Tioga Pass, in the eastern-central Sierra Nevada, we re-interpret a proposed Triassic caldera
(Schweickert and Lahren, 1999) as a Late Cretaceous hypabyssal complex. We present new field mapping,
U-Pb zircon ages, and whole-rock element and isotopic geochemical data from Tioga Pass to characterize
magmatic and tectonic histories from Triassic to Late Cretaceous time. We demonstrate that both the Tioga
Lake quartz monzonite and intrusive dacite-rhyolite porphyry are ca. 100 Ma and represent co-magmatic
parts of a punched laccolith shaped body (steep-sided and flat-topped, with vertical walls discordant to host
rock structure; Corry 1988) that intruded blocks of faulted and tilted Triassic and Jurassic volcanic and
sedimentary strata. The Tioga Pass system shares compositional and structural affinity with other Late
Cretaceous intrusive and volcanic rocks in the east-central Sierra Nevada region and is one of several
hypabyssal intrusions found across a ~50-kilometer-wide belt in the central Sierra Nevada arc section.
These intrusions are key structural markers and may represent feeders to volcanic eruptions or stalled late
melts from plutons.
2. Previous work
2.1 Regional Background
The Mesozoic Sierra Nevada Batholith is an upper- to lower-crustal composite arc section, dominated
by magnesian, calc-alkaline, metaluminous tonalite, hornblende-biotite granodiorite and locally
peraluminous granite. In the central Sierra Nevada, emplacement depths of Triassic-Cretaceous intrusive
suites range from 6–10 km based on Al-in-hornblende barometry (Ague and Brimhall, 1988b; Anderson et
al., 2008; Foley et al., 2007). The 120–85 Ma Cretaceous flare-up resulted in emplacement of voluminous
upper-crustal intrusive complexes (e.g., Bateman, 1992; Paterson and Ducea, 2015). Cretaceous volcanic
sections are locally exposed in several pendants of the Sierra Nevada Batholith, including the Ritter Range,
Saddlebag Lake and Piute pendants (Fig. 4.1) (Greene and Schweickert, 1995; Schweickert and Lahren,
2006; Memeti et al., 2010b; Chapman et al., 2012; Paterson and Memeti, 2014), although most of the
Cretaceous volcanic cover is interpreted to have been eroded. Calculated magma addition rates, using
Cretaceous plutonic and volcanic rock areas, resulted in plutonic-volcanic ratios between 20:1 to 30:1
during flare-up magmatism (Paterson and Ducea, 2015).
2.2 Geology of the Saddlebag Lake Pendant
The Saddlebag Lake pendant exposes Paleozoic deep marine framework rocks and a sequence of
Mesozoic arc volcanic and sedimentary packages (Brook, 1977; Schweickert and Lahren, 2006; Barth et
al., 2012; Paterson and Memeti, 2014; Attia et al., 2018) (Figs. 4.1 and 4.2). It is intruded by the Triassic
Scheelite Intrusive Suite in the east and the Cretaceous (95-85 Ma) Tuolumne Intrusive Complex to the
81
Figure 4.1: Simplified geologic map of the east-central Sierra Nevada region, modified from the Geologic Map of
Yosemite National Park by Huber et al. (1989), illustrating the extent of Cretaceous plutonic (pink), hypabyssal
porphyry (orange) and volcanic (green) fields. Dashed rings include areas of the Minarets and Merced Peak calderas
as reported by: [1] Fiske and Tobisch (1994) and [2] Lowe (1995). Numbers indicate names of known Cretaceous
hypabyssal intrusions: 1-Tioga Pass Porphyry-see Figure 4.2 for detailed map; 2-Shellenbarger Lake Porphyry; 3-
Post Peak Porphyry; 4-Star Lakes Porphyry; 5-Red Peak Porphyry; 6-Ireland Lake Porphyry; 7-Johnson Peak
Porphyry; 8-Beartrap Lake Porphyry. Locations of metavolcanic and metasedimentary pendants are labelled, and the
extent of Triassic-Jurassic volcanic rocks within each pendant is shown in the cross-hatched pattern. The boundaries
of the Piute Meadow, Saddlebag Lake and northern Ritter Range pendants are shown with a dotted line. Uncolored
areas indicate regions of Paleozoic metasediments or Quaternary cover. At the southern end of the map, the
uncolored region marks the southern extent of the study area. Inset map shows the Sierra Nevada Batholith and
location of the study area (box).
82
83
Figure 4.2 (previous page): Geologic map of the Tioga Pass area, based on new 1:10,000 scale mapping and
compilations of maps by Russell (1976), Greene (1995), and McColl (2017). Unit colors indicate rock type (see
legend) and units are labelled by estimated age (Pz, TR, J, K) and rock type. In some cases, units are labelled by age
and unit name if given: TRsl – rhyolite tuff of Saddlebag Lake; TRbm – rhyolite tuff of Black Mountain. Uncolored
regions denote areas of Quaternary cover, for example, alluvium, talus, or glacial moraine features. Transects A-A’
and B-B’ and C-C’ refer to cross sections through the Saddlebag Lake pendant and northern contact of intrusive
rocks, through the southern extent of the Saddlebag Lake pendant and intrusive rocks, and through the Dana
sequence, respectively (see Figure 4.3A-C). Cross section of transect D-D’ is shown in Figure 4.6. Sense of slip on
major faults is labeled where constrained. Stereographic projections: Great circles of the mean bedding of the Tioga
Lake section (grey line; N=8 measurements; strike/dip = 323/81), Saddlebag Lake pendant (black line; N=75;
163/86), Dana sequence section (dotted line; N=26 measurements; 317/79) and porphyry intrusion magmatic layers
(dashed line; N=17; 141/71); (right): Great circles of the mean foliation within the Tioga Lake section (N=12;
323/81), in the surrounding Saddlebag Lake pendant strata (N=75;160/85) and the porphyry/quartz monzodiorite
magmatic foliation (N=26; 147/83). [1] in legend refers to the location of the published age of the rhyolite tuff of
Saddlebag Lake by Schweickert and Lahren (1999). For U-Pb zircon sample information, refer to Table 4.1.
Samples TP-7 and TP-14 not shown, located ~500 m north of area shown in Figure 4.2.
84
west (Bateman, 1992; Memeti et al., 2010a; Barth et al., 2011). Regional metamorphism pervasively altered
rocks to greenschist facies (Kerrick, 1970; Brook, 1977); the prefix “meta-” is hereafter omitted from rock
descriptions where the protolith is known. Unconformable contacts between westward-younging major
stratigraphic packages were subsequently reactivated by strike-slip faulting (Paterson and Memeti, 2014;
Cao et al., 2015; Attia et al., 2018). Tioga Pass is located at the southeastern end of the Saddlebag Lake
pendant (Figs. 4.1 and 4.2), marking the transition of the Saddlebag Lake pendant into the northern Ritter
Range pendant.
The distinctive lithologic units exposed west of Tioga Lake were previously mapped and studied by
Kistler (1966a), Brook (1977), and Schweickert and Lahren (1999) (Fig. 4.2). Interbedded sandstone and
tuff units, discordant to regionally NW-striking volcanic units, could not be stratigraphically correlated to
nearby Saddlebag Lake or northern Ritter Range pendants. Geologic interpretations from field relations
varied widely, from hypabyssal pluton emplacement along a normal fault, a shallow volcanic vent structure,
and a caldera structure (Brook, 1977; Schweickert and Lahren, 1999). Schweickert and Lahren (1999)
interpreted a Triassic age for the proposed caldera collapse event, caldera fill sequence, and resurgent
intrusions based on contact relationships between an intracaldera vent structure and the dated 222 Ma tuff
of Saddlebag Lake (TRsl; Fig. 4.2).
The Dana sequence is a ~1 km thick layered package of volcanic and sedimentary units located southeast
of Tioga Pass, in the vicinity of Mt. Dana, and has been linked to units at Tioga Lake (Fig. 4.2; Kistler and
Fleck, 1994; Greene, 1995; Schweickert and Lahren, 1999). Russell (1976) found basal shallow marine
interbedded volcanic and sedimentary deposits, grading upwards into terrestrial ash-flow tuff, cross-bedded
sandstone and conglomerate. Intrusive units mapped at Tioga Pass, such as the quartz monzodiorite and
dacite porphyry unit, were also found at Mt. Dana (Kistler, 1966a; Greene 1995). Several intrusions and
faults mask original contacts of the Dana sequence with surrounding units. Kistler (1966a) and Russell
(1976) assigned a lower Jurassic age to the Dana sequence, based on lithologic correlations to the Dunlop
Formation (Nevada) and nearby Ritter Range pendant. Schweickert and Lahren (1999) reinterpreted this
section as part of the Triassic caldera-fill package.
3. Methods
We remapped the Tioga Pass area at 1:10,000 scale, compiled map data from the literature (Kistler,
1966; Russell, 1976; Greene, 1995; Schweickert and Lahren, 1999) and collected U-Pb zircon laser
ablation-inductively coupled plasma mass spectrometry (LA-ICPMS) data to temporally constrain regional
stratigraphy. We dated tuff units at the base and the top of volcanic packages, defined by regional
85
unconformities or faults, to bracket the range of possible ages for coherent volcanic sequences, and dated
sedimentary units to estimate maximum depositional ages. We dated cross-cutting intrusions to provide
crystallization age estimates. Detailed 1:1,000 scale mapping, petrography and whole-rock major, trace and
rare earth element geochemistry focused on the Tioga Pass intrusive units. Stereographic projections were
created using the program Stereonet 9.8 (Allmendinger et al., 2013)
3.1 U-Pb Zircon Geochronology
LA-ICPMS U-Pb zircon analyses were performed at the University of Arizona Laserchron Center using
conventional lab procedures (e.g., Gehrels et al., 2008, 2009; Gehrels, 2014). Full analytical methods and
data tables are included in supplementary materials. Sri Lanka, FC-1, and R33 zircon grains were used as
primary age standards (Gehrels et al., 2008). For igneous samples, 20–30 zircons were analyzed where
possible, and a weighted mean age calculated, reported at the 2σ level. We included in our weighted mean
calculation only concordant zircons with U/Th < 10 as defined by Gehrels et al. (2009) for igneous zircons.
This excluded a small number of zircons (n=17 grains) from the total sample set of igneous samples (n=199
grains). For detrital zircon ages from sedimentary units approximately 100 zircons were analyzed, where
possible. The location of the youngest zircon peak, made of at least 3 zircons with overlapping uncertainty,
is interpreted as the maximum depositional age of the sample (Dickinson and Gehrels, 2009). For published
LA-ICP-MS ages where sample data are available (e.g., Paterson and Memeti, 2014), we re-interpreted age
estimates using the above protocol in order to improve consistency across samples.
3.2 Whole-rock Geochemistry
Sixteen samples were collected for whole rock elemental geochemistry to characterize the porphyry
(n=10) and the quartz monzodiorite unit (n=6). Fourteen samples were analyzed by X-ray fluorescence
(XRF) at Pomona College using lab procedures outlined in Lackey et al. (2012). Five samples were
analyzed by inductively coupled plasma mass spectrometry (ICP-MS) for major oxides and/or trace
elements at Activation Laboratories, Ancaster, under the “4LithoResearch” package. Three samples were
analyzed for trace elements by laser ablation-inductively coupled plasma mass spectrometry (LA-ICP-MS)
at Texas Tech University. Two samples were selected for whole rock isotopic analysis performed at the
University of Arizona, following methods in Otamendi et al. (2009). The standard error for isotope samples
is between 0·0007%-0.0008% for
87
Sr/
86
Sr and 0·0009-0.0018% for
143
Nd/
144
Nd, and between 0.003-
0.008% for Pb isotopes. Detailed analytical methods are included in supplementary materials.
86
4. Results
4.1 Stratigraphy of the Study Area
The wide-ranging geologic interpretations at Tioga Pass highlights the importance of combining a
detailed stratigraphic and field-based record with modern geochronology to resolve outstanding
uncertainty. Stratigraphic, field, and geochronologic relationships are described below for the intrusive
complex and the surrounding host rocks. We use these to establish which units are spatially and/or
temporally associated, restore tectonic effects, and examine the contact relationships between units. We use
previously defined regional terms, such as the Triassic ‘Koip sequence’ and Triassic-Jurassic ‘Dana
sequence’, to describe and group temporally and spatially related strata (Kistler, 1966a; Bateman, 1992).
The upper Paleozoic deep marine section forms the eastern boundary of the Tioga Pass complex, where
fine-grained sandstones are interbedded with chert and phyllite (Pzs; Figs. 4.2, 4.3A-C). A detrital zircon
maximum depositional age of 435 Ma was attained from a fine-grained sandstone sample northwest of
Tioga Lake (sample TP-8; Table 4.1, Figs. 4.2, 4.4 and 4.5). The lithology and age of overlying Triassic
Koip sequence metavolcanic rocks in the Saddlebag Lake pendant have been described by Brook (1977),
Barth et al. (2011), Paterson and Memeti (2014), and Cao et al. (2015). At Tioga Pass, the rhyolite tuff of
Black Mountain (TRbm; Figs. 4.2, 4.3A, 4.4) defines the base of the Koip sequence and yielded an age of
229.6 ± 1.8 Ma (sample TP-14, Fig. 4.5). This age estimate is overlapping with a U-Pb zircon sensitive
high-resolution ion microprobe (SHRIMP) age from Barth et al. (2011) of 232 ± 2 Ma. The overlying
Cooney Lake conglomerate (TRcl; Figs. 4.2, 4.3A-B, 4.4) contains meter-wide pebbly layers interbedded
with coarse sandstone. Detrital zircon ages for Cooney Lake conglomerate samples form a narrow unimodal
distribution with minimum peaks at 219-220 Ma (samples KA-9, TP-7, TP-1; Fig. 4.5). The overlying 224
± 1 Ma rhyolite tuff of Saddlebag Lake (TRsl; Figs. 4.2, 4.3A-B, 4.4) (Barth et al., 2011) separates the
Cooney Lake conglomerate from a package of crystal-rich flows of andesite-dacite composition (TRan;
Figs. 4.2, 4.3A, 4.4). The uppermost section of the Koip sequence (TRuv; Figs. 4.2, 4.3A-B, 4.4) includes
layers of clastic rhyolite-dacite tuff, crystal-rich andesite flows and volcanic breccia, thinly inter-bedded
with volcanoclastic conglomerate and sandstone. One sample of a rhyolite breccia east of the Maul Lake
fault yielded an age of 220.0 ± 2.2 Ma and constrains the top of the preserved section of Koip sequence
metavolcanic rock units at Tioga Pass (Figs. 4.4 and 4.5; sample KA-6).
The ten-meter-wide, brittle Maul Lake fault (NW corner of Fig. 4.2) juxtaposes Koip sequence units
(TRuv) with the ca. 95 Ma Kuna Crest granodiorite (Kkc; Figs. 4.2, 4.3A, and 4.4; Memeti et al., 2010a).
Jurassic metasedimentary rocks and Cretaceous volcanic rocks form part of the stratigraphy further north
87
Figure 4.3: Cross sections from (A) the Saddlebag Lake pendant, (B) Tioga Lake, and (C) Dana sequence transects.
See Figure 4.2 for locations and legend. Unit abbreviations are listed in the Figure 4.2 legend and within the text. In
cross sections, vertical scale = horizontal scale. MLF= Maul Lake Fault and GPF= Gaylor Peak Fault, marked on
Figure 4.2.
88
Table 4.1: Summary of LA-ICP-MS U-Pb zircon data for igneous and detrital samples collected from Tioga Pass. The age range of zircon
indicates the range of ages found in zircons analyzed per sample. See methods for an explanation of procedure, and supplementary materials
for additional sample information.
89
Figure 4.4: Schematic stratigraphy of units in the Tioga Pass area. Refer to legend in Figure 4.2 for explanation of
colors and unit labels. Approximate thicknesses of units are labelled. Bold lines represent faults. Wavy lines
represent unconformities reactivated by faulting. Pluton emplacement is shown on the left-hand side of the columns
where it intrudes stratigraphic packages. Stars indicate where a U-Pb zircon age (igneous and detrital) was obtained.
Blue stars represent ages of plutonic units, black stars represent ages of volcanic and sedimentary units. Interpreted
correlation between sections is shown with a solid line. Jurassic and Cretaceous rocks in the Tioga Pass section are
extrapolated from maps of nearby Sawmill Canyon (~3 km north of Tioga Pass; Paterson and Memeti, 2014).
Abbreviations: MLF = Maul Lake Fault; GPF = Gaylor Peak Fault; d.z. = average of youngest peak of detrital
zircon population. [1] refers to the published age from the tuff of Saddlebag Lake by Barth et al. (2011) north of
Tioga Pass. [2] refers to re-interpreted ages for volcanic rocks from Paterson and Memeti (2014) (see Methods and
Supplementary Materials).
90
Figure 4.5: Normalized age probability plots illustrating sample age distribution for Paleozoic-Jurassic sedimentary
samples and Triassic to Cretaceous volcanic samples. Ages determined by LA-ICPMS U-Pb zircon geochronology.
See Table 1 for sample locations, rock descriptions and additional information. Note change in scale in center plot at
300 Ma.
91
in the Saddlebag Lake pendant (e.g., Sawmill Canyon, Paterson and Memeti, 2014; Virginia Canyon, Cao
et al., 2015) (Figs. 4.1, 4.4), and south in the Ritter Range pendant (Fiske and Tobisch, 1978), however they
are not exposed at Tioga Pass. They were likely entirely removed during intrusion of the Tuolumne
Intrusive Complex (Fig. 4.4) (Paterson and Memeti, 2014; Cao et al. 2016).
4.2 Tioga Lake and Dana Sequence Sections
Tioga Lake Section
A coherent, NW-striking and SW- to NE-dipping section of volcanic and volcaniclastic strata forms an
approximately 800 x 400 m block exposed west of Tioga Lake (TRvs, TRan, and TRrh; Fig. 4.2). This
volcanic block is intruded on all exposed sides by a dacite-rhyolite porphyry unit and quartz monzodiorite
unit (Kdp, Ktm; Figs. 4.2, 4.3B, 4.4). Units include volcaniclastic sandstone, pebbly, monomict
conglomerate, and andesite and rhyolite composition lava flows and tuff. The section has a total thickness
of ~850 m. Meter-wide dacite porphyry dikes intrude these units at a high angle. A maximum depositional
age of 217 Ma from a volcaniclastic sandstone in the intra-porphyry package (sample TCL 13-5; Figs. 4.4,
4.5) is similar to the youngest Koip sequence units in the Saddlebag Lake pendant.
Dana Sequence Section
The Dana sequence is a fault-bounded package of interbedded metavolcanic and metasedimentary units
that is restricted to the area around Mt. Dana. It is surrounded by the Paleozoic chert-argillite unit. The
package grades from mudstones and intermediate to felsic tuffs and flows (TRuv and TRtm; Figs. 4.2, 3C)
into a ~200-meter-wide rhyolite tuff (TRrh; Figs. 4.2, 4.3C) and overlying volcaniclastic sandstone unit
(TRvs; Figs. 4.2, 4.3C), consistent with descriptions from previous studies (Russell, 1976; Greene, 1995).
An andesite tuff from the exposed base of the Dana sequence yielded a weighted mean age of 212.8 ±
6.4 Ma, however, this sample yielded only 7 zircons, 6 of which were concordant (sample MD-8C; Figs.
4.2, 4.4, 4.5). A rhyolite flow ~200 m below the summit of Mt. Dana yielded an age of 195.1 ± 2.1 Ma
(sample L101; Figs. 4.2, 4.4, 4.5). A fault-bounded rhyolite tuff yielded an age of 221.7 ± 2.5 Ma and could
be part of either the Koip or Dana sequence (sample MD-9; Figs. 4.2, 4.3C, 4.4, 4.5).
92
4.3 Tioga Pass Intrusive Units
Quartz Monzodiorite
The Tioga Lake quartz monzodiorite (Ktm; Fig. 4.2) crops out north, west and east of Tioga Lake and
has an elongate shape in map view (Fig. 4.2). It was previously named the granodiorite of Tioga Lake
(Bateman, 1992; Schweickert and Lahren, 1999, ‘unassigned granitic rocks’ by Kistler, 1966a) and ranges
in composition from quartz monzodiorite to granodiorite (Le Maitre et al., 2004). It is an equigranular
biotite- (~5-10%) and hornblende-rich (~10%), plagioclase dominated (60–70%) unit with quartz (15-
30%); grain size varies between 0.5 and 4 mm. The unit is finer grained and porphyritic towards the western
intrusive contact with the porphyry unit (Figs. 4.2 and 4.6). The quartz monzodiorite typically grades into
the porphyry intrusion close to the contact and contains abundant rounded to sub-angular porphyritic
enclaves throughout (Fig. 4.7A).
A ~50-meter-wide mutually intrusive contact zone separates the quartz monzodiorite unit from the
adjacent porphyry; however, in some outcrops, the units have a sharp intrusive boundary (Figs. 4.6, 4.7B).
The quartz monzodiorite intrudes Paleozoic metasediments along a near vertical contact on northern and
eastern exposures and contains stoped blocks of metasediments and metavolcanic tuff from 1 to 100-meter
scale (Fig. 4.3A-B). Close to the contact with the Paleozoic quartzite, plutonic samples contain clots of
biotite.
Petrography: Plagioclase is the dominant mineral in the hypidiomorphic quartz monzodiorite unit. Distinct,
resorbed cores and well-defined oscillatory rim zoning are observed throughout the pluton (Fig. 4.7C).
Plagioclase crystals have euhedral to subhedral habit, show frequent plagioclase-plagioclase contacts and
are almost always sutured with other plagioclase crystals (Fig. 4.7C-D). Euhedral to subhedral hornblende
crystals contain inclusions of magnetite. Biotite is interstitial, < 0.5 mm in size, with minor replacement by
chlorite. Quartz occupies interstitial spaces between feldspar crystal clusters and shows undulose extinction
and recrystallization indicating low temperature crystal plastic deformation.
Porphyry Intrusion
The dacite-rhyolite porphyry (Kdp; Fig. 4.2) crops out as several large NW-striking elongate intrusive
bodies across the Tioga Pass area. The largest continuous body is exposed west of Tioga Lake. The
porphyry has an aphanitic quartz-biotite groundmass and euhedral plagioclase and alkali feldspar (~75–
80%), biotite (~10–15%) and hornblende (~5%) phenocrysts. The unit grades from a low silica dacite to
high silica rhyolite to the west.
93
Figure 4.6: WSW-NNE cross section (Line D-D’ in Figure 4.2) between the porphyry and quartz monzodiorite units, summarizing
field observations. Letters A-L refer to photos in Figure 4.7 (Photo A is not located along the transect and this is indicated by the
arrow). Field observations from the surrounding areas were projected onto this cross-section. Vertical scale = horizontal scale. Each bar
on the grain size scale represents a field measurement of average phenocryst size (mm). Filled bars represent mafic minerals, e.g. biotite
and hornblende, whereas unfilled bars represent the measurement of feldspar phenocrysts.
94
Figure 4.7: Field photos and photo micrographs of Tioga Pass samples. Letters correspond to locations in Figure
4.6. Photos A-D are from the quartz monzodiorite unit. (A) Rounded to angular porphyritic enclaves are common
within the quartz monzodiorite unit (photo from Mt. Dana area, outside of Figure 4.6 transect); (B) Tioga Lake
quartz monzodiorite unit at sharp contact with dacite porphyry dike; (C) Relict cores and zoning within a plagioclase
crystal from the quartz monzodiorite, sample KA23; (D) Typical microstructure of the Tioga Lake quartz
monzodiorite unit, with euhedral-subhedral plagioclase and interstitial quartz, sample KA23. Photos E-L are from
the porphyry unit. (E) Crystal rich portion of the porphyry unit has 2-3 mm phenocrysts of feldspar set in an
aphanitic grey groundmass; (F) Mingling zone at western edge of the porphyry.
95
Figure 4.7 (cont.): Field photos and photo micrographs of Tioga Pass samples. Letters correspond to locations in
Figure 4.6. (G) Clastic layers are found in the porphyry, with parallel orientation to flow bands; (H) porphyry clast
within rhyolitic portion of intrusive porphyry unit; (I) View of flow banded rhyolite portion of the porphyry unit.
Dashed line marks the contact between a dark and light flow band, sample KA13; (J) Typical microstructure of the
porphyry, sample KA17; (K) Glomerocryst cluster in the porphyry unit, sample 254; (L) Flow banded rhyolite in the
porphyry unit, with fractured feldspar phenocryst. Fractures contain interstitial quartz and alkali feldspar. In the top
right of the image, an example of a glassy inclusion ~0.2 mm in size. Smaller examples are found in the matrix,
sample TP16-4b.
96
The porphyry is internally structurally complex at the micro- and meso-scale (Fig. 4.6). Quartzite, pelitic
schist, rhyolite tuff, basaltic-andesite and andesite lava flows make up displaced blocks (up to tens of meters
scale) that are incorporated into, and intruded by, the porphyry. Meter-scale rounded inclusions of the quartz
monzodiorite unit are also observed within the porphyry intrusion (Fig. 4.6). Abundance of phenocrysts
varies from <5% to 50%, while phenocryst size ranges from 1-4 mm. Crystal-rich (45-50%) dacite is
observed near the contact with the Tioga Lake quartz monzodiorite, as well as in dispersed zones throughout
the porphyry unit (Fig. 4.7E). Crystal-poor rhyodacite to high-silica rhyolite (<10% phenocrysts) forms the
western part of the intrusion, with extensive mingling zones of intermediate composition magma and flow-
banded rhyolite (Figs. 4.6 and 4.7F). Grain-size generally decreases to the west. Massive and layered flow-
banded zones of the porphyry unit are interpreted to be devitrified obsidian, from outcrop and micro-scale
observations of phenocryst and matrix grain size, composition and local extent of the layering. Volcanic
clasts are prominent in some layers of the unit (Figs. 4.6, 4.7G). Clasts of crystal-rich dacite are observed
within the rhyolite portions of the porphyry (Fig. 4.7H).
The porphyry has a discordant intrusive contact with Paleozoic metasediments and Lower Koip
sequence units at Tioga Lake. In the east, the porphyry intrudes the Dana sequence and Paleozoic
metasediments. Contacts are sharp and near-vertical (Fig. 4.2, 4.3A-B). The western contact of the porphyry
with Saddlebag Lake pendant is offset <100 m by the Gaylor Peak fault. The contact between the porphyry
and quartz monzodiorite is typically gradational, except where the porphyry cross-cuts the quartz
monzodiorite unit as dikes or sheets.
Petrography: In thin section, the porphyry has a microcrystalline groundmass (Figs. 4.7I-L). Most
groundmass appears devitrified to quartz, feldspar, mica, and oxides, although glass remains intact as
inclusions or as amorphous aggregates 0.2–0.5 mm (Fig. 4.7L). In flow banded layers the groundmass
wraps around aligned phenocrysts forming a magmatic foliation (Fig. 4.7I, 4.7L). Hornblende phenocrysts,
observed in samples of intermediate composition, are euhedral-subhedral in shape and 1–3 mm in size.
Some hornblende phenocrysts are replaced by biotite at their rims. Biotite phenocrysts are often euhedral
and tabular mm-size grains, in some cases deformed (Fig. 4.7J), while groundmass biotite is
microcrystalline (up to 0.1 mm). Hornblende and biotite decrease in abundance from the contact zone
towards the western exposed edge of the porphyry unit.
Plagioclase phenocrysts are ubiquitous in the porphyry unit (Figs. 4.7I-K). Plagioclase is the most
abundant feldspar in the section of the porphyry unit closest to the contact, at approximately a 3:1 ratio of
plagioclase to alkali feldspar. Westwards, the proportion of alkali feldspar phenocrysts increases to 1:4 ratio
of plagioclase to alkali feldspar (sample KA13; Fig. 4.6). Alkali feldspar phenocrysts in sample KA13, a
97
flow-banded rhyolite portion of the porphyry (Figs. 4.6 and 4.7I) are ~1 mm in size and are coarser in the
dark flow bands relative to the light bands. Feldspars have oscillatory zoning preserved throughout the unit
and there is a population of grains identified by truncated zones and contact melting points between two
grains. Some phenocrysts are fractured and infilled with melt (Fig. 4.7L). Plagioclase phenocrysts are often
arranged in clusters (Fig. 4.7K). Signs of local alteration by fluids include quartz-epidote veins and chlorite
lenses in the groundmass (Fig. 4.7L).
U-Pb zircon ages of intrusive units
The Tioga Lake quartz monzodiorite yielded a U-Pb zircon weighted mean age of 101.0 ± 1.5 Ma
(sample TIOGA; Fig. 4.8A), older than the 96 Ma U-Pb zircon age (n=3 samples) reported by Schweickert
and Lahren (1999). The porphyry unit was dated at the northern and southern extents of the unit and yielded
ages of 100.0 ± 1.3 Ma and 99.5 ± 1.6 Ma (samples KA17 and TP17), indistinguishable in age within
uncertainty from the Tioga Lake quartz monzodiorite (Fig. 4.8B-C). Low MSWD values for these samples
indicate overestimation of uncertainty (Wendt and Carl, 1991) (Table 4.1). The 99.1 ± 1.4 Ma age estimate
of an outcropping granite unit (sample 3-3; Kg on Fig. 4.2) at the Kuna Crest granodiorite margin also
overlaps in age with the quartz monzodiorite and dacite porphyry unit. In this sample the age of the youngest
peak, at 98-99 Ma, is interpreted as the unit age, as the high MSWD value indicates that the current
weighted-mean model does not suitably fit the data (Wendt and Carl, 1991; Reiners et al., 2018) (Fig. 4.8D,
Table 4.1). This age estimate for sample 3-3 is older than 93-95 Ma isotope-dilution thermal ionization
mass spectrometry (ID-TIMS) age estimates for the Kuna Crest margin ~10 km south of Tioga Pass
(Memeti et al., 2010a). In summary, the granite is contemporaneous with the dacite porphyry and quartz
monzodiorite, but the petrogenetic relationship of the granite to these units remains uncertain.
4.4 Structural Features of the Study Area
Structural data from Mesozoic volcanic and sedimentary rocks and Cretaceous intrusive rocks are used
to evaluate the timing and impacts of local tectonism in deforming and displacing units related to the
intrusions, and to provide information on the three-dimensional structure of the intrusive complex.
A bedding-parallel foliation and cleavage is well exposed in Paleozoic and Mesozoic rocks. with an
average foliation of strike/dip=160/85 in the mapped area (Fig. 4.2). Mineral stretching lineation (measured
in biotite and quartz) plunges between 70–90 °. The Triassic Tioga Lake block (Figs. 4.2, 4.3B and 4.4)
displays bedding rotated ~30° counterclockwise relative to the average orientation of Saddlebag Lake
pendant units in the mapped area, with an average orientation of 133/83, also seen in foliation measurements
(average orientation 323/81; Fig. 4.2). Dana sequence units are eastward-dipping and gently to tightly
98
99
Figure 4.8 (previous page): LA-ICPMS U-Pb zircon age distributions for Cretaceous intrusive units at Tioga Pass.
From base of section to top: (A) sample TIOGA from the Tioga Lake quartz monzodiorite unit, (B) sample KA17 of
the porphyry unit, (C) sample TP17 of the porphyry unit, and (D) sample 3-3 of the granite unit. Probability density
function plot shown in a red line. The zircon grain age in sample 3-3 represented by a filled circle was identified as
an outlier in Isoplot (Ludwig, 2003). Horizontal bars for individual grains represent 2σ uncertainty, circles represent
grain ages. Vertical dashed lines indicate the calculated weighted mean age, grey bars show the analytical
uncertainty. In sample 3-3, the blue line represents the age of the youngest peak, which is slightly younger than the
mean weighted age. See Table 4.1 for sample locations, descriptions and additional information.
100
folded, with an average orientation of 318/79 (Fig. 4.2). Magmatic foliations (147/83) and well-defined
flow bands (141/71) in the porphyry unit are oriented in a NW-SE direction (Figs. 4.2 and 4.6). Along a
SW-NE transect, layering of flow-bands, clastic layers and mingling zones generally dip between 55–70°,
shallower than the regional structural grain (Figs 4.2 and 4.6).
The Sierra Crest Shear Zone is a ~300 km-long dextral-transpressive shear zone active in the eastern
Sierra Nevada, recording ductile deformation from at least 95 Ma to 84 Ma based on pluton ages and biotite
cooling ages, switching to brittle behavior between 84-80 Ma (Tikoff et al., 2005; Jiang and Bentley, 2012;
Paterson and Memeti, 2014; Cao et al., 2015; Hartman et al., 2018). The porphyry intrusion has a solid-
state foliation defined by deformed quartz grains and biotite phenocrysts that roughly parallels the earlier
magmatic foliation. This solid-state foliation occurs at the western margin of the porphyry, where it
intersects with the eastern margin of the Sierra Crest Shear Zone (‘Cretaceous ductile shear zone’ in Fig.
4.2). The eastern margin of the Sierra Crest Shear Zone also deforms Triassic andesite flows and breccias
in the Saddlebag Lake pendant. The western edge of the shear zone extends to the margin of the Kuna Crest
granodiorite of the Tuolumne Intrusive Complex. Shear bands, boudinaged quartz veins and σ-
porphyroclasts are common structures within sheared parts of the porphyry intrusion and generally indicate
dextral shear.
Faults in a NW striking brittle fault system are locally quartz-filled or associated with mineralization of
tourmaline and epidote. Dextral strike slip motion is the dominant fault style based on offset markers
between stratigraphic units and measured slickenlines. The Gaylor Peak fault, a Jurassic thrust fault
(Schweickert and Lahren, 1999; 2006), is one example of a reactivated through-going brittle fault in the
area. The mapped fault trace cross-cuts the ca. 100 Ma porphyry unit (Fig. 4.2), suggesting that the fault
was re-activated in the Late Cretaceous. Sub-horizontal slickenlines are found along the fault trace,
indicating strike-slip movement on the fault. Several smaller, near-vertical strike-slip brittle faults cross-
cut the porphyry unit with minor offsets (<100 m). These faults represent the transition from ductile to
brittle behavior of the shear zone (Hartman et al., 2018).
4.5 Whole-rock Geochemistry of Tioga Pass intrusive rocks
Major and trace element whole-rock geochemistry from Tioga Pass intrusive rocks (Fig. 4.9) illustrates
wide inter- and intra-unit compositional variation. Fields of Cretaceous plutonic, hypabyssal, and volcanic
rock samples, between 105 and 95 Ma, from the central Sierra Nevada (data sources listed in Fig. 4.9 and
supplementary materials) are shown for comparison with the Tioga Pass units. All four groups show similar
101
102
Figure 4.9 (previous page): Elemental geochemistry data. See Table 4.2 for sample names and analytical data. (A)
MgO (wt. %) vs. SiO 2 (wt. %); (B) K 2O (wt. %) vs. SiO 2 (wt. %); (C) Ba (ppm) vs. K 2O (wt. %); (D) Sr (ppm) vs.
CaO (wt. %); (E) V (ppm) vs. TiO 2 (wt. %); (F) Sc (ppm) vs. CaO (wt. %); (G) Y (ppm) vs. SiO 2 (wt. %); (H) Rare
earth element patterns normalized to chondrite (Boynton, 1984). Fields include Cretaceous volcanic, hypabyssal and
plutonic samples in the central Sierra Nevada. Data sources for fields: Peck and Van Kooten (1983); Lowe (1995);
Ratajeski et al. (2001); Memeti (2009); Cao et al. (2015); Ardill et al. (2018) and this study. Blue arrow points
towards the top of the exposed section in the porphyry unit. This is not shown on C, as there is no clear trend from
the base of the unit to the top.
103
Table 4.2: Whole rock major oxide, trace element and isotope compositions of samples collected from the Tioga
Pass intrusive complex. XRF data were collected at Pomona College (Method a), and Activation Laboratories
(Actlabs) (Method b). ICP-MS data were collected at Actlabs (Method c), and LA-ICP-MS data were collected at
Texas Tech University (Method d). Isotopic data were collected at the University of Arizona. B.D.L = Below
detection limit. - = Not analyzed. See Table 4.2 for sample specific method and Supplementary materials for
detailed analytical methods.
104
trends with overlap in composition. Plutonic rocks span the widest range in composition, while hypabyssal
samples show a narrow range in composition.
Major elements
Quartz monzodiorite samples contain between 55 and 64 wt.% SiO 2, whereas the silica content in the
porphyry unit ranges between 63-73 wt.%. Al 2O 3, CaO, Fe 2O 3 (total Fe expressed as Fe 2O 3), TiO 2, MnO,
P 2O 5 and MgO decrease in abundance with increasing SiO 2 contents in both units (Fig. 4.9A; Table 4.2).
Samples of the porphyry unit are dominantly peraluminous with ASI between 1.02 and 1.15 (ASI= molar
(Al 2O 3)/[(CaO-3.33*P 2O 5) + Na 2O + K 2O)]). One metaluminous porphyry sample (sample KA20) with ASI
0.86 crops out at the contact between the pluton and the porphyry. The quartz monzodiorite has an ASI
between 0.90 and 1.06. The single peraluminous sample in this unit (KA26) is also from the contact zone.
These samples from the contact have similar major element compositions and also share a compositional
affinity with porphyry and quartz monzodiorite samples collected from the southwest and northwest slopes
of Mt. Dana, respectively.
One sample of the quartz monzodiorite (sample KA23; 56 wt.% SiO 2) has high amounts of the major
elements, e.g., 6.64 wt.% CaO, 9.42 wt.% Fe 2O 3 and 1.19 wt.% TiO 2, compared to the other samples. A
porphyritic mafic magmatic enclave (sample TP16-1a) with mm-size feldspar phenocrysts has a similar
SiO 2 content to KA23, but has higher MgO, Al 2O 3 and lower Fe 2O 3. Within the porphyry unit there is a
compositional trend in major elements linked to spatial location, as the porphyry becomes gradually more
felsic towards the west (Fig. 4.6; Fig. 4.9 arrows); a pattern that is not observed in the quartz monzodiorite
unit.
Alkali enrichment
Samples in both units show a trend of increasing K 2O and decreasing Na 2O with increasing SiO 2 (Fig.
4.9B). Barium and Rb also increase with increasing K 2O (Fig. 4.9C). Sample KA13L contains anomalously
high concentrations of K 2O and Ba (7.41 wt. % and 3208 ppm respectively). In addition, elevated K/Na,
K/Al and Ba/Ti in sample KA13L suggests that this sample experienced pervasive alkali metasomatism
(Sorensen et al., 1998).
Minor and trace elements
Strontium decreases with increasing SiO 2, and with decreasing CaO in the porphyry unit (Fig. 4.9D),
both decreasing towards the top of the mapped section (Fig. 4.9 arrow). In the quartz monzodiorite, Sr
varies from 500-600 ppm but does not vary with CaO. In both units V and Sc decrease with increasing SiO 2
105
and increase with increasing TiO 2 and CaO (Fig. 4.9E, F). With increasing SiO 2, Y, Zr, and Rb in the
porphyry unit increase, but Y decreases slightly in the quartz monzodiorite unit; Zr and Rb in the quartz
monzodiorite increase (Fig. 4.9G). Compositional overlap between units is evident in all the above elements
between 62-65 wt.% SiO 2. Rare earth element (REE) patterns of the porphyry unit and the quartz
monzodiorite display similar normalized La/Lu ratios of 7.3-11.1 and 8.5-15 respectively, although samples
from the porphyry have a slight negative Eu anomaly that is absent in quartz monzodiorite samples (Fig.
4.9H).
Isotopes
Initial strontium and neodymium isotope ratios (
87
Sr/
86
Sr i and εNd) for the quartz monzodiorite sample
(KA23) are 0.704997 and -0.29, respectively. In the porphyry sample (KA17) these ratios are 0.705494 and
-1.33. εNd and
87
Sr/
86
Sr i values are within the mantle array (Zindler and Hart, 1986) and are within the
isotopic range of Cretaceous Sierra Nevada peridotite and pyroxenite xenoliths (Ducea and Saleeby, 1998)
(Fig. 4.10A). Initial Pb isotopic ratios (
207
Pb/
204
Pb,
208
Pb/
204
Pb,
206
Pb/
204
Pb) in the quartz monzodiorite are
15.67433, 38.97684 and 19.15353, respectively. The porphyry unit has initial Pb ratios (
207
Pb/
204
Pb,
208
Pb/
204
Pb,
206
Pb/
204
Pb) of 15.65512, 38.67369, and 18.9539, respectively. Lead isotopes from the porphyry
and quartz monzodiorite unit are also consistent with xenolith data (Ducea, 1998) and are within the field
defined by western US passive margin deep marine sediments (Zartman, 1974) (Fig. 4.10B).
5. Discussion
The Tioga Pass area of the central Sierra Nevada contains a Cretaceous hypabyssal magmatic complex
intruding into regionally extensive host-rock strata of the Saddlebag Lake and northern Ritter Range
pendants, as well as the Tioga Lake and Dana sequence sections (Fig. 4.11). Evidence for this new
interpretation includes: 1) ca. 100 Ma U-Pb zircon ages of the porphyry and quartz monzodiorite; 2) the
gradational contact between these two units and overlap in compositions indicating co-magmatic
emplacement; and 3) field and petrographic features characteristic of shallow (sub-volcanic) emplacement.
Field evidence combined with element and isotope compositions are permissive of a ‘magma feeder system’
model (Fig. 4.11) where fractionation was significant in producing the compositional zoning and
peraluminous rhyolite from mantle-derived magmas. Specifically, amphibole (ASI ~0.5; Zen, 1986) and
clinopyroxene (ASI ~0; Zen, 1986) fractionation can drive the resulting melt to peraluminous compositions
(e.g., Zen, 1986; Nandedkar et al., 2014; Clemens et al., 2020).
106
Figure 4.10: Whole-rock isotopes from Tioga Pass and select regional constraints. (A) Age-corrected
87
Sr/
86
Sr i vs.
εNd with mantle reservoirs (DM, HIMU, BSE, EMI, EMII), mid-ocean ridge basalt (MORB) and altered oceanic
crust (AOC) labelled (Zindler and Hart, 1986). Cretaceous Sierra Nevada peridotite and pyroxenite xenoliths (Ducea
and Saleeby, 1998) are included for comparison. (B)
208
Pb/
204
Pb vs.
206
Pb/
204
Pb showing the location of mantle
reservoirs (DM, EMI, EMII, HIMU) (Zindler and Hart, 1986). Western Cordilleran Pb isotope sources are drawn
from Zartman (1974). Central Sierra Nevada fields drawn from Chen and Tilton (1991) and pyroxenite xenolith field
drawn from Ducea (1998). Data sources [1] Ardill et al., (2018). [2] Lowe (1995). [3] Memeti et al. (2014). [4] Gray
et al. (2008).
107
Figure 4.11: Cartoon of the Tioga Pass magmatic complex at 100 Ma, showing the relationship between the quartz
monzodiorite, the dacite-rhyolite porphyry unit and surrounding host rocks. The magmatic complex is in the shape
of a punched laccolith, interpreted from cross-sections through the complex. The fault shown is the Gaylor Peak
Fault, reactivated after intrusion of the complex. Colors correspond to units in the Figure 4.2 map and Figure 4.2
legend. The Tioga Pass system most closely represents a ‘magma feeder system’ model.
108
5.1 Tioga Pass Cretaceous Magmatic Complex
Weighted mean ages from the Tioga Lake quartz monzodiorite, granite and porphyry unit are
indistinguishable within uncertainty. As the contact between the granite and the other Cretaceous intrusions
is concealed, interpretation of the relationship to the granite unit is hindered. However, field evidence from
the quartz monzodiorite and the porphyry unit demonstrate that they are co-magmatic across a contact that
varies from gradational to sharp along strike. In addition, inclusions of the quartz monzodiorite are found
within the porphyry and vice versa. The overlap in major and trace elements between units (at 62-65 wt.%
SiO 2) near the contact zone, and the overlap in REE patterns, is consistent with a co-magmatic relationship.
Observations consistent with a hypabyssal interpretation include the discordant intrusive contact of the
porphyry unit and quartz monzodiorite with the surrounding host rocks, the porphyritic texture, the glassy
inclusions and clasts, as well as the glomerocrysts and fractured crystals. The glass inclusions in particular
are indicative of rapidly cooled magma. Devitrification is interpreted where the glassy inclusions are found
together with the recrystallized matrix (Lofgren 1971, 1974). Layers including clastic material suggests
interaction with the surface, or near-surface, environment, either during magma ascent or possible eruption.
Plagioclase and alkali feldspar phenocrysts are clustered, forming the distinctive glomero-porphyritic
structure of the porphyry unit. In this case, the preferential contacts of like minerals are observed (e.g.,
plagioclase-plagioclase clusters are abundant), which is a common feature of hypabyssal rocks (e.g., Vance
and Gilreath, 1967). Melt extraction from a crystal mush and/or the tectonic disruption of the mush could
produce feldspar glomerocrysts in porphyries and volcanic rocks (Seaman, 2000; Bachmann and Bergantz,
2004; Bachmann et al., 2007; Beane and Wiebe, 2012). Many of the phenocrysts are fractured, which could
occur by volcanic eruption, a decrease in pressure due to magma ascent, or the rapid “ungluing” of crystal
groups (Vance, 1969; Bachmann et al., 2002; Beane and Wiebe, 2012; Cashman et al., 2017).
The quartz monzodiorite transitions from an equigranular to porphyritic texture close to the contact with
the porphyry unit, but the largest textural and structural variation is within the porphyry unit. The quartz
monzodiorite and porphyry share similar mineralogy but differ in modal proportions of quartz, alkali
feldspar and biotite (more abundant in the western extent of the porphyry) and there is a decrease in
hornblende and biotite phenocryst abundances westwards in the porphyry unit. The porphyry unit is
stratified by rock type and structure, with layers uniformly dipping (Fig. 4.6). Topographically, the
porphyry grades from a crystal rich dacite to a flow banded rhyolite at higher levels, and many of the major
and trace elements are zoned from the base of the section to the top (Figs. 4.9 and 4.11). This suggests that
layers in the porphyry unit preserve their original orientation and have not been steeply tilted. Our tests of
sample alteration (e.g., Sorensen et al., 1998, alkali enrichment) suggest that this is a primary chemical
109
gradation. Further, the compositional zoning implies that either the porphyry is made up of multiple pulses,
or gradually fractionated from a dacite to a high-silica rhyolite (Fig. 4.12).
Plagioclase appears to be an important mineral defining the intermediate-felsic composition of the
intrusive units. It is likely controlling the decrease in Sr, and the high concentration of Sr and CaO in some
quartz monzodiorite and porphyry samples suggests that these samples accumulated plagioclase. The quartz
monzodiorite does not show the negative Eu anomaly of the porphyry unit, nor a positive anomaly, but
retains a subtle signature of feldspar accumulation at upper-crustal levels. This interpretation is consistent
with observations of the elevated plagioclase content, plagioclase-plagioclase contacts in the quartz
monzodiorite and plagioclase glomerocrysts of the porphyry unit. Zoning of K 2O and Rb across the mapped
section suggest an important role for K-feldspar fractionation, except in sample KA13L.
The porphyry is distinct from the quartz monzodiorite in that it has a peraluminous composition (with
exception of one sample at the contact zone reported above). In addition to the role of fluids and volatiles
(e.g., Zen, 1988), one explanation is that the porphyry may contain a higher proportion of highly
fractionated melt (preserved in the rock as glass or recrystallized glass matrix). The metaluminous quartz
monzodiorite on the other hand is likely a cumulate that lost melt and accumulated feldspar. Another
possibility is that the porphyry unit assimilated some of the surrounding host rocks (Triassic volcanic rocks
or Paleozoic chert/argillite unit), supported by field observations of stoped blocks (Figs. 4.6 and 4.11).
However, these blocks are found in both units, so assimilation of the blocks does not entirely explain the
peraluminosity of the porphyry unit. Generating peraluminous rhyolite compositions in the porphyry unit
from a mantle-derived magma (e.g., Lackey et al., 2008; discussion below) suggests that fractionation
played a significant role in generating the compositional variety observed in the porphyry unit.
A depleted mantle or oceanic crust (MORB) magma source (εNd > +7.5; Zindler and Hart, 1984) does
not fully explain the magmatic source of the Tioga Pass intrusions, which have negative εNd values, unless
the depleted mantle source was progressively enriched and modified during earlier arc activity. In addition,
the samples do not share isotopic affinity with enriched mantle reservoirs (EMI and EMII; Fig. 4.10A).
However, Sr, Nd and Pb isotope ratios of the porphyry unit and quartz monzodiorite match Cretaceous
lower crust and mantle lithosphere xenolith compositions (Ducea and Saleeby, 1998), suggesting that the
porphyry and quartz monzodiorite magmas are largely mantle-derived. Further, Lackey et al. (2008)
measured δ
18
O in zircon from the Tioga Lake quartz monzodiorite (5.98‰) and found values slightly
elevated above the mantle range (5.3±0.3‰; Valley et al., 1998). Initial Pb isotopes indicate that the
magmas incorporated deep-water marine sediments of the western Cordilleran passive margin (Zartman,
1974; Fig. 4.10B), which is compatible with field observations of the Paleozoic chert-argillite unit in contact
110
Figure 4.12: Cartoon illustrating differences between plutonic-volcanic models. In magma feeder system models
(e.g., Lipman, 1984; Bachmann and Bergantz, 2004), processes such as crystal fractionation and melt extraction to
higher levels results in a complementary relationship between volcanic and hypabyssal and plutonic rocks, or, if
magmas ascend between levels without significant differentiation, then volcanic rocks may be equivalent to
hypabyssal rocks. In the failed eruption model (e.g., Glazner et al., 2015), plutonic rocks do not feed volcanic
eruptions, or share genetic characteristics, except perhaps those attained at the magma source region. In this case,
hypabyssal rocks may represent the melt-rich cap of a pluton, or a rapidly chilled roof zone.
111
with the Tioga Pass intrusions, and detrital zircon studies that show that this type of unit is a regionally
extensive part of the arc framework in the eastern-central Sierra Nevada (e.g., Attia et al., 2018 and
references therein). Although crustal contamination was limited, based on Rb-Sr, Sm-Nd, and O isotope
systems, the marine sediments (and the underlying basement/lithospheric mantle) represent one likely
assimilant.
Isotopes from Tioga Pass samples are relatively primitive compared to isotopic compositions of
intrusive suites from the central Sierra Nevada, which are on average
87
Sr/
86
Sr i >0.706 and εNd <-1.7 (e.g.,
Kistler and Peterman, 1973; DePaolo, 1981; Kistler et al., 1986). This signal is not unique to Tioga Pass,
and other Late Cretaceous intrusions with a similar isotopic composition, emplaced south of Tioga Pass,
include the ca. 97 Ma quartz monzodiorite of Rush Creek, which is comparable to the Tioga Lake quartz
monzodiorite in composition and size (Kistler and Swanson, 1981; Bateman et al., 1984; Fig. 4.10A), and
the hypabyssal ca. 100 Ma Shellenbarger Lake granite porphyry (Lowe, 1995; Figs. 4.10A, B). To the north
of Tioga Pass, the 97 Ma Solider Lake granodiorite (Cao et al., 2015) and a Cretaceous volcanic sample are
isotopically similar to the Tioga dacite porphyry unit. The eastern margin of the Kuna Crest granodiorite
and Kuna Crest lobe, emplaced at 95-94 Ma (Memeti et al., 2010a, Memeti et al., in review) contain the
most primitive compositions of the Tuolumne Intrusive Complex (Gray et al., 2008; Fig. 4-12 in Memeti
et al., 2014; Memeti et al., in review), and represent much larger volumes of magma than the other
examples. These intrusions and volcanic rocks are all located within an arc-parallel belt in the eastern-
central Sierra Nevada that is also defined by δ
18
O zrc approximating, or slightly higher than, mantle values
between 5.5-6.5‰ (Lackey et al., 2008). Lackey et al. (2008) interpreted this as a zone of minimal crustal
contamination, and extensive recycling of mantle lithosphere and upper-mantle (see also Coleman et al.,
1992). The similarity in isotope compositions suggests each of these intrusions had a similar mantle-
derived magma source to the Tioga Pass system, including a minor component of deep marine sediments
(Fig. 4.10). It further indicates that the whole-rock Sr, Nd and Pb isotopes across this eastern belt capture
the basement composition of the arc framework rocks (e.g., Kistler, 1990).
5.2 Alternate Interpretation of the Tioga Pass magmatic complex
An alternate interpretation of the field and geochronologic data is that the porphyry represents the sub-
volcanic (intrusive) roots of a lava dome (e.g., Swanson et al., 1989). The lack of vesicles within this unit
suggests that the magma was already degassed, or the microstructure was overprinted by Cretaceous
regional tectonic shortening (Eichelberger et al., 1985; Bonnichsen and Kauffman, 1987; Sampson, 1987;
Cao et al., 2015). Flow banding structures are characteristic features of lava flows and domes, both in the
upper parts of magma feeder conduits and at the surface (Gonnermann and Manga, 2005; Tomek et al.,
112
2016), a feature which is ubiquitous in the porphyry unit. Glomerocrysts have been recorded in extrusive
lava flows and domes (e.g. Seaman, 2000). The sharp, discordant intrusive contact between the intrusive
units and older host rocks requires this to be a sub-volcanic system but could have fed an extrusive structure
such as a lava dome or stratovolcano (e.g., Kistler and Swanson, 1981).
5.3 Tectonic Implications
New geochronology from volcanic and sedimentary units at Tioga Pass expands our understanding of
the regional early Mesozoic stratigraphy of the Saddlebag Lake and northern Ritter Range pendants. In
combination with field relationships, this resolves differences between Triassic and Cretaceous tectonic
histories, and aids in refining the local structural history for this area. Triassic and Jurassic deposition is
recorded in the Tioga Lake and Dana sequence sections, preserving the transition from Triassic emergent
volcanism to shallow marine sedimentation. We suggest that correlative in-situ units from the Saddlebag
Lake and northern Ritter Range pendants are not observed because they were removed prior to, or during,
emplacement of the 95-85 Ma Tuolumne Intrusive Complex, which truncates the ca. 220 Ma volcaniclastic
unit at Tioga Pass. Some motion of the Tioga Lake and Dana sequence sections may have been
accommodated by the Gaylor Peak fault, as proposed by Greene (1995) and Schweickert and Lahren (1999).
If so, this must postdate Dana sequence formation (195 Ma) and precede the emplacement of the magmatic
complex at 100 Ma, which intrudes these units. Subsequent strike-slip re-activation of this fault offsets the
porphyry unit (Fig. 4.11). Rotation of bedding and foliation in the Tioga Lake block could alternatively
have occurred during the emplacement of the porphyry unit; rotated structures are also observed in <50 m
stoped host-rock blocks within the porphyry unit.
Repeated episodes of deformation and tilting of strata occurred in the Triassic and Jurassic (e.g.,
Paterson and Memeti, 2014; Cao et al., 2015), however, Cretaceous units of the Tioga Pass intrusive
complex are less tilted and less strained than the older rocks in the same area. This may indicate that by 100
Ma (at the peak of the Cretaceous flare-up), tectonically driven steepening and shortening was waning. In
intrusive rocks, the porphyry preserves evidence for ductile shear during Sierra Crest Shear Zone activity
that occurred after porphyry emplacement at ca. 100 Ma, but prior to brittle faulting between 84-80 Ma
(e.g., Hartman et al., 2018).
The Kuna Crest granodiorite, Tioga Lake quartz monzodiorite, hypabyssal porphyry and Cretaceous
volcanic rocks are structurally juxtaposed at the present-day surface (Figs. 4.1, 4.3A-C). This tectonic
association between plutons and volcanic rocks of similar age has been identified at other localities in the
central Sierra Nevada (e.g., Anderson et al., 2008; Greene et al., 2017) and is attributed to downward
113
transfer of host rock (e.g., Tobisch et al., 2000; Paterson and Farris, 2008; Memeti et al., 2010b; Cao et al.,
2016). Using the minimum and maximum U-Pb zircon ages for each unit, and Al-in hornblende barometry
of the Kuna Crest granodiorite (Ague and Brimhall, 1988; Memeti et al., in review), the rate of downward
transfer is estimated between 1.6–4 km/m.y., consistent with pendant-wide estimates from Cao et al. (2016).
Whether downward transfer was episodic (over geologic timescales) or continuous remains to be tested.
5.4 Implications for Hypabyssal Intrusions in the central Sierra Nevada and the Volcanic-Plutonic
Connection
The Tioga Pass magmatic complex is one example of a shallowly emplaced, porphyritic magma feeder
system in the central Sierra Nevada (Fig. 4.11). Other Cretaceous hypabyssal intrusions scattered along a
~50-kilometer-wide belt of the central Sierra Nevada are recognized from a synthesis of published literature
and our ongoing field studies using the field and textural criteria outlined in this study. Known hypabyssal
intrusions are labelled in Figure 4.1 and are reported as far south as Kings Canyon National Park in the
southern Sierra Nevada (not shown; Saleeby et al., 1990; Sisson and Moore, 2013; Ryan Davis et al., 2019).
Below we outline some of the general features of these intrusions in the context of our findings at Tioga
Pass:
Hypabyssal intrusions across the central Sierra Nevada are typically small in areal extent (10 m –10 km
in diameter) and have a range of 3D shapes, including punched laccoliths (flat-topped, steep-sided
intrusions: Corry, 1988, Fig. 4.11), pipes, and porphyritic dikes (Bateman, 1992; Fiske and Tobisch, 1994;
Tobisch et al., 2000). They sometimes intrude through slightly older metavolcanic rocks; the ca. 100 Ma
Shellenbarger Lake granite porphyry cuts across the Minarets caldera fill deposit and is interpreted as a
hypabyssal resurgent dome to the Minarets caldera (Fig. 4.1; Fiske and Tobisch, 1994; Tomek et al., 2016).
Field observations of hypabyssal intrusions highlights the diversity of structures found at shallow levels
and suggests that a range of magmatic conditions that may be captured in these systems. In the central
Sierra Nevada, hypabyssal intrusions are dominantly intermediate to felsic in composition, such as andesite,
dacite, monzodiorite, granodiorite and granite. Basaltic andesite compositions are locally found (Fiske and
Tobisch, 1994; S. Attia, pers. comm). Textures are largely porphyritic, with phenocrysts of alkali feldspar,
plagioclase, and biotite common (Peck, 1980; Lowe, 1995; Peck, 2002). Phenocrysts are set in a
groundmass that ranges from medium grained (akin to typical plutonic equigranular textures) to fine-
grained or microcrystalline (volcanic aphanitic textures). In the latter case, the only differentiating feature
of a hypabyssal stock from a volcanic deposit is a discordant, intrusive contact with surrounding strata or
older plutonic bodies (Saleeby et al., 1990; Bateman, 1992). As exemplified at Tioga Pass, hypabyssal
114
bodies often have complex macro- to micro-scale internal structures. In many instances they contain
miarolitic cavities and granophyric microstructures, which are compatible with volatile saturation and
shallow emplacement at approximately 1 kbar (Candela, 1991, 1997), although miarolitic cavities are also
found in the Tuolumne Intrusive Complex (6–10 km; Memeti et al., 2014).
Compositions of the hypabyssal intrusions can be used to reconstruct source characteristics, test crystal
accumulation processes, and explore the possible genetic associations between plutons, hypabyssal
intrusions and volcanic rocks in the upper crust. Cretaceous plutons in the central Sierra Nevada are
magnesian, calc-alkaline and metaluminous to peraluminous, granodiorite and granite in composition
(Bateman, 1992). Porphyry and volcanic fields overlap considerably with plutonic compositions (Fig. 4.9).
However, the hypabyssal samples trend towards peraluminous and alkali-calcic compositions and are
generally more compositionally restricted than plutons, between 62–76 wt.% SiO
2 (e.g., Peck and Van
Kooten, 1983; Bateman et al., 1984; Lowe, 1995; Memeti, 2009). Our findings at Tioga Pass extend the
compositional range of known hypabyssal intrusions in the central Sierra Nevada (Fig. 4.9), and in some
cases results are quite distinct from regional patterns (e.g., Yttrium; Fig. 4.9). In addition to the Tioga Pass
system, other Late Cretaceous plutons and hypabyssal intrusions emplaced along the eastern edge of the
central Sierra Nevada have relatively primitive bulk isotopic compositions (Fig. 4.10A) compared to
average values for the arc section, illustrating a regional control on mantle and crustal components in the
magma source that extends into the upper crust (Kistler, 1990; Lackey et al., 2008).
Field relationships combined with chemical compositions of plutonic, hypabyssal and volcanic rocks
have been studied in other settings to characterize volcanic-plutonic systems and refine existing volcanic-
plutonic models (Fig. 4.12). For example, caldera structures that formed during the Cenozoic ignimbrite
flare-up document a wide variety of structural and chemical volcanic-plutonic relationships, made possible
due to exposures of the magma plumbing from the surface to 9 km depths (e.g., Lipman, 1984; Lipman,
2007; Bachmann et al., 2007; Watts et al., 2016; Colgan et al., 2018). Volcanic and shallow plutonic rocks
within each system typically share a ‘compositional affinity’ (e.g., Watts et al., 2016) and, in some cases,
shallow plutons are interpreted as the residual magmas to erupted volcanic deposits (e.g., Zimmerer and
McIntosh, 2013; Deering et al., 2016; Fig. 4.12-magma feeder system model). Often, the exposed shallow
plutons are resurgent into the erupted volcanic rocks, and thus did not directly feed volcanism (e.g., Steven
et al., 1974; Johnson et al., 1989; Colgan et al., 2018; Fig. 4.12-failed eruption model). In each example,
shallow intrusions (0-6 km depth) provide windows into different levels of the magmatic system through
time. Study of these intrusions is particularly important in areas where surface deposits and the uppermost
crust are largely eroded, such as the Sierra Nevada.
115
6. Conclusions
1) The Tioga Lake quartz monzodiorite and dacite-rhyolite porphyry are co-magmatic intrusive units
overlapping in age at ca. 100 Ma and are not related to a Triassic caldera system. Our favored interpretation
is that the intrusions were emplaced at shallow crustal levels and represent part of the subvolcanic magmatic
roots of once-extensive Cretaceous volcanic deposits, now partially preserved in host-rock pendants across
the central Sierra Nevada. Thus, the only known Sierran caldera structures are the Cretaceous Minarets and
Merced peak calderas.
2) Whole-rock isotopes from Tioga Pass, and nearby intrusions and volcanic rocks, along the eastern belt
of the central Sierra Nevada are sensitive to the composition of regional arc framework rocks, including the
underlying basement and mantle sources. Mantle-derived magmas along this eastern belt contain a minor
crustal component consistent with the composition of the host Paleozoic deep marine sediments. The Tioga
Lake and Dana sequence sections represent ~2 km of Triassic and Jurassic strata that is otherwise not
preserved at Tioga Pass, due to surface or magmatic erosion. Cretaceous ductile and brittle activity re-
activated earlier structures, and rapid downward flow of rocks juxtaposed volcanic, hypabyssal and plutonic
rocks.
3) Subvolcanic porphyry intrusions in the Cretaceous central Sierra Nevada are widespread and are also
documented in other parts of the arc section (e.g., Saleeby et al., 1990; Sisson and Moore, 2013). Combining
this dataset with temporally and spatially associated volcanic and plutonic rocks provides a multi-level view
of the uppermost parts of an arc crustal column, owing to either differential erosion and/or downward
transfer processes that transported volcanic rocks erupted at the surface to 6–10 km depth before intrusion
of broadly coeval plutons. In this regard, porphyry intrusions represent key structural markers in tectonic
studies, where they are constrained by a crystallization age and emplacement depth (between approximately
0–6 km). We posit that these hypabyssal intrusions are also significant in that they may physically and
chemically relate to volcanic deposits, and/or the late melts drained from plutons.
7. Acknowledgements
We acknowledge support from National Science Foundation grants EAR-1624847 awarded to Scott
Paterson, and EAR-1624854 awarded to Vali Memeti. We thank Mark Pecha and research scientists from
the Arizona LaserChron Center (NSF EAR-1649254) for help obtaining U-Pb zircon ages. We thank James
McColl, Luke Ardill, and Maria Cuevas for field assistance, Aly Angulo for help preparing XRF beads,
and Melanie Barnes and Kevin Werts for helping us obtain trace element data. Cal Barnes, Wenrong Cao,
and Snir Attia are thanked for useful discussions and comments. Snir Attia provided samples of mafic
116
hypabyssal rocks to include in our compilation. Helpful reviews by Jade Star Lackey, Kathryn Watts,
Andrew Barth, Lily Claiborne, Joseph Colgan, Nancy Riggs and Gerardo J Aguirre-Diaz improved earlier
versions of this manuscript. We thank Sarah Roeske for editorial handling.
117
Chapter 5: Schlieren-bound magmatic structures record crystal flow-
sorting in dynamic upper-crustal magma-mush chambers
This chapter is in press at Frontiers in Earth Science: Petrology
Ardill, K., Paterson S.R., Stanback, J., Alasino, P.H., King, J.J., and Crosbie, S., (in press). Schlieren-bound
magmatic structures record crystal flow-sorting in dynamic upper-crustal magma-mush chambers. Frontiers
in Earth Science, doi:10.3389/feart.2020.00190.
Abstract
The size, longevity, and mobility of upper-crustal magma mushes, and thus their ability to mix and
interact with newly arriving batches, are key factors determining the evolution of magma reservoirs.
Magmatic structures in plutons represent local sites of structural and compositional diversity and provide
an opportunity to test the extent of physical and chemical processes that operated through time. Regional
compilation of compositionally-defined magmatic structures, specifically those involving schlieren, in the
Tuolumne Intrusive Complex (TIC), yields a synthesis of ~1500 schlieren-bound structure measurements.
Field observations, petrography, and whole-rock geochemistry were integrated to test schlieren formation
mechanisms.
At a local scale (1 mm – 1 m), we find that schlieren-bound structures formed from the surrounding host
magma during dynamic magmatic processes such as crystal flow-sorting, magmatic faulting, and folding.
Fluidization of the magma mush, interpreted from 1 m – 1 km wide domains of clustered schlieren-bound
structures, appears to have operated within a hydrogranular medium, or ‘crystal slurry’ (Bergantz et al.,
2017). At the regional scale (10’s km), outward younging patterns of troughs, migrating tubes, and plumes
indicate that the mush convected, driven by intrusion of new pulses. Troughs and planar schlieren are
weakly oriented parallel to nearby major unit contacts, which could be related to internal mush convection
or effects of high thermochemical gradients at internal unit boundaries. We hypothesize that these younging
patterns and orientations have the potential to constrain the size of mobile magma mixing regions, that in
the TIC extended to a minimum of 150 km
2
(~1500 km
3
) and were long-lived (>1 m.y). These require the
generation of extensive melt-present reservoirs that could flow magmatically, formed from the
amalgamation of intruding magma pulses, and precludes dike, sill, or laccolith emplacement models. We
conclude that schlieren-bound structures are faithful recorders of the multi-scale, hypersolidus evolution of
upper-crustal magma bodies, and represent useful tools for studying plutonic systems.
118
1. Introduction
On the path to solidification, magma reservoirs reside in the crust as a crystal-rich mush (Marsh, 1981).
Current debates relate to the size and longevity of magma mushes, and the extent to which they can be
mobilized by internal (magmatic) and external (tectonic) forces (e.g., Annen et al., 2015; Paterson et al.,
2016; Bachmann and Huber, 2016; Bartley et al., 2018; Holness, 2018; Jackson et al., 2018). Each of these
factors controls the importance, and spatiotemporal extent, of physical and chemical magmatic processes
that drive differentiation within magma reservoirs (e.g., Spera and Bohrson, 2018). Thus, these factors have
implications for understanding transcrustal magmatic plumbing systems, and how material is transported
and differentiated vertically in the crust (e.g., Lipman, 2007; Cashman et al., 2017). Additionally, the size,
longevity, and mobility of magma mushes in the upper-crust is closely coupled with eruptive behavior at
volcanic centers (e.g., Bachmann and Bergantz, 2004).
Investigating the evolving behavior of magma reservoirs is a multi-faceted issue due to the multiphase
nature of magmas as well as the effects of incremental emplacement, magma-host rock interactions, and
local and regional stresses on the magmatic plumbing system. In addition, magma experiences a substantial
transition in effective viscosity during crystallization, between 10-20 orders of magnitude, emphasizing that
a wide range of behaviors are possible (e.g., Petford, 2003; Sparks et al., 2019). Early studies suggested
that magma composition and crystallinity were primary indicators of magma viscosity, as well as whether
the magma could erupt, or “lock up” (e.g., Van der Molen and Paterson, 1979, Rosenberg & Handy, 2005;
Petford, 2009). More recently, the primary focus has been on describing magmatic systems from the
mechanical perspective of hydrogranular, dense, crystal-rich slurries, with or without exsolved volatiles
(e.g., Schleicher et al., 2016; Bergantz et al., 2017; McIntire et al., 2019; Degruyter et al., 2019). This
perspective is particularly relevant for magma mushes as it incorporates the effects of particle-particle
interactions and deformation on the state of the mush, both of which are well-documented in field-based
studies of magmatic structures in plutons (e.g., Paterson et al., 2018).
The Late Cretaceous Tuolumne Intrusive Complex (TIC), in the central Sierra Nevada, California is one
example of a long-lived, composite, upper-crustal pluton, incrementally emplaced over ~10 m.y. at 6-10
km depth (Ague and Brimhall, 1988; Bateman, 1992; Coleman et al., 2004; Memeti et al., 2010; Paterson
et al., 2016). The emplacement history and evolution of the TIC continues to be intensely studied and
debated (e.g., Memeti et al., 2014; Paterson et al., 2016; Bartley et al., 2018). On one hand, the TIC has
been proposed to consist of sheets with ephemeral mush systems that have little interaction with each other
during or after emplacement (e.g., Coleman et al., 2012; Bartley et al., 2018). In contrast, other studies have
proposed that the TIC formed several long-lived (i.e., 0.5 – 2 m.y.) magma chambers, regions of
119
interconnected magma mush that experienced convection and return flow (e.g., Bateman, 1992; Burgess
and Miller, 2008; Solgadi and Sawyer, 2008; Memeti et al., 2010; Paterson et al., 2011, 2016). Each model
has distinct implications for the origin of diverse magma compositions, the thermal history of the complex,
the expected volume of eruptible magma, as well as magma rheology and the structural evolution of the
complex.
The TIC contains an unusually wide range of magmatic structures; studying them can start to address
issues surrounding the magmatic history of the complex. Compositionally defined magmatic structures, and
specifically schlieren-bound structures, are the focus of this study (Cloos, 1936; Bateman, 1992; Reid et
al., 1993; Žák et al., 2007; Solgadi and Sawyer, 2008; Paterson, 2009; Memeti et al., 2014; Paterson et al.,
2018). In plutonic systems generally, the processes driving compositionally-defined magmatic structure
formation and consequent implications for magma rheology remain controversial, with contrasting models
that rely on either physical (active) or chemical (static) mechanisms (e.g., Weinberg et al., 2001; Paterson,
2009; Hodge et al., 2012; McBirney and Noyes, 1979; Boudreau, 2011). Thus, understanding the formation
and significance of magmatic structures is an important step towards understanding the physical and
chemical evolution of upper-crustal magma reservoirs.
This paper examines compositionally defined magmatic structures in the TIC from field, structural, and
geochemical perspectives to investigate the evolution and dynamics of an upper crustal magma storage
region at scales ranging from individual structures to pluton-wide dimensions. We find that schlieren-bound
structures formed by physical flow-sorting in an active crystal-mush environment, triggered by local flow
instabilities. Dense minerals were selectively accumulated, to varying degrees, along both steeply- and
gently-dipping boundaries. Schlieren-bound structures may form highly intricate and disorganized patterns
at the outcrop scale, reflecting the complexity of crystal-melt and crystal-crystal interactions in a
hydrogranular medium (e.g., Bergantz et al., 2017). Within mappable domains, the clustering of structures
reflects the spatial and temporal heterogeneity of the magma mush, where fluidized (more mobile crystal
mush) regions promote structure formation and preservation, and less dynamic zones are devoid of
schlieren-bound structures. At the regional scale, schlieren-bound structures show weak to moderate
alignment with contacts and outward younging directions that we propose are the result of internal return
flow and convection of the mush. We use these patterns to estimate the maximum and minimum sizes of
active magma chambers.
120
2. Background
2.1 Compositionally defined magmatic structure models
Magmatic structures may be broadly grouped by the preferred alignment of objects (e.g., fabrics),
deformation related structures (e.g., folds and faults), structures of growth, younging, or kinematics (e.g.,
cross-cutting boundaries, S-C structures), and structures involving internal contacts and layering (e.g., dikes
and hybrid zones) (e.g., Paterson et al., 2018). A fifth group includes compositionally-defined magmatic
structures. For example, the heterogeneous clustering of mineral populations, such as K-feldspar megacryst
clusters (Vernon and Paterson, 2006; Memeti et al., 2014; Rocher et al., 2018), plagioclase clusters
(Oppenheim et al., in review), and schlieren (Barrière 1981; Reid et al., 1993).
The formation of compositionally-defined magmatic structures in mafic magmatic systems has been the
focus of studies for decades (e.g., reviews by Namur et al., 2015; O’Driscoll and VanTongeren, 2017,
Holness et al., 2017). For example, the Skaergaard intrusion preserves spectacular rhythmic layering as
well as channel-shaped trough structures (e.g., Wager and Brown, 1968; Irvine et al., 1998; Vukmanovic
et al., 2018). Fewer studies in silicic magmatic systems, the focus of this study, are of great interest as they
demonstrate that compositionally-defined magmatic structures may form over a wide compositional and
rheological range (e.g., Wiebe and Collins, 1998; Weinberg et al., 2001; Paterson, 2009).
One subset of compositionally-defined magmatic structures involves the formation of schlieren.
Schlieren are igneous modally defined layers that may contain abundant hornblende, biotite, magnetite,
apatite, titanite, and zircon, amongst other accessory minerals. They may be modally- or grain-size-graded.
While observed in granitoid plutons worldwide, a few localities are noted for their rich abundance, and
variety, of schlieren-bound magmatic structures: the Ploumanac’h massif, Brittany (Barrière, 1981), the
Tavares pluton, Brazil (Weinberg et al., 2001), the Halifax pluton, Nova Scotia (Smith, 1974; Clarke, 2003;
Clarke et al., 2013), the Vinalhaven granite, Maine (Wiebe and Collins, 1998; Wiebe et al., 2007) and the
Tuolumne Intrusive Complex, California (e.g., Cloos, 1936; Bateman, 1992; Reid et al., 1993; Paterson,
2009).
Schlieren are commonly found in planar form (e.g., Bateman, 1992; Žák and Klomínský, 2007; Burgess
and Miller, 2008; Pinotti et al., 2016), but also delineate curved boundaries of structures such as: channel-
shaped magmatic troughs (e.g., Wahrhaftig, 1979; Barrière, 1981; Solgadi and Sawyer, 2008; Žák and
Paterson, 2010; Alasino et al. 2019; see also Wager and Brown, 1967, Vukmanovic et al., 2018 for mafic
systems), stationary and migrating tubes, also called ladder dikes (e.g., Reid et al., 1993; Weinberg et al.,
121
2001; Wiebe et al., 2007; Dietl et al., 2010; Clarke et al., 2013), meter-scale diapirs and plume heads (e.g.,
Weinberg et al., 2001; Paterson, 2009) and mafic ellipsoids (e.g., Memeti et al., 2014).
In general, studies of schlieren-bound magmatic structures aim to address two related questions: 1) under
what magmatic conditions and by which process(es) do schlieren form (Fig. 5.1a); and 2) how do these
processes influence, or contribute to, the variety of schlieren-bound structures of different geometries and
characteristics (Fig. 5.1b)?
Models describing schlieren formation generally fall within the range of three end-member processes,
or ‘schools of thought’ (Fig. 5.1a; see also Solgadi and Sawyer, 2008; Barbey, 2009). The first considers
schlieren as a product of incomplete mixing of new (mafic) magma injections. One example of this process
would be the disaggregation of mafic magmatic enclaves into layers (e.g., Barbey et al., 2008; Farner et al.,
2018). The second school of thought calls upon in-situ thermochemical diffusion and recrystallization as
principal mechanisms to generate modal layering in schlieren. Processes such as oscillatory nucleation and
Ostwald ripening are important in these models (lower left Fig. 5.1a; e.g., McBirney and Noyes, 1979;
Glazner, 2014). The third school of thought views schlieren as resulting from magmatic flow against a more
rigid boundary, in its most dynamic form by magmatic currents or avalanching, to sort crystals by size
and/or density (e.g., Irvine et al., 1998; Solgadi and Sawyer, 2008; Alasino et al., 2019; Fig. 5.1a).
Under this framework, a range of schlieren-forming processes have been proposed that involve different
combinations of these three end-members (Fig. 5.1a). These processes include thermal and compositional
convective instabilities that could result from intrusion of new pulses or thermochemical gradients within
a solidification front (Griffiths, 1986; Weinberg et al., 2001; Pons et al., 2006; Wiebe et al., 2007; Paterson,
2009; Rocher et al., 2018; Alasino et al., 2019), shear/hydrodynamic flow sorting (Cloos, 1936; Bhattacharji
and Smith, 1964; Barrière, 1981; Pitcher, 1997; Clarke et al., 2013) together with compaction and filter
pressing at rheologic boundaries to extract melts (Weinberg et al., 2001; Žák et al., 2007; Paterson, 2009),
gravitational settling (e.g., Coats, 1936; Wager and Deer, 1939), and liquid immiscibility (e.g., Glazner et
al., 2012).
Physical processes have been proposed to create the wide variety of schlieren-bound structures; structure
types have been distinguished and explained by variations in magma flow geometry and intensity, magma
rheology, and the importance of thermal and compositional gradients (Fig. 5.1b) (e.g., Weinberg et al.,
2001; Paterson, 2009; Wiebe et al., 2017). These variations suggest that planar schlieren could form
horizontally in a largely static environment, with repeated filter pressing and/or compaction. Compaction
would produce a distinct microstructural record of viscous deformation (Holness et al., 2017). However,
122
Figure 5.1: Summary of formation mechanisms for schlieren and different magmatic structures. (A) Layer
formation mechanisms. Processes listed in the box can occur in the field inside the triangle and represent
combinations of end-member processes (bold). (B) Structure formation mechanisms. A summary of proposed
formation mechanisms and how different schlieren-bound structures have been interpreted. Schematic diagrams
show the approximate 3D geometries of each structure. Grey arrows represent interpreted motion/flow directions.
Red arrows represent younging or migration directions.
123
this is not observed in Skaergaard troughs (Vukmanovic et al., 2018). Where layers are steeply dipping,
graded, and deformed, magmatic flow conditions are considered to be more dynamic (Fig. 5.1b) (Paterson,
2009). For example, troughs represent magmatic erosion and boundary flow processes within a channel-
shaped, basal mushy surface where minerals are deposited and flow is channelized, as shown by magmatic
hornblende lineations plunging down-axis in trough basal surfaces (Solgadi and Sawyer, 2008; Paterson,
2009; Vukmanovic et al., 2018). Where relative buoyancy and density effects are significant in the magma
chamber, structures such as tubes and plumes may be favored (Fig. 5.1b) (e.g., Wiebe et al., 2007; Paterson,
2009; Clarke et al., 2013). Models to describe tube formation suggest that they represent vertical magma
(or volatile) pathways (e.g., Weinberg et al., 2001; Paterson, 2009; Dietl et al., 2010; Clarke et al., 2013),
are a record of sinking objects (e.g., Wiebe et al., 2007; Clarke et al., 2013) or are a package of stacked
trough channels (e.g., Wiebe et al., 2017). Plume heads represent examples of (now variably-oriented)
Rayleigh-Taylor instabilities in their dome-like heads, detached from the surrounding host (Griffiths, 1986;
Paterson, 2009). The rheology of the surrounding host magma must play a role in forming and deforming
structures (e.g., Wiebe and Collins, 1998; Vernon and Paterson, 2006; Hodge et al., 2012), but it is unclear
if it controls the types of structures that can form.
2.2 The Tuolumne Intrusive Complex
The Tuolumne Intrusive Complex (TIC) is a ~1100 km
2
partially nested, composite intrusive complex,
emplaced over ~10 m.y. (Calkins, 1930; Bateman and Chappell, 1979; Bateman, 1992; Coleman et al.,
2004; Memeti et al., 2010; Paterson et al., 2016). From oldest to youngest, major units are: the 95-92 Ma
Kuna Crest Granodiorite (KC; Fig. 5.2), the 92-90 Ma equigranular Half Dome Granodiorite (eHD; Fig.
5.2), the 90-88 Ma porphyritic Half Dome Granodiorite (pHD; Fig. 5.2), the 88-85 Ma Cathedral Peak
Granodiorite (CP; Fig. 5.2). Sheet-like to irregularly-shaped leucogranite bodies are found throughout all
units, and the largest known is the Johnson Granite porphyry (JP; Fig. 5.2) (Kistler and Fleck, 1994;
Coleman et al., 2004; Bracciali et al., 2008; Memeti et al., 2010, 2014; Paterson et al., 2016). Memeti et al.
(2014) published a zircon isochron map of the TIC, using CA-ID-TIMS zircon ages and structural
observations, that illustrates the spatial relationships in zircon crystallization ages. Regions of steep age
gradients, representing large age ranges over a short spatial distance, include the western TIC margin, and
the Kuna Crest lobe (Fig. 4-3 in Memeti et al., 2014). These regions are associated with internal contacts,
involving magmatic erosion and recycling along boundaries (Paterson et al., 2016). In contrast, the inner
portion of the CP unit contains larger regions of similar zircon ages, depicted by shallow gradients, and
indicates a general younging towards the north (Memeti et al., 2010; Fig. 4-3 in Memeti et al., 2014).
124
Figure 5.2: Synthesis of magmatic structures in the Tuolumne Intrusive Complex. (A) Magmatic fabric summary. Foliation data are summarized by
trend lines. Lineations are summarized on the stereonet inset. Inset map shows the location of the Tuolumne Intrusive Complex. (B). Map summary of
schlieren-bound structural data. Points represent the locations of schlieren-bound structures. Zones of concentrated structures are highlighted in blue,
and observations of younging/migration are labelled with arrows. Unit abbreviations as follows: KC= Kuna Crest granodiorite; eHD= equigranular
Half Dome granodiorite; pHD= porphyritic Half Dome granodiorite; CP=Cathedral Peak Granodiorite; JP=Johnson Granite Porphyry. The locations of
the four mapped domains are labelled. GA=Glen Aulin; YL= Young Lakes; TP=Tenaya Peak; LC=Lyell Canyon. Data sources for map and structural
data: Bateman et al., (1988); Chesterman (1975); Loetterle (2004); Burgess and Miller, (2008); Paterson et al. (2008); Solgadi and Sawyer (2008);
Memeti et al. (2010, 2014); Žák et al., (2009); Paterson et al. (2016); Oppenheim et al. (in review). We acknowledge Vali Memeti, Bob Miller, and Jiri
Žák for sharing field data for this synthesis.
125
Major and minor boundaries within the TIC are defined by a range of internal contact types (Memeti et
al., 2010, 2014; Paterson et al., 2016). These include sharp and gradational contacts, and contacts defined
by a change in composition and/or microstructure (Memeti et al., 2014; Paterson et al., 2016). Contacts are
variable both along and across strike; for example, at Glen Aulin (GA in Fig. 5.2), the boundary between
eHD and pHD is characteristically gradational over 10-25 m, while the younger pHD- CP contact alternates
between a gradational and sharp contact along strike at the 100 meter scale (Paterson et al., 2016). Contacts
in some cases are defined by compositionally defined magmatic structures and may have contact-parallel
magmatic or solid-state fabrics (Žák et al., 2007; Paterson et al., 2008; Žák et al., 2009). They are interpreted
to record information about magma mingling and mixing, erosion, recycling, and preserve information
about thermal or rheological magmatic conditions (Memeti et al., 2010; Paterson et al., 2016).
The complex is generally more mafic in older units and more felsic in younger units; however, every
unit is highly composite, and a variety of rock types occur in each unit (Bateman, 1992; Memeti et al., 2014;
Paterson et al., 2016). Four lobes at the margins of the TIC are each normally zoned (Memeti et al., 2010).
KC rocks are both the most mafic and most varied in composition, from generally equigranular, minor
gabbro and diorite to tonalite, granodiorite, granite, and leucogranite. The eHD and pHD are granodiorite
units containing euhedral cm-size hornblende and biotite, and the addition of < 3 cm K-feldspar phenocrysts
in the pHD (Bateman and Chappell, 1979; Bateman, 1992; Memeti et al., 2014). CP rocks are dominantly
granodiorite to granite with K-feldspar megacrysts 2-12 cm forming the porphyritic texture (Bateman,
1992; Memeti et al., 2014). The most evolved, peraluminous rocks in the TIC are found in the youngest,
NE corner of the CP unit (Memeti et al., 2010).
TIC rocks are metaluminous to locally peraluminous, calc-alkaline compositions, spanning a wide range
of SiO
2 contents (48-79 wt.%) (Bateman and Chappell, 1979; Bateman, 1992; Memeti et al., 2014).
Hornblende analyses indicate that all units experienced crystal accumulation and/or melt loss to varying
degrees, contributing to the wide variation in rock type (Barnes et al., 2016c; Werts et al., 2020). The
considerable compositional overlap between units in major and trace elements, as well as isotopes, has been
attributed to a combination of fractionation and mixing processes, involving multiple magma sources; the
depth that these processes occur is debated (e.g., Kistler et al., 1986; Bateman, 1992; Gray et al., 2008;
Coleman et al., 2012; Memeti et al., 2014; Barnes et al., 2016c).
Mapping by Cloos (1936), Bateman et al. (1983), Žák and Paterson (2005), and Žák et al., (2007)
revealed multiple, time-transgressive magmatic fabrics within the TIC, defined by the alignment of
hornblende, biotite, plagioclase, and quartz and K-feldspar to a lesser extent. Four magmatic fabrics were
distinguished: Type 1 fabric is associated with local compositionally defined structures (e.g., schlieren) and
126
is generally discordant to, or overprinted by, additional regionally extensive fabrics. Type 2 foliations,
approximately margin parallel, but in detail at an angle to internal contacts, are typically the earliest formed
(Žák et al., 2007). While type 3 (NW-SE-striking) and 4 (E-W-striking) fabrics may show contradictory or
ambiguous overprinting relationships to each other at a single outcrop, Žák et al. (2007) concluded that the
type 4 fabric is generally younger than type 3. Magmatic lineation plunges steeply in all cases (Žák et al.,
2007).
2.3 Schlieren-bound magmatic structures in the Tuolumne Intrusive Complex
The earliest studies of schlieren-bound magmatic structures in the TIC by Cloos (1936), Bateman
(1992), and Reid et al. (1993) documented field characteristics of planar schlieren, troughs, and tubes
(referred to as ladder dikes). These included planar to concave upward schlieren geometries with horizonal
to sub-vertical dips, reverse (size) graded layering, and cross-bedding. Schlieren were proposed to form by
either gravitational settling of crystals (e.g., Coats, 1936) or crystal flow-sorting (e.g., Cloos, 1936;
Bhattacharji and Smith, 1964). Reid et al. (1993) suggested that the schlieren bounding tubes “settled from
crystal slushes flowing through narrow vertical channels”. Repeated events of schlieren deposition and
erosion, and gradients in magma flow intensity were proposed to explain field observations of the layers
(Bateman, 1992; Reid et al., 1993). Bateman (1992) first observed the map-scale pattern of outward-dipping
planar schlieren towards older, outer units and interpreted this pattern to result from magma tilting and
shearing on the partly-solid margins of the magma chamber, driven by emplacement of new magma at the
core of the complex.
Reid et al. (1993) noted that migrating tubes in the TIC are spatially clustered, and that the minerals in
schlieren defining tubes, while highly concentrated, matched the assemblage of the surrounding host.
Whole-rock major and trace element data supported interpretations that schlieren are cumulate rocks
derived from main-sequence TIC magmas, with some derived from more mafic magmas. Notably, schlieren
are depleted in Al
2O 3, and other elements necessary to form plagioclase, relative to the main-sequence TIC
rocks (Reid et al., 1993). Because of this, crystal accumulation associated with schlieren formation was not
considered a viable mechanism to generate the compositional variety of main-sequence TIC rocks.
Subsequent studies investigated the interaction between structure formation and local- to regional-scale
magmatic deformation. Loetterle and Bergantz (2003) interpreted deformation-related structures such as
schlieren slumping to represent loading of a crystal mush at low melt fractions between 20-50%. Žák and
Paterson (2005) mapped sheeted zones in the TIC and found outward younging planar schlieren, troughs,
and tubes that formed during and after assembly of the sheets. These structures were subsequently deformed
by magmatic faults and folds and overprinted by magmatic fabrics (Žák and Paterson, 2005). Domains with
127
the most complex structural and petrologic relationships were found at irregularly-shaped contacts, such as
corners (Žák and Paterson, 2005).
Paterson et al. (2008) interpreted magmatic structures within the Sawmill Canyon sheeted zone as local
thermal-mechanical instabilities in crystal-rich magmas. Schlieren are enriched by 3-4-fold in the high-
field-strength elements (HFSE) relative to the surrounding host, but K, Na, Sr, and Rb remain constant,
consistent with the enrichment and depletion trends identified by Reid et al. (1993). Least-squares
modelling of whole-rock data (Albarède, 1995) required ~30% accumulation of the constituent minerals to
form schlieren. Leveraging zircon saturation temperatures calculated for TIC magmas (Miller et al., 2007),
the authors proposed schlieren formation to occur between ~760°C and the solidus. The structures are
locally intensely magmatically deformed, but also record the regional strain by overprinted magmatic
fabric, which likely became effective when the local magma flow field waned (Paterson et al. 2008). The
authors emphasized the importance of rheologic flow and stress gradients over gravitational forces alone in
forming steeply-dipping schlieren.
The Sawmill Canyon structures were also studied from a microstructural and chemical perspective by
Solgadi and Sawyer (2008). Within planar schlieren and trough structures, they identified multiple
hornblende populations corresponding to the major host rock units in the area (KC, HD and CP units), and
thus interpreted schlieren to represent crystal mixtures. Hornblende did not show intra-layer compositional
gradients, and thus did not form from a gradually evolving melt composition or in-situ fractionation
(Solgadi and Sawyer, 2008). All minerals within schlieren gradually were rotated away from basal, layer-
parallel orientations towards the top of the layer. These observations together with field observations of
erosional features led to the interpretation that dense minerals were mobilized and segregated into schlieren
by high-energy “hyper-concentrated (magmatic) sedimentary gravity flows”. As the flow waned, crystals
were deposited, and as crystal separation became less efficient, graded layers formed. Changing flow
intensities or conditions could explain different types of schlieren layers and structures.
Burgess and Miller (2008) identified a densely-concentrated, heterogeneously distributed 1-2 km wide
zone of planar schlieren and tubes near Tenaya Lake. The majority of planar schlieren are oriented at low-
angles to the adjacent contact between pHD and CP. They viewed schlieren forming in areas of strong
rheological contrasts (along boundaries), and disrupted schlieren (faulted, folded, re-intruded) as evidence
for a destabilized crystal mush. In addition to whole-rock elemental trends comparable with earlier schlieren
studies (Reid et al., 1993; Paterson et al., 2008; Solgadi and Sawyer, 2008), whole-rock Sm-Nd and Rb-Sr
isotope compositions between a schlieren sample and host sample were indistinguishable in ƐNd
i, but the
128
schlieren had slightly lower
87
Sr/
86
Sr i values than the host CP. Their interpretations focused on the mixing
processes that generated the relatively narrow range of isotope compositions in the CP unit.
Paterson (2009) described the key characteristics of several TIC schlieren-bound structures, including
the addition of plume heads, diapirs, and pipes, and outlined their structure-specific dynamic features.
Paterson (2009) noted that structures were clustered in space across each unit of the TIC, with certain
patterns: Troughs young outwards, potentially favoring low-stress sites, and tube structures are statistically
vertically plunging, and thus may be useful paleo-vertical indicators. The implications of this study were
that this group of structures together preserve a time-transgressive history of magma mobility in a crystal-
rich, already-constructed magma chamber. Future, open questions concerned the spatial clustering of
structures, and whether schlieren-bound structures might record tectonic information.
Žák et al. (2009) modeled whole-rock schlieren compositions in the Mammoth Peak sheeted zone to
show that schlieren were enriched in major and trace elements from ~150% (large-ion-lithophile elements,
or LILE) to >500% (MREE, Fe, Mg, Mn) over the adjacent host granitoids. Some minerals in schlieren
form glomerocrysts (e.g., magnetite), separated by interstitial quartz and K-feldspar. The clots (if formed
prior to structure formation) have a greater density than the individual minerals and thus may enhance the
effects of gravitational settling and flow-sorting processes (Žák et al., 2009). Žák and Paterson (2010)
further explored the dynamic interplay between regional tectonics and magmatic processes at magmatic
internal contacts, showing that the magmatic tubes were statistically elongated parallel to the regional NW-
SE magmatic foliation, indicating that the local-flow structures were experiencing the regional stress when
they formed.
Hodge et al. (2012) explored the spatial clustering of tube structures at Glen Aulin, finding regular
spacing between clusters of tubes of 1-10 m. In this area, tubes were statistically oriented perpendicular to
the nearby CP-pHD contact. Using detailed maps and structural measurements, they proposed that tube
structures were buoyant magma plumes that were oriented and deformed by rheological contrasts between
the buoyant plume and the host magma. They defined regimes in which structures were most likely to be
deformed (ponding or tension) and estimated a yield strength of the magma on the order of 10
3
Pa.
In contrast to previous studies that view schlieren-bound structures originating from physical, flow-
dominated processes, Glazner et al (2012) and Bartley et al (2013) proposed a chemically driven mechanism
for schlieren formation in migrating tube structures. In this model, the felsic inner core represents a conduit
for late, Si-rich immiscible liquids, while the dense, Fe-rich liquid is deposited at the tube base, forming
129
schlieren. These studies do not address the field observations of concentric magmatic fabrics in the schlieren
walls of the tube, as well as processes that could produce the migrating pattern.
Memeti et al. (2014) reviewed the wide array of TIC magmatic structures in the context of regional
geochronology and petrology of the main-sequence rocks. They mapped a regionally-extensive planar
schlieren zone in the north of the CP unit (Fig. 5.2). U-Pb zircon thermal ionization mass spectrometry
dates from zircons in schlieren and surrounding host overlapped, and thus were interpreted to be locally
derived. Paterson et al. (2016) expanded these datasets and outlined several lines of evidence for magmatic
erosion and recycling in the TIC main sequence rocks, that included schlieren-bound structures.
Extrapolating from the known occurrence and approximate density of structures, they estimated that the
TIC contains >9000 schlieren-bound magmatic structures with local erosional features.
Wiebe et al. (2017) proposed a model for the formation of migrating tube structures involving
progressive downward flow and stacking of schlieren-bounded channels (akin to trough channels). They
suggest that tubes migrate outwards (towards older units) when new magma pulses are intruded in the center
of the complex (e.g., Bateman, 1992).
Bartley et al. (2018) interpreted schlieren layers at Tenaya Lake, which grade towards older units in the
west, as one part of “kilometer-scale lithologic cycles” where regional E-W extension makes space for the
emplacement of N-S striking dikes. In this model, the schlieren (representing early crystallizing
components) concentrate in the lower part of the dike to form modal, graded layers. This study does not
address the spatially variable orientations of schlieren, deformation of schlieren, more complex geometries
observed in the same area as planar, graded schlieren (e.g., tubes-Burgess and Miller, 2008), as well as the
lack of graded layering in other parts of the Tenaya Lake area (e.g., the host equigranular Half Dome unit).
Building on previous studies of schlieren-bound magmatic structures in the TIC, this study will focus
on characterizing and quantifying observations of planar schlieren, troughs, tubes and plumes at multiple
spatial scales.
3. Methods
3.1 Field data
3.1.1 Regional compilations
Ongoing regional compilations of TIC datasets are summarized here in order to provide context to
schlieren-bound structure data. Existing regional mapping (>1:24,000 scale) and new 1:10,000 scale maps
130
of the TIC, including information about the nature of contacts from previous studies (Žák and Paterson,
2005; Paterson et al., 2008; Žák et al., 2009; Memeti et al., 2010, 2014; Paterson et al., 2016), were compiled
and digitized using ESRI ArcGIS 10.6.1 (for a list of map sources, see Fig. 5.2). The compilation of TIC
fabric data is the result of >50 years of field mapping by the USGS (Bateman 1983; Chesterman, 1975;
Graham, 2012) and others (Žák and Paterson, 2005; Žák et al., 2007, 2009; Paterson et al., 2008; Burgess
and Miller, 2008; Memeti et al., 2010, and unpublished). Fabric data are discussed in more detail in a
subsequent paper (Paterson et al., forthcoming). Compositionally defined magmatic structures, such as K-
feldspar megacryst clusters, are reported widely in the TIC, however they do not form part of the digital
database. Schlieren-bound magmatic structure data were digitized from the same sources as the fabric data,
and assigned locations, orientations and defining characteristics (e.g., structure type), where available. In
legacy data (pre- 2009), the schlieren-structure type was not always specified (e.g., planar schlieren, trough,
tube, plume). In these cases, strike and dip measurements were assigned as planar schlieren by default.
3.1.2 Mapping domains
Mapping in four domains at 1:5,000 to 1:10,000 scale was performed over four field seasons (see Fig.
5.2 for domain locations). Domains were selected in a symmetrical pattern around the center of the southern
CP region in locations that were variably proximal to contacts with the older pHD unit (e.g., domains are
adjacent to, traversed across, or distanced from, the major unit contact) to encompass a range of magmatic
conditions. Domains were mapped in order to provide information on the distribution of magmatic
structures across areas 2-12 km
2
in size. Mapping focused on characterizing the location, type, and
orientation of schlieren-bound structures within each domain, and documenting their relationships to
lithology, magmatic fabrics, and contacts. Structures studied include planar schlieren, troughs, tubes
(stationary and migrating) and plumes in each domain. Less emphasis was placed on mafic ellipsoids, K-
feldspar clusters and K-feldspar pipes; the location of these structures was noted in the field.
3.1.3 Grid mapping
Individual structures or outcrops were selected for grid mapping at 1:10-1:100 scale to highlight key
field relationships or as a tool to unravel complicated structure. Each grid square was drawn out with a tape
measure and string. Measurements were taken of magmatic fabrics and all schlieren-bound structures in
each square. Photographs aided in digitizing the map.
3.1.4 Taking structural measurements and determining younging/migration directions
Measurements of the strike, dip and intensity of the magmatic fabric(s) were made at approximately
regular intervals across domains. When recording information about schlieren-bound structures
131
specifically, a protocol was implemented as follows: 1) record structure type and approximate dimensions;
2) measure strike and dip of planar features (e.g., planar schlieren strike and dip); 3) measure trend and
plunge of linear features in the structure (e.g., trough or tube axis); 4) measure younging or migration
directions (see below and Fig. 5.1b); and 5) note additional relative timing indicators if observed, such as
overprinting relationships of fabrics, or relationships between structures.
Planar schlieren younging directions are measured if they are reverse graded; the younging azimuth is
towards coarser grained, less modally concentrated compositions, perpendicular to the strike of the layer
(Fig. 5.1b; Fig. 5.3a). Troughs show evidence of layer erosion and truncation of older troughs. These cross-
cutting relationships define a relative younging direction that is measured perpendicular to the trough axis
(Fig. 5.1b; Fig. 5.3b-d). A strike and dip measurement can also be taken at this point to constrain 3D
geometry. In stationary tubes, usually the younging is inward, and so a migration direction is not measured
(Fig. 5.4a). Sharp erosional contacts between crescent-shaped schlieren rings define the migration direction
of a (migrating) tube structure (Fig. 5.1b; Fig. 5.4b-d). To measure plume head migration, azimuthal
younging is measured in the direction perpendicular, and towards the convex side of the plume head (Fig.
5.1b; Fig 5.5a-d). Measurements are made in 3D where possible, but working on glacially polished surfaces,
resulted in some measurements in 2D (e.g., azimuthal younging/migration directions).
3.1.5 Statistical analysis of structural measurements from domains
Statistical analysis of structural measurements (strike, dip, younging or migration directions) was
performed for each type of structure (foliation, planar schlieren and troughs, tubes, and plumes) in each
domain. Structural data was separated by unit before analysis. Stereonets and rose diagrams were created
using Stereonet 10 software by Richard W. Allmendinger, and mean vectors and cylindrical best fits of the
data within the program were used to estimate average orientations.
Angular difference was calculated as [foliation strike – schlieren strike] (planar schlieren and troughs
only) at each field station. The angular difference was converted to a 0-90° scale (i.e., ignoring dip
direction). In cases where no fabric measurement was taken, the fabric strike from the adjacent station was
used. For stations with two or more fabrics, the angular difference was calculated for each foliation
measurement at that station. The same applies for multiple schlieren strike measurements at one station.
Foliation measurements were assigned to a Type (2, 3, or 4) based on their orientation, and following the
protocol used in the regional scale map.
132
Figure 5.3: Field photos of planar schlieren and trough structures. (A) Modally and grain-size graded schlieren
layering in eHD, Sawmill Canyon area. Layers have a sharp, fine-grained base, and increase in grain size while
reducing the amount of hornblende and biotite towards the right of the photo before the next layer forms. Pen for
scale. (B) Erosion and re-deposition between meter-scale schlieren trough packages in CP. The younger, upper
layers truncate older, lower layers. The package is reintruded by leucogranite dikes. Height of outcrop imaged ~5 m.
(C) Megacryst-bearing schlieren troughs showing low-angle truncation. Photo courtesy of J. McColl. (D) Complex
schlieren showing modal/grain-size heterogeneity and complex schlieren trends.
133
Figure 5.4: Field photos of magmatic tubes. (A) Stationary tube truncated by host CP in the Glen Aulin area. This
tube contains at least 6 rings and unusually, the K-feldspar megacrysts within the structure are larger than the
surrounding host megacrysts and also oriented at high angle to the schlieren rings. (B) Bifurcated migrating tube in
the Glen Aulin area. This steeply-dipping tube is ~10 m in length, younging towards the field mapper. The diameter
of the tube varies in size as it migrates, and in places is re-intruded by host CP. Photo from Paterson et al. (2018).
(C) 6 m long migrating tube, migrating towards the bottom of the photo. These densely-concentrated rings contain
K-feldspar megacrysts. (D) 11 m migrating tube and 3 m migrating tube, Glen Aulin area. Segments of the longer
tube have been re-intruded by host CP magma and rotated. The K-feldspar megacrysts align with the mafic minerals
in the schlieren rings. Migration direction is towards the bottom of the photo (see also Hodge et al., 2012 for more
information on this structure).
134
Figure 5.5: Field photos of plumes. (A) Plume with a ~0.5 m diameter and thick schlieren head, Glen Aulin area.
The thickness of the schlieren varies around the margins of the plume, is sharpest at the contact with host CP, and
more diffuse towards the interior. (B) ~1 m wide plume head, Young Lakes area. As in (A), the schlieren is sharpest
at the contact with host CP and is diffuse towards the interior. The plume is re-intruded by host CP. (C) Diffuse
plume with tail preserved, Lyell Canyon area. This diffuse type of schlieren is common in the central regions of the
CP unit, (see also Pothole Dome area; Paterson, 2009). (D) Plume with well-defined head and elongate tail. The
interior of this plume has a clear compositional difference with the surrounding host CP.
135
3.2 Sample collection and analysis
Samples were collected from each of the main TIC units, from a range of schlieren-bound structure
types (sample locations provided in supplementary materials). The largest group consists of planar schlieren
and troughs, with a smaller subset of tube and plume samples. In many cases, multiple samples were
collected from a single structure, or outcrop, and divided into ‘schlieren’, ‘felsic’ and ‘adjacent host’
components. Samples were used for petrographic analysis and whole-rock geochemistry.
3.2.1 Petrography
Seven representative samples were selected for petrographic description and mineral mode estimation:
two samples each from the pHD and KC units, and three samples from the CP unit. In each unit, one
schlieren sample from a trough and an adjacent host sample were collected. In the CP unit, one additional
schlieren sample was collected from a migrating tube structure adjacent to the sampled trough. Thin section
images were traced and scanned into ImageJ software where they were subsequently processed using the
Color Threshold tool to isolate individual phases. We then used ImageJ to calculate the proportion of
mineral phases in each sample. Quartz and feldspar proportions were estimated by difference (subtracted
the mafic phases) and then estimated under an optical microscope (Table 5.1).
3.2.2 Whole rock geochemistry
A subset of 19 samples, consisting of 9 schlieren, 5 felsic components and 5 host samples, were analyzed
by ICP-MS at Activation Laboratories, or by XRF and ICP-MS at the University of Washington,
GeoAnalytical Laboratory for major, trace, and rare-earth element (REE) concentrations (samples are
labelled by method in Table 5.2). KCL major, trace, and REE sample data and materials were provided by
V. Memeti. Six of these samples, of a trough and adjacent host sample each from the CP, pHD, and KC
units (the same samples analyzed for mineralogy), were selected for isotope analysis. Four additional
samples were selected from the Pothole Dome area of schlieren and host (not paired). Samples selected for
Sr, Nd and Pb isotope analyses were analyzed at the University of Arizona. Full laboratory conditions and
methods for isotope analyses are described in Otamendi et al. (2009). Elemental and isotope data from
schlieren, felsic components, and nearby host samples in the TIC were compiled from the literature to
supplement our dataset (Data sources: Loetterle, 2004; Burgess and Miller, 2008; Paterson et al. 2008;
Solgadi and Sawyer, 2008; Memeti, 2009; Žák et al., 2009; Memeti et al., 2014).
136
Table 5.1: Modal mineralogy of representative schlieren, felsic components, and TIC host samples.
137
Table 5.2: Representative whole-rock compositions of schlieren, felsic components, and surrounding TIC host
samples. Full dataset included in the Supplementary File.
138
4. Results
4.1 Regional syntheses
4.1.1 Magmatic foliations and lineations
6485 magmatic foliation measurements and 1933 lineation measurements, covering >70% of the
intrusive complex, are summarized in Figure 5.2a. Building on the findings of Žák et al. (2007), we mapped
the distribution of Type 2-4 foliations and identified one additional N-NE striking foliation (Type 5; Fig.
5.2a). Many domains contain multiple magmatic foliations. Magmatic lineations are steeply plunging; 80%
of measurements plunge ≥60° (Fig. 5.2a).
4.1.2 Compositionally defined magmatic structures
Compositionally-defined magmatic structures, heterogeneous crystal clusters, are found in all TIC units
(Paterson, 2009). K-feldspar megacryst clusters are especially common in the CP unit, even in regions of
sparse to absent schlieren (e.g., Paterson et al., 2005; Vernon and Paterson, 2006). K-feldspar megacryst
clusters are found both inside and outside mapped domains, in zones of concentrated schlieren-bound
structures as well as in zones with sparse schlieren-bound structures (see below). They form a wide range
of shapes from irregular, to pipes and sheets (Paterson, 2009; Memeti et al., 2014).
4.1.3 Schlieren-bound magmatic structures
The compilation identified >1500 schlieren-bound magmatic structures (planar schlieren, troughs, tubes,
and plumes). Schlieren-bound structures are also widely distributed and appear to be either sparsely or
densely clustered (Paterson, 2009; Fig. 5.2b). In addition, the types of structures vary between units. Planar
schlieren and troughs, tubes, and plumes are found in all major TIC units, are less common in Kuna and
rare in late leucogranites. Tubes (both stationary and migrating) and plumes are abundant in the CP unit,
and found less commonly in the pHD, eHD, and KC units (Paterson, 2009).
Younging indicators suggest that schlieren orientations largely young or migrate outwards, towards
older contacts (~70% of the younging dataset; n=470) (see also Bateman 1992, Burgess and Miller, 2008;
Paterson et al., 2008; Solgadi and Sawyer 2008; Paterson, 2009; Žák et al., 2009). They often migrate or
young at high angles to nearby contacts (Hodge et al., 2012; Wiebe et al., 2017). However, a smaller
population (~30% of the dataset; n=202) show inward younging or contact-parallel orientations, in part
reflecting some of the heterogeneity found across smaller spatial scales (see section 4.3).
139
4.1.4 Summary
Compositionally defined magmatic structures, including schlieren-bound magmatic structures, are
widespread in the TIC and found in all units spanning the entire duration of the complex (i.e., they are times
transgressive). Schlieren-bound structures are not evenly distributed in the TIC; instead they are broadly
clustered in packages on the order of 5-15 km
2
in size, irrespective of unit type, distance to nearest contacts,
or contact orientation. They are found across both sharp and gradational contact zones. Relative younging
directions of tubes and troughs are generally outward towards contacts. Fabric type spatially varies
throughout the TIC, defining several domains with multiple, different fabrics preserved. Schlieren-bound
magmatic structures are not associated with any particular regional fabric configuration at the regional
scale.
4.2 Domain-scale structure patterns
In each of the four study areas, schlieren-bound magmatic structures are abundant (Fig. 5.2b, blue
squares). Planar schlieren and troughs represent ~60% of the structures reported in the four domains.
Stationary and migrating tubes represent 22% of the structures in the four study regions and plumes make
up 18% of the dataset. For each domain, we describe: (1) rock types, ages, and contact relationships; (2)
magmatic foliation and lineation patterns; (3) planar schlieren and trough orientations and their relationship
to fabrics; (4) tube and plume structure patterns; and (5) structure clustering.
4.2.1 Glen Aulin domain
The Glen Aulin domain is situated at the western margin of the CP unit, crossing the contacts with pHD
and eHD units (Fig. 5.2, 5.6). The CP unit at Glen Aulin is characterized by hornblende biotite granodiorite
in the west and becomes increasingly granitic eastwards towards Pothole Dome, as hornblende and biotite
decrease in abundance. Hornblende in the eastern part of the domain is replaced by biotite, except in
schlieren where it is sometimes preserved. K-feldspar megacrysts range in size from 3-10 cm and generally
increase in size to the east. K-feldspar clusters and leucogranite dikes are widely found. Contacts between
the CP and pHD units, and pHD and eHD units both strike NE-SW (azimuth: 033) in the north, and bend
to an E-W (azimuth: 090) orientation at the western margin of the study area. The contact between the CP
and pHD transitions from a sharp to gradational contact southwards along strike. The nearest U-Pb in zircon
ages range from 88.1 ± 0.2 Ma in the CP unit to 89.79 ± 0.15 Ma in the pHD (Coleman et al., 2004; Paterson
et al., 2016). There are steep age gradients across the contacts in the isochron map (Memeti et al., 2014).
Regional E-W striking (Type 4) and local NE-SW striking (Type 2) approximately margin-parallel
magmatic foliations are dominant, and the regional NW-SE fabric (Type 3) is locally and weakly expressed
140
Figure 5.6: Magmatic structures field data summary. Bold lines mark major unit contacts. Unmapped areas shown
in hatched pattern. The trend of steeply-dipping (70-90°) magmatic foliations is shown by dashed lines, depending
on fabric intensity. Magmatic lineations are steeply plunging, regardless of the strike of the foliation. The location of
different schlieren-bound structures are shown by colored dots (see key). Y or arrow symbols are oriented in the
direction of measured migration/younging. Inset box shows domain-scale magmatic foliation summary as on a
stereonet, and a summary of average orientations for younging. For Glen Aulin and Young Lakes, angular difference
measurements between the foliations and schlieren are shown as a histogram in the inset box. eHD-Equigranular
Half Dome granodiorite. pHD-Porphyritic Half Dome granodiorite; CP-Cathedral Peak granodiorite; tCP-
Transitional Cathedral Peak granodiorite. Glen Aulin area, NW quadrant. Co-ordinate system used: UTM NAD27
zone 11S.
141
in the north of the domain (Fig. 5.6). Lineations plunge steeply, with a best fit of 83235 (plunge trend;
n=18).
Planar schlieren and troughs are observed in all orientations and range from shallow to sub-vertical dips.
In the CP unit, the cylindrical best fit orientation is 212/78 (n=87) for planar schlieren and 221/76 (n=75)
for troughs, although there is a wide scatter. Trough younging (perpendicular to trough strike), similarly is
varied in orientation. The largest population of troughs in the CP unit young toward the NW, perpendicular
to, and towards, the CP-pHD contact at an average azimuth of 304 ± 24° (Fig. 5.6; 2σ uncertainty). Troughs
in the southwestern part of the CP unit, where the contact strikes E-W, have a younging direction to the
north, perpendicular to the adjacent contact. pHD and eHD planar schlieren (n=38) and troughs (n=8) trend
N-S to NE-SW, dip to the west, and young on average towards 333, perpendicular and towards older
contacts of eHD and KC, respectively. There is no clear pattern between the orientations of the regional
Type 3 and 4 foliations and schlieren orientations (Fig. 5.6 histogram). However, schlieren orientations and
the contact-parallel foliations (Type 2) are more similar, as shown by the number of schlieren with a low
angular difference to the foliation.
The average azimuthal orientation of tube migration in the CP unit is 339 ± 23° (n=113), perpendicular
and towards the contact with the pHD, a result consistent with previous studies of TIC tubes in this area
(Fig. 5.6) (e.g., Hodge et al., 2012; Wiebe et al., 2017). pHD or eHD tubes are less commonly found north
of Glen Aulin. Plume heads are abundant in the central and eastern parts of the study area, within the CP
unit (n=131). Most observations of plumes are on horizontal to shallowly dipping planes, and in 2D they
have variable migration directions, with a mean azimuth towards 293 ± 17° (Fig. 5.6). No plumes were
identified in the pHD or eHD.
Consistent with regional patterns, schlieren-bound structures at Glen Aulin are non-uniformly
distributed throughout the map area (Fig. 5.6). Schlieren-bound structures are concentrated in zones that
range in size between 0.3-1.1 km
2
. These clusters are separated by zones where schlieren-bound structures
are rare to absent, ranging in size from 0.02-0.2 km
2
. Both zones contain magmatic fabrics, and other
compositionally-defined structures such as K-feldspar megacryst clusters. To a first order, structures are
weakly clustered by type, with plumes most abundant in the eastern part of the domain (e.g., Pothole Dome),
and tubes most abundant in the western part. Planar schlieren and troughs are, in contrast, evenly distributed
within structure-rich zones.
142
4.2.2 Young Lakes domain
The Young Lakes domain is situated in the east-central part of the CP unit (Fig. 5.2, 5.7). The dominant
rock type is biotite granodiorite, with rare hornblende (altered to biotite). K-feldspar megacrysts are
between 3-6 cm in size. Mafic enclaves and K-feldspar clusters are common, and rare leucogranite dikes
are observed. Young Lakes is located 2-3 km west of the nearest contact, which is the gradational internal
contact of the CP unit with the pHD and eHD units at Sawmill Canyon, striking NW-SE (azimuth: 300).
This contact transitions to a sharp contact with the eHD southwards along strike (Paterson et al., 2008). The
nearest U-Pb in zircon age is 88.53 ± 0.12 Ma from Sawmill Canyon (Paterson et al., 2016). On the isochron
map of Memeti et al. (2014), the domain is characterized by a shallow age gradient of broadly similar ages.
Both Type 3 and 4 foliations are found at Young Lakes and appear to be equally dominant (Fig. 5.7). A
weaker NE-SW striking foliation (Type 5) is also identified locally. Lineations are steeply plunging (see
Fig. 5.3 in Paterson et al., 2008).
Planar schlieren (n=13) and troughs (n=117) have a strong preferred orientation, striking NE-SW and
dipping NW, with cylindrical best fit orientations of 009/50 and 013/41 for planar schlieren and troughs,
respectively. Troughs young predominantly towards the ESE (mean azimuth: 102 ± 17°) in the direction
of, and broadly perpendicular to, the CP-pHD contact at Sawmill Canyon (Fig. 5.7). A smaller population
(~27%) young in other orientations. The angular difference between the Type 3 and 4 foliations and
schlieren orientations is high (median between 45-60°) and each foliation type has a distinct angular
difference distribution (Fig. 5.7 histogram). Migrating and stationary tubes (n=2) as well as plumes (n=4)
are rare at Young Lakes and do not show any preferred orientation.
Structural diversity is relatively low at Young Lakes, making it difficult to estimate clustering by
structure type. However, planar schlieren and troughs are clustered in space. Zones where schlieren are rare
to absent range in size from 0.07-0.6 km
2
. These domains are smaller than the extent of zones where
schlieren-bound structures are abundant (0.6-1 km
2
). Within the structure-rich clusters, structures appear
more sparsely distributed than in Glen Aulin.
4.2.3 Tenaya Peak domain
The Tenaya Peak domain is situated in a narrow N-S trending belt of the pHD unit (Fig. 5.2, 5.8a). The
pHD unit is a biotite-hornblende granodiorite with hornblende phenocrysts and biotite books 3-8 mm in
size. K-feldspar megacrysts are 2-3 cm in size, increasing in size towards the CP unit (see also Bateman
1992; Memeti et al., 2014). The mapped domain crosses the N-S trending sharp contact of the CP-pHD in
the east (350-000; Fig. 5.8a) and is ~200 m from the N-S trending gradational contact of the pHD-eHD in
143
Figure 5.7: Magmatic structures field data summary. Bold lines mark major unit contacts. Unmapped areas shown
in hatched pattern. The trend of steeply-dipping (70-90°) magmatic foliations is shown by dashed lines, depending
on fabric intensity. Magmatic lineations are steeply plunging, regardless of the strike of the foliation. The location of
different schlieren-bound structures are shown by colored dots (see key). Y or arrow symbols are oriented in the
direction of measured migration/younging. Inset box shows domain-scale magmatic foliation summary as on a
stereonet, and a summary of average orientations for younging. For Glen Aulin and Young Lakes, angular difference
measurements between the foliations and schlieren are shown as a histogram in the inset box. eHD-Equigranular
Half Dome granodiorite. pHD-Porphyritic Half Dome granodiorite; CP-Cathedral Peak granodiorite; tCP-
Transitional Cathedral Peak granodiorite. Young Lakes area, NE quadrant. Co-ordinate system used: UTM NAD27
zone 11S.
144
Figure 5.8: Magmatic structures field data summary. Bold lines mark major unit contacts. Unmapped areas shown
in hatched pattern. The trend of steeply-dipping (70-90°) magmatic foliations is shown by dashed lines, depending
on fabric intensity. Magmatic lineations are steeply plunging, regardless of the strike of the foliation. The location of
different schlieren-bound structures are shown by colored dots (see key). Y or arrow symbols are oriented in the
direction of measured migration/younging. Inset box shows domain-scale magmatic foliation summary as on a
stereonet, and a summary of average orientations for younging. eHD-Equigranular Half Dome granodiorite. pHD-
Porphyritic Half Dome granodiorite; CP-Cathedral Peak granodiorite; tCP-Transitional Cathedral Peak granodiorite.
(a) Tenaya Peak, SW quadrant (Aldiss, 2017). (b) Lyell Canyon, SE quadrant (McColl, 2017). Co-ordinate system
used: UTM NAD27 zone 11S.
145
the west. The eastern part of the domain, in the CP unit, was previously mapped and studied by Burgess
and Miller (2008). Ages range from 87.3 ± 0.7 Ma to 89.79 ± 0.15 across the contact (Burgess and Miller,
2008; Paterson et al., 2016), represented by a smoothly varying age gradient on the zircon isochron map
(Memeti et al., 2014).
The E-W Type 4 foliation is dominant over the NW-SE Type 3 foliation (Fig. 5.8a). A N-S to NE-SW
striking foliation (Type 2) is locally well-defined in the pHD unit, but weak-to-absent in other areas.
Burgess and Miller (2008) report steeply plunging lineations between 70-90°.
Planar schlieren and troughs in the pHD unit, on average, strike parallel to the contact (165/40; n=26
schlieren, 147/44; n=35 troughs; Fig. 5.8a), although there is a wide range in orientation, with strikes
between 021-358. Trough younging orientation is variable, with a weak preference towards the west (mean
vector: 245 ± 50°). Including measurements by Burgess and Miller (2008), troughs in the CP unit (n=7)
vary in orientation between NW-SE striking to E-W striking. Two troughs measured in the adjacent CP
unit young westward, and one trough youngs towards the east. The small number of foliation and planar
schlieren/trough measurements from this domain precludes statistical analysis of angular difference.
Qualitatively, the cylindrical best fit orientations for planar schlieren and troughs aligns with the
orientations for Type 2 and Type 3 foliations.
Migrating tubes in the pHD are concentrated at the eastern margin of the pHD, adjacent to the contact
with the CP. A small number of measurements (n=13) show a preference for westward migration (mean
vector: 266 ± 26°; Fig. 5.8a), perpendicular to the nearby eHD contact. In the CP unit, including
measurements by Burgess and Miller (2008), tubes are also westward-younging (n=9; not shown). Thirteen
plumes mapped across the pHD have a widely variable orientation, with a mean vector of 284 ± 99° (Fig.
5.8a). Plumes were not found at the western margin of the CP unit.
Within this domain, the structures are not strongly clustered by type. However, tubes seem to be most
abundant close to the pHD-CP contact. Areas ranging in size from 0.1-0.4 km
2
do not contain schlieren-
bound structures. Domains that contain abundant structures are ~0.5-1 km
2
in area.
4.2.4 Lyell Canyon domain
The Lyell Canyon domain is situated in a wide gradational zone at the southeastern margin of the CP
unit, bordering the pHD unit to the east (Figs. 5.2, 5.8b) (Oppenheim et al., in review). The pHD unit here
is a porphyritic granodiorite with biotite>hornblende, and locally contains biotite books and K-feldspar
megacrysts (2-3 cm in size). In the transitional CP unit, the K-feldspar megacrysts increase in size and
number. The gradational contact strikes NE-SW (azimuth: 015). The gradational contact between the
146
transitional CP and main CP unit is found at the western edge of the mapped domain and also strikes NE-
SW. There are no age estimates for this section of the CP unit. The interpolation on the isochron map of
Memeti et al. (2014) suggests that this region has broadly similar ages (a shallow age gradient) of about 87
Ma.
Regional Type 3 and 4 foliations are observed in equal intensity and cross-cut the CP-pHD contact (Fig.
5.8b). A broadly contact-parallel, NE-SW striking foliation (Type 2) is also found in both units. Lineations
were not recorded in this domain.
The Lyell Canyon domain is the smallest of the four mapped domains, and the number of structures
observed is also the smallest. Thus, less emphasis is placed on statistically average orientations (for
comparison to other domains, average orientations are given in summary diagram Fig. 5.8b). A small
number of planar schlieren and troughs in the CP unit (n=16) vary widely in orientation, but the largest
group are NW to NE striking. A small number of graded schlieren measurements (n=5) indicate younging
eastwards towards the pHD unit and troughs also young towards the contact, between 060 and 160. In the
pHD unit, planar schlieren and troughs are both striking NW (n=37): 337/61 (n=30) and 294/79 (n=7),
respectively. Planar schlieren grading in the pHD unit indicates younging westwards (n=9), with a smaller
number younging to the NE (n=2). Trough azimuthal younging directions are between 180 and 330,
opposite to structures in the CP unit. Angular relationships between foliations and schlieren orientations
were not calculated due to the limited size of the dataset.
Tubes and plumes are found across the contact in both units. In the CP unit, tubes (n=24) and plumes
(n=11) migrate eastwards, with plumes showing a wider scatter in orientation. In the pHD unit, migration
directions are reversed, as tubes (n=12) and plumes (n=10) migrate westwards (Fig. 5.8b).
Domains where structures are rare to absent are 0.1-0.3 km
2
in area, similar to the other study areas, but
spatially restricted to the southern parts of the domain. The rest of the domain has abundant structures,
approximately evenly-spaced. With a larger dataset covering a broader area, clustering by type could be
evaluated.
4.2.5 Summary
Domain-scale mapping revealed structure clustering on the order of 0.3-1 km
2
in size, separated by zones
of sparse to absent schlieren-bound structures. Both zones are characterized by magmatic fabrics and K-
feldspar megacryst clusters. Structure-rich and structure-poor domains form highly irregular shapes, with
no apparent preferred orientation in map view. The clusters show no clear spatial relationship to nearby
contacts or magmatic fabric orientations. In a subset of domains (Glen Aulin and Tenaya Peak) plume and
147
tube structures are weakly grouped by type. In contrast, planar schlieren and troughs appear in all clusters
of the four domains. Planar schlieren and troughs are weakly aligned with contact-parallel magmatic fabrics
at Glen Aulin, and younging or migration directions of all schlieren-bound structure types in the CP unit is,
on average, outward younging and perpendicular to contact orientations (with overall a wide range in
orientations). There is additional complexity to resolve in the pHD unit.
4.3 Outcrop-scale patterns
4.3.1 Magmatic structure patterns
Grid mapping of several outcrops illustrates several features of schlieren-bound structures: 1) Schlieren-
bound structures include a component of schlieren, but also include components that are more felsic, or
slightly different composition, than the surrounding host (e.g., layers between schlieren bases of troughs,
tube centers, Figs. 5.3b, 5.4a; herein called ‘felsic components’); 2) Schlieren-bound structures may be
deformed, evidenced by magmatic faulting and folding (Figs. 5.3a, 5.3d, 5.4d, 5.9a), and structures broken
and re-intruded by the surrounding host pluton (Figs. 5.4b, 5.4c, 5.4d, 5.9a, b, c, see also Hodge et al.,
2012); 3) While structures are typically classified as tubes, troughs, and plumes (or others, not part of this
study), outcrops less commonly show structures that represent transitions, or combinations, of structures
(Fig. 5.9b); 4) K-feldspar megacrysts, where observed in schlieren-bound structures, are often strongly
aligned with local schlieren-parallel foliation or sorted by size (Figs. 5.9a-c). This is not common in the
surrounding host, where K-feldspar megacrysts are clustered or weakly aligned in a magmatic fabric (e.g.,
Paterson, 2009).
Local schlieren-parallel foliations (Type 1) in schlieren-bound structures are discordant to magmatic
foliations in the surrounding host (Type 2-5). The schlieren-parallel foliation in the mafic (schlieren)
component of structures is rarely overprinted by other magmatic foliations. However, minerals in the felsic
components of schlieren-bound structures are rotated to Type 2, 3, 4, or 5 orientations. In addition, the
presence of schlieren-bound structures at an outcrop does not modify the orientation of the magmatic
foliations in the host (Figs. 5.9a, 5.9b).
4.3.2 Effects of magmatic deformation and rotation
Field evidence described above in each of the mapped domains indicates that some structures were
magmatically deformed (i.e., translated, rotated, and distorted), and no longer record their original
orientations. This deformation can also produce multiple younging directions within a single outcrop (Fig.
5.9a). Even in outcrops where structures appear undeformed, there may be multiple schlieren-bound
148
149
Figure 5.9 (previous page): Outcrop-scale grid maps showing complex schlieren-bound structures. (A) Pothole
Dome, Glen Aulin area (NW quadrant). Schlieren are sharply truncated, or re-intruded, by Cathedral Peak magma,
and are locally folded. The Cathedral Peak unit contains K-feldspar megacrysts from ~4-6 cm in size at Pothole
Dome. At this outcrop, K-feldspar megacrysts are densely clustered (stippled pattern) and are sorted by grain size
into funnel- or channel-shaped structures bounded by schlieren. K-feldspars are included in schlieren, and where
observed, they are aligned with schlieren-parallel fabric defined by hornblende and biotite. This is an example where
there are multiple younging directions preserved at one outcrop. The regional fabric is locally distorted into-
schlieren-parallel orientations, although a clear regional E-W orientation and contact-parallel NE-SW orientation is
still dominant. (B) Young Lakes area (NE quadrant). The northern section of the map shows a package of cm wide
schlieren troughs, with a consistent younging direction to the NNE. K-feldspar megacrysts (3-8 cm) are abundant
within schileren layers and align in a schileren-parallel orientation. Within some layers (yellow) the schlieren are
generally more megacryst-rich. The trough package is truncated by an elliptical-shaped tube structure which is
stationary (youngest rings are in the center of the tube). K-feldspar megacrysts in the tube are aligned parallel to
schlieren. K-feldspar defined fabric within the CP is weakly defined, except in and around schlieren-bound
magmatic structures. (C) Young Lakes area (NE quadrant). Folded schlieren troughs younging to the east. Schlieren
include K-feldspar megacrysts that are variably aligned- most well defined foliation in thin elongate layers, and
weak foliation in irregularly shaped schlieren.
150
structures with varying orientations and several, opposing younging indicators (Figs. 5.4b, 5.9b). Thus, the
effects of rotation are not easily quantified and generate scatter in the structural data.
However, even with this scatter, the observation that multiple types of structures show similar outward-
younging trends at regional and domain-scales indicates that the most prominent younging direction, or
orientation, measured in the outcrop is meaningful. There is no systematic rotation of the structures in the
dataset in space (e.g., changing strikes and dips towards contact zones). In addition, schlieren foliations
(type 1) are discordant to, and thus not rotated by, strain producing the host magmatic fabric (types 2-5),
suggesting they are in a primary orientation (see above). The possibility of rotated structures underscores
the importance of using large datasets at the domain- or regional-scales when interpreting patterns in
orientation or younging.
4.3.3 Schlieren petrography
Schlieren in the TIC contain between 2-8 times more hornblende, biotite, magnetite, apatite, titanite and
zircon than the surrounding plutonic material (e.g., Wahrhaftig 1979; Reid et al., 1993; Table 5.1). The
relative modal proportions of these minerals in the schlieren structures varies with the mineralogy of the
host unit. Within the TIC, the ratio of biotite to hornblende increases towards the interior of the complex
(e.g., Bateman and Chappell, 1979; Memeti et al., 2014). This trend is mirrored in the schlieren found in
different TIC units (Table 5.1). The proportion of quartz to feldspar also increases from the KC to the CP
in both schlieren and host samples.
The abundance of magnetite, titanite and apatite increases from the outer KC unit to the interior CP unit
in both host and schlieren samples, yet the relative proportions of these minerals are highly unit-dependent
(Table 5.1). For example, both the Kuna Crest host and schlieren are generally titanite-poor, relative to
younger units. In contrast, the Cathedral Peak unit is richer in titanite (~1%; Table 5.1), which is also
pronounced in the schlieren (~7%; Table 5.1). Magnetite forms crystal clusters in all schlieren samples,
which also includes titanite clusters in the CP schlieren. Zircon is sometimes observed in thin section but
is generally too small to measure quantitatively. Previous workers have also identified allanite in select
samples (e.g., Reid et al., 1993; Paterson, 2009). These changes towards the TIC interior coincide with an
increase in the grain size of the accessory minerals, and in all studied samples, the grain size of accessory
minerals is, on average, larger in the schlieren than in the host.
In thin section, schlieren-bound magmatic structures record the same type of microstructures found in
the surrounding host and are dominantly magmatic in origin. For example, interstitial quartz is pinned by
phenocrystic feldspar, biotite, and hornblende. Crystal clots of feldspar, magnetite, and titanite are common
151
in all studied examples. Crystal-plastic deformation is limited to undulose extinction in quartz, and rare
cases of deformation twinning in feldspar.
4.3.4 Whole-rock geochemistry
4.3.4.1 Major and trace elements
TIC schlieren have a wide range in SiO 2 contents (40-70 wt.%; Fig. 5.10a-d), encompassing much of
the variation in SiO 2 found in the main TIC units (48-79 wt.% SiO 2), except at the high-silica end >70 wt.%
SiO 2. As described in earlier studies (Reid et al., 1993; Solgadi and Sawyer, 2008; Paterson et al., 2008;
Žák et al., 2009), for many elements TIC schlieren samples plot at a high angle to the main TIC trend in
SiO 2-element space. Schlieren are enriched in TiO 2, MnO, and P 2O 5, relative to the host TIC for a given
SiO 2 contents and have a variable FeO t contents at SiO 2 < 62 wt.% (Fig. 5.10b). Schlieren are depleted in
Al 2O 3 (Fig. 5.10a) and CaO relative to the surrounding host, and Na 2O ranges from depleted to overlapping
with the host at higher SiO 2 content. K 2O is enriched in schlieren relative to the host at low SiO 2 and then
compositionally overlaps with the host at SiO 2 >60 wt.%. Felsic components of schlieren-bound structures
compositionally overlap with the surrounding TIC rocks. SiO 2 contents of the felsic component are between
57-78 wt.%, falling at the felsic end of schlieren compositions and in some cases matching compositions
of TIC leucogranite samples (Fig. 10). The intersection of linear trends formed by the schlieren samples
and felsic or host samples is consistently between 65-70 wt.% SiO 2.
Some large-ion-lithophile-elements such as Rb, Cs, and Eu are enriched in the schlieren relative to the
host for a given SiO 2 content, while Sr (Fig. 5.10c) and Ba are depleted (Table 5.2). Most high-field-
strength-elements, such as Y, Nb, and Zr (Fig. 5.10d) are enriched relative to the host (Table 5.2). Pb and
Sc contents in both schlieren and host are approximately equal (Table 5.2), forming one trend line.
4.3.4.2 Rare Earth Elements (REE)
The main TIC units have classic arc-type REE slopes, with La N/Yb N values between 2-44. The REE
patterns vary between units, with the CP samples showing the most fractionated patterns. TIC schlieren
generally have REE patterns parallel to the host unit they are situated in, with the exception of schlieren
samples in the Kuna Crest granodiorite (Fig. 5.11a). A key feature that distinguishes schlieren from the
main TIC units is that the schlieren have REE abundances up to 10 times greater than the surrounding host.
Felsic components also have REE patterns parallel to those of the host but display lower REE contents (as
much as 10 times lower than the host) (Fig. 5.11a-c). A felsic component sample from the Half Dome units
shows a slight positive Eu anomaly, which is not found in the sample of the adjacent host (Fig. 5.11b).
152
Figure 5.10: Major and trace element data from schlieren, felsic components, and host TIC samples. Best fit trend
lines of schlieren compositions (trends 1 and 3) are at high angle to the main TIC trend, and the trend of felsic
components (trend 2). Samples are colored based on the sample type; TIC host, schlieren, felsic components,
mixtures of schlieren and felsic material (from Solgadi and Sawyer (2008). (A) Al 2O 3 v. SiO 2 (B) FeO t v. SiO 2. (C)
Sr v. SiO 2. (D) Zr v. SiO 2. Data sources: Bateman and Chappell (1979); Peck and Van Kooten (1983); Bateman et
al., (1984); Kistler et al., (1986); Loetterle (2004); Burgess and Miller, (2008); Economos et al., (2008); Gray et al.,
(2008); Paterson et al. (2008); Solgadi and Sawyer (2008); Memeti, (2009); Žák et al., (2009); Coleman et al.
(2012), and new data (Table 5.2).
153
Figure 5.11: (A-C): Rare Earth Element patterns in TIC host, schlieren and felsic components. Plots are divided up
by major unit. (A) Kuna Crest granodiorite (sample data from Memeti (2009)). (B) Half Dome granodiorite units.
eHD and pHD have been combined to maximize data coverage. (C) Cathedral Peak granodiorite. Matching colors
indicate that the samples were collected from the same outcrop, at adjacent layers. Normalized after Sun and
McDonough (1989). (D) Whole rock isotopes
87
Sr/
86
Sr i v.ƐNd i. TIC samples are color-coded by unit. They define a
broad isotopic range (Memeti et al. 2014). Schlieren-host pairs collected at the same outcrop are circled. Fields
defined by host rock samples as shown in Memeti et al. (2014). Standard errors are smaller than the size of the
symbols and provided in Table 5.2. Data sources: Bateman and Chappell (1979); Peck and Van Kooten (1983);
Bateman et al., (1984); Kistler et al., (1986); Loetterle (2004); Burgess and Miller, (2008); Economos et al., (2008);
Gray et al., (2008); Paterson et al. (2008); Solgadi and Sawyer (2008); Memeti, (2009); Žák et al., (2009); Coleman
et al. (2012), and new data (Table 5.2).
154
4.3.4.3 Sr and Nd isotopes
The TIC has a large range initial
143
Nd/
144
Nd i (ƐNd i notation used) between -2.0- and -9.5 and initial
87
Sr/
86
Sr values (Sr i notation used), between 0.7055 and 0.7077, with the CP unit and late leucogranite
samples producing the most evolved compositions (Kistler et al., 1986; Memeti et al., 2014; Fig. 5.11d).
Within one unit, the isotopes vary up to 5 epsilon units (ƐNd i), with smaller variations in Sr i. The KC is the
most isotopically heterogeneous unit, and the pHD is the most homogeneous unit (Fig. 5.11d; Memeti et
al., 2014). Schlieren and host pairs are isotopically similar, corresponding well with the isotopic
composition of the unit they are found in. However, the schlieren and host pairs are not identical.
Differences in ƐNd i between paired schlieren and host samples are <1 epsilon unit, and in pHD and CP
samples this difference is minimized to within 0.2 epsilon units. The schlieren samples in all but one
example have higher Sr i ratios than their adjacent host (Fig. 5.11d). These differences are greater than the
2σ analytical uncertainty (0.0006-0.0025%, smaller than the size of the symbols in Fig. 5.11d).
4.4 Relative timing indicators
The zircon isochron map of Memeti et al. (2014) provides the only absolute timing constraint on
schlieren-bound structure formation; notably, schlieren-bound structures span the entire 10 m.y. history of
the complex. At the scale of a single schlieren trough, Memeti et al. (2014) also found identical ages
between zircons contained in the schlieren and zircons in the surrounding host.
Several examples have been found where intrusive contacts sharply truncate schlieren-bound structures
(e.g., Fig. 5.4a) indicating that these schlieren-bound structures formed prior to arrival of a new magma
pulse. Examples have also been found where materials of one unit slumped off a contact into the adjacent
unit to form a compositionally-defined magmatic structure, requiring that they formed immediately after
juxtaposition of two magma pulses. Elsewhere, schlieren-bound structures seem to cluster along internal
contacts indicating the contact already existed (e.g., Žák and Paterson, 2005). Together, particularly
combined with the temporal history established from zircon ages, these observations indicate that
compositionally defined magmatic structures formed from many repeated events both prior to, during, and
after juxtaposition of magma pulses, over the entire 10 m.y. history of the TIC.
The relative timing between different types of schlieren-bound structures (i.e., planar schlieren, troughs,
tubes and plumes) is not clear. This is because different schlieren-bound structures are generally spaced ~1-
10 m apart and do not cross-cut each other (see also Hodge et al., 2012), except in rare cases (Fig. 5.9b).
Often, transitional-type structures are observed, for example, trough structures that grade into tubes, or
bifurcating tubes (Fig. 5.4b), where a local relative timing relationship can be determined (e.g., Paterson,
155
2009). Overprinting relationships between type 1 magmatic fabrics in schlieren-bound structures and fabric
types 2-5, described above, suggest that schlieren-bound structures and magmatic fabric types 2-5 formed
synchronously.
In addition, schlieren major and trace element compositions consistently intersect TIC host compositions
between 65-70 wt.% SiO 2. We used these intersections to estimate the bulk composition of a TIC magma
that we could model under fixed pressure and varying temperature conditions. Werts et al. (2020) and
Barnes et al. (2020), using the thermometer of Putirka, (2016a), calculated hornblende crystallization
temperatures ranging from >800°C to ~650°C across all TIC units. As hornblende is abundant in all
schlieren studied, the temperature range 800-700°C was considered relevant to schlieren formation to a first
order. Further, a study of apatite in TIC rocks (also abundant in schlieren) suggested that at apatite saturation
(~900°C) magma crystallinity was 9% (Piccoli and Candela, 1994), which we use as a minimum
crystallinity here. Using the estimated bulk composition from schlieren-host TIC intersections (Table 5.3),
with 4 wt.% H 2O (due to the presence of hornblende) and 65 wt.% SiO 2, we calculated effective viscosities
of the bulk composition containing between 10% to 90% crystals from 800-700°C using the method of
Murase et al. (1985)(equation 3). At this bulk composition at 800°C, effective viscosity increases from 10
6.9
Pa s to 10
9.3
Pa s with increasing crystal contents and reaches a maximum of 10
10.3
at 700°C (Table 5.3).
Werts et al (2020) demonstrated that the bulk composition of many TIC samples is generally more mafic
than the composition of melts in equilibrium with hornblende due to crystal accumulation. We compared
our bulk-rock results to a model using melt compositions in equilibrium with CP hornblende, containing 4
wt.% H 2O and 76 wt.% SiO 2 (equations from Zhang et al., 2017). This model resulted in calculated effective
viscosities ranging between 10
8.2
Pa s to 10
10.5
Pa s at 800°C, with a maximum of 10
11.8
Pa s at 700°C.
5. Discussion
Our field and geochemical results allow us to analyze schlieren-bound magmatic structure formation
across several spatial scales. First, outcrop-scale observations and sample compositions are discussed,
forming the basis for physical and chemical models of schlieren formation. Kilometer-scale domainal
clustering and 10’s kilometer scale regional patterns are then discussed to characterize the scale and
mobility of the magma mush system(s).
156
Table 5.3: Magma viscosity summary for bulk composition and calculated melt composition, using equation (3) of
Murase et al. (1985).
157
5.1 Outcrop scale
5.1.1 Evaluating models to form schlieren
5.1.1.1 New (mafic) magma injection
Models concluding that schlieren form from distinct, mafic magma batches (e.g., from a similar source
as mafic magmatic enclaves) cannot describe the origin of most TIC schlieren (Fig. 5.1a). The varied
geometries and field patterns of schlieren from outcrop to regional scales are difficult to explain by injection
of mafic magma batches alone. In addition, mineralogy and whole-rock geochemistry suggests that
schlieren are sourced from the surrounding host magma, rather than an ‘exotic’ mafic magma. Schlieren
mineral assemblages reflect, to a first order, the assemblage of the surrounding host, with different modal
abundances (Table 5.1, see also Reid et al., 1993; Alasino et al., 2019). Parallel REE patterns further suggest
that the mixture of REE-bearing minerals in schlieren is matched in the host and indicates a common magma
source, distinct from enclaves (Reid et al., 1993), which is supported by the similarity in host and schlieren
isotope compositions (Fig. 5.11d). The unique bulk-rock compositions of schlieren, together with modal
assemblages, indicate that they are cumulates (Fig. 5.10, Table 5.1, see below). Thus, their present bulk
compositions are not representative of the more evolved parental magma from which they were derived.
One interpretation of this parental magma composition is that it falls between 65-70 wt.% SiO
2, where the
best fit line of schlieren compositions intersects the main TIC trend. Another, determined from calculated
melts in equilibrium with TIC hornblende compositions, suggests a range from 67-79 wt.% SiO 2 (Werts et
al., 2020).
5.1.1.2 Diffusion and re-crystallization
Thermochemical diffusion and recrystallization end-member models are also not supported by our
combined structural and geochemical datasets (Fig. 5.1a). Field patterns of schlieren, including cross-
cutting relationships, are not explained by this model. Solgadi and Sawyer (2008) demonstrated, using
hornblende Mg-number, that schlieren lacked any compositional gradients from base to top, which either a
thermal or chemical gradient should impose. In-situ mineral analyses of samples across all TIC units reveal
complex zoning patterns, which suggests thermochemical diffusion did not play a significant role in their
history (Memeti et al. 2014; Barnes et al., 2016c; Werts et al., 2020). In a liquid immiscibility model, a
form of chemical diffusion (e.g., Glazner et al., 2012; Bartley et al., 2019), isotope compositions should be
identical between schlieren and host, and the felsic component to the mafic schlieren should be enriched in
alkalis, Rb, Cs, Sr and Ba (Ryerson and Hess, 1978; Hurai et al., 1998), neither of which are observed in
TIC samples. Instead, the felsic component has a similar composition to the host. Mixing schlieren and
felsic components together as demonstrated by analyses of Solgadi and Sawyer (2008) and shown on Figure
158
5.10 produces a compositional array parallel to schlieren trends, distinct from the host. Furthermore, if
mafic schlieren represented the Fe-rich liquid of an immiscible pair, the density of the schlieren should be
considerably greater than the host and would be expected to separate from felsic components and the host,
rather than be interlayered with them.
5.1.1.3 Magmatic flow
Our proposed model to explain field and geochemical datasets involves hydrodynamic flow-sorting of
crystals in a mush to produce schlieren layering (Fig. 5.12c). TIC schlieren contain dense minerals and
reduced abundance of low-density feldspar. Together with evidence of common mode and size grading of
schlieren layers, these observations are consistent with mineral accumulation that depends on the density,
size, and shape of individual crystals (e.g., Bhattacharji and Smith, 1964; Kawabata et al., 2013). Further,
outcrop scale cross-cutting relationships, truncation of structures by the host magma, and consistent patterns
of outward-younging orientations within and across domains that span a range of crystallization ages
requires physical, magmatic erosion and flow processes to operate synchronously and repeatedly. Schlieren
form an array of major and trace element compositions at high-angle to the main TIC trend, which are
interpreted here to represent the varying degrees of crystal accumulation of hornblende, biotite, apatite,
titanite and zircon that dominate schlieren compositions. As a suitable complementary felsic fractionate is
not observed (see above), we interpret that any melts resulting from schlieren accumulation were mixed
back in with surrounding host magma, to form the felsic components of schlieren-bound structures (see
also Alasino et al., 2019).
Small shifts towards radiogenic Sr
i in schlieren, and variable differences in εNd i between schlieren and
host are also compatible with a flow-sorting crystal accumulation model. Measurable isotopic shifts at the
scale of the schlieren and adjacent host could be caused by: 1) crystal-melt fractionation from an isotopically
hybrid magma (e.g., Beard, 2008); 2) selective crystal accumulation of Rb-bearing minerals, or Sm-bearing
minerals; or 3) alteration of minerals by late magmatic fluids (e.g., Alasino et al., 2019; Li et al., 2019).
Diffusion, or metasomatism by late magmatic fluids, should leach mobile elements, such as Rb, from
samples, leading to low Rb/Sr. However, the relationship between pairs of related samples is not the same
for all samples studied or compiled; in most cases (5 samples out of 6), Sr i is more radiogenic in schlieren,
but εNd i is variable. In addition, the preservation of textural and compositional zoning in Rb-bearing phases
such as K-feldspar, and Sm and Nd-bearing phases such as hornblende (and apatite) demonstrates that
diffusion of fast- and slow-diffusing elements was limited in the TIC (Memeti et al., 2014; Barnes et al.
2016c; Werts et al., 2020). Thus, isotopic differences between schlieren and host are interpreted as primary
magmatic features of the samples resulting from processes 1) and 2), rather than secondary features (3).
159
Figure 5.12: Summary of multi-scale dynamic processes operating in mobile magma mushes. (A) Mixing bowl
scale. Return flow of the mush is triggered by intruding pulses in the center of the complex. Thermal/mechanical
gradients aid convection. (B) Domain scale activity summarized as a 3D diagram. Structures form in fluidized zones
that are more melt-rich due to local changes in pore-pressure, viscosity, and stress, while crystal rich zones represent
zones that are absent of schlieren-bound structures. They still contain melt as they contain magmatic fabrics, dikes
and other structure. (C) Structure-scale processes demonstrated for a magmatic trough. Flow sorting, filter pressing,
crystal accumulation and melt escape are all important processes at the mm-cm scale. Flow along the trough channel
produces a lineation in hornblende.
160
As the TIC is isotopically heterogeneous (Kistler et al., 1986; Gray et al., 2008; Memeti et al., 2014),
models that fractionate crystals from a hybrid magma and create further heterogeneity, as illustrated by
Beard (2008) are considered likely to explain schlieren-host isotopes in combination with field data and
mineral assemblages. Furthermore, selective fractionation and accumulation of minerals could amplify the
difference in isotope composition between schlieren and the surrounding host (e.g., Kendall et al., 1995;
Blum and Erel, 1997). Rubidium-bearing minerals such as biotite, abundant in the CP schlieren, could result
in high Sr i ratios, while accumulation of Sm-bearing minerals such as apatite and hornblende could result
in high
143
Nd/
144
Nd i and thus more primitive εNd i values. An example of the latter is the KC schlieren
sample, which contains ~40% hornblende.
5.1.2 The significance of forming schlieren between 65-79 wt.% SiO 2
TIC host magmas were already mixed before schlieren formation based on whole-rock isotope trends
(Memeti et al., 2014). The flow-sorting model requires that most, if not all, of the highly concentrated
minerals (biotite, hornblende, magnetite, titanite, apatite, and zircon) were saturated in the host magma
before schlieren minerals were separated and requires sharp rheological boundaries to exist in the mush.
Melt presence is required to deform schlieren-bound structures by magmatic faulting and folding, for the
felsic component of schlieren accumulation to mix back in with the host magma, and also to form magmatic
fabrics (discussed below). The structure must have formed in a sufficiently strong magma to erode and
sometimes break structures without losing definition in individual layers and also to form sub-vertically
dipping schlieren (e.g., Bergantz, 2000; Weinberg et al., 2001; Paterson, 2009; Alasino et al., 2019). Each
of these observations suggest that the magma was chemically evolved and crystal-rich during schlieren
formation.
The intersection of major and trace element best-fit trends of schlieren with TIC host and felsic
component samples lies consistently between 65-70 wt.% SiO
2. One possibility is that this represents the
bulk composition of the parental magma when the schlieren and felsic components were formed. However
calculated melt compositions from hornblende, using the Zhang et al. (2017) chemometer are between 67-
79 wt.% SiO 2, highlighting the effects of crystal accumulation (Werts et al. 2020; Barnes et al., 2020). In
both cases we can place initial constraints on magmatic conditions during schlieren formation, specifically
the magma viscosity. During the temperature interval of 800-700°C, when schlieren likely formed,
calculated effective viscosities increased from ~10
7
to 10
11
Pa s in both starting compositions, with
considerable overlap between models, and the calculated melt compositions providing the highest values.
Although these values appear on the high end to facilitate magmatic flow, previous studies have
demonstrated that small variations in pore pressure or strain rate may manifest as large variations in magma
161
viscosity, without substantially reducing temperature or crystallinity of the magma to generate local
rheologic boundaries (e.g., Weinberg et al., 2001; Sparks et al., 2019, Bergantz et al., 2017). Numerical
modeling of hydrogranular crystal networks suggest that, at the crystal scale, particle-particle-melt
interactions and forces such as melt lubrication, crystal clustering, jamming, and the migration of particle
force chains all can aid or inhibit movement of crystals and fluids under magmatic conditions and thus
instantaneously and dramatically modify mush rheology to promote or inhibit magmatic flow (Bergantz et
al., 2017; Carrara et al., 2019; McIntire et al., 2019). Thus, bulk or effective viscosity does not fully describe
the mobility or the dynamics of the system. These processes that locally and transiently reduce viscosity
may also enhance melt migration out of the schlieren and into the surrounding host mush; potential
mechanisms include filter-pressing or porous flow (e.g., Weinberg et al., 2001). Similar magmatic flow
models were previously proposed in granitoid systems by Weinberg et al. (2001), Paterson (2009), and
Alasino et al. (2019), who invoked deformation-assisted mineral-melt separation to form schlieren
structures in the Tavares pluton, the TIC, and the Sonora Pass Intrusive Suite, respectively.
5.1.3 Evaluating models to form different schlieren-bound structures
The results of this study supports a range of physical processes to form different schlieren-bound
structures, as described by Weinberg et al. (2001) and Paterson (2009). These include boundary flow, filter
pressing, and Rayleigh-Taylor instabilities. Combinations of these processes might operate to create
transitional-type structures described in section 4.4 (e.g., troughs grading into tubes). Instabilities triggering
fluidization of the mush and structure formation may be driven by sinking of dense objects or magma
mushes (e.g., Weibe et al., 2007), or the rise of buoyant objects (e.g., exsolved volatiles, extracted melts:
Clarke 2003; Dietl et al., 2010), both of which have been proposed to form individual tube structures. In
addition, the intrusion of new magma pulses into the chamber may promote thermal, chemical and
mechanical instability (e.g., Pons et al., 2006). This behavior has been suggested to drive crystal avalanches
(forming troughs; Solgadi and Sawyer, 2008; Alasino et al., 2019), as well as local rising of magmas
(forming tubes or plumes; Paterson, 2009).
The diversity in the types of observed magmatic structures within most map domains (GA, TP, LC)
indicates that at a single outcrop (~10-20 m scale), local flow fields were complex enough to form
geometrically distinct troughs, tubes and plumes. For example, at several localities at Glen Aulin there may
be both vertical magma flow through tube structures, and moderately down-dipping flow in troughs. This
raises the question if different structures may require or favor different magmatic conditions, or driving
mechanisms, to form (e.g., Fig. 5.1b). In contrast, within the Young Lakes area, tubes and plumes are sparse
162
and planar schlieren and troughs are dominant, which could suggest the magmatic flow field was distinct,
and perhaps more uniform.
There is a weak spatial control on the types of structures that form. At Glen Aulin, plumes are generally
concentrated in the east, and tubes are concentrated in the west. The eastern edge of the domain is the
youngest edge of the domain, formed across a steep age gradient (Memeti et al., 2014 isochron plot), which
suggests that this could result from a temporal shift in physical flow mechanisms, in addition to a spatial
shift.
5.2 Domain scale
5.2.1 Magma chamber domains are spatially heterogeneous as a result of flow instabilities
TIC schlieren-bound structures are spatially clustered between large (~0.1-0.5 km
2
) zones where
structures are sparse to absent. The difference between the structure-rich and structure-poor zones could
reflect: 1) zones which were ‘fluidized’, more melt-rich and therefore more mobile mush, producing
abundant structures, separated by mushy zones that were unable to flow as dynamically and therefore lack
schlieren-bound structures (a form of partitioning); 2) structures were uniformly present throughout the
chamber and were subsequently locally erased by later mixing and homogenization.; 3) structure rich
domains represent zones of sharp flow velocity gradients, that could have increased mobility due to positive
feedbacks as structures started forming; or 4) structure rich zones represent channels of porous flow. All
scenarios imply that magma dynamics was spatially heterogeneous at the 1-10 km scale. It appears more
difficult to completely erase the evidence of schlieren-bound structures, because that would require
thorough mixing at high crystal contents. Field observations show that schlieren-bound structures can
deform semi-rigidly (by faulting and folding and re-intrusion) or mingle with the host magma, suggesting
it may be mechanically challenging to fully erase the schlieren once they form. This is supported by the
weak to absent overprinting of local type 1 fabrics by regional type 2-5 fabrics in schlieren, but more
pervasive/intense overprinting in the felsic parts of the structures.
Spatial clustering of schlieren-bound structures is documented in other plutons. In the Vinalhaven
intrusive complex, schlieren-bound structures (mostly tubes) “occur widely but sparsely”, wherever sinking
enclaves could sufficiently shear the surrounding crystal mush (Wiebe et al., 2007). The pluton was strongly
vertically stratified, controlling the path of the enclave (Wiebe et al., 2007). Clarke et al. (2013) suggested
that tube structures were preferentially found near the pluton margins in the Halifax pluton, Nova Scotia, a
result also found in K-feldspar pipes associated with schlieren by Rocher et al. (2018) in the San Blas
intrusive complex, Argentina. Clarke et al. (2013) further posited that the cause of structures clustering
163
within the contact zone could be attributed to “restricted zones of permeability”, which could represent the
‘fluidized zones’ described here.
5.2.1.1 Processes triggering domain-scale flow instabilities
The occurrence of domain-scale magmatic structure patterns suggests that broader-scale dynamic
instabilities were present in TIC mushes, and not restricted to the outcrop-scale where individual structures
formed. Several processes may have operated in the TIC. The intrusion of new pulses, or upwards migration
of volatiles, could lead to mobilization and convection of the mush (e.g., Bachmann and Bergantz, 2006;
Burgisser and Bergantz, 2011; Pinotti et al., 2016; Fig. 5.12a). In computational models, the heterogeneous
distribution of particle force chains maps transient high and low stress zones in the mush, that has
implications for pore-pressure variability, and overall mush porosity (e.g., Estep and Dufek, 2012; Bergantz
et al., 2017). Focused porous flow could also create zones of increased permeability that promote flow
instability (Keleman et al., 1995; Paterson et al., 2012). In other cases (e.g., Sawmill Canyon) structure-
rich domains may be spatially restricted to magma sheets or cracks (e.g., Clarke and Clarke 1998; Paterson
et al., 2008). Pluton roofs and internal contacts may generate vertical domain-scale instabilities. Roof
instability can lead to sinking of crystalline mush, which could then be further segregated by density (e.g.,
‘crystal-rich drips’; Bergantz and Ni, 1999; Rocher et al., 2018; Carrara et al., 2019). We infer that the
present-day exposed levels of the TIC are representative of the upper levels of the magma body, close to
the roof. Evidence includes local exposures of roof flaps and an approximately uniform elevation plane of
highest (plutonic) peaks across the TIC (e.g., Bateman, 1992; Cruden et al., 2017). Sub-vertical, internal
contacts are ideal locations for solidification fronts, leading to mush collapse and avalanching along steeply
dipping walls (e.g., Marsh 1996; Bergantz, 2000; Rocher et al., 2018; Alasino et al., 2019). Domain-scale
solidification fronts may play a larger role in the Glen Aulin area, relative to the Young Lakes or Lyell
Canyon areas, as major unit contacts are closely spaced, with potentially greater thermal, chemical, and
rheologic gradients across them. While all of these processes likely were active during the evolution of the
TIC, the lack of similarity in domain size, shape and orientation, together with our outcrop- and regional-
scale structural patterns favor forming domains by spatial heterogeneity in particle-particle interactions,
driven by new pulses.
5.2.2 Planar schlieren and troughs weakly align with nearby contacts due to boundary conditions
In the Glen Aulin domain, there is a weak preference for planar schlieren and trough orientations in a
NE-SW direction that aligns with the well-defined Type 2 foliation, and with the adjacent contacts of the
Cathedral Peak unit with the older Half Dome units. This is also qualitatively suggested in N-S striking
contact-parallel schlieren at Tenaya Peak, although this interpretation is complicated by a similarly-oriented
164
Type 3 foliation. In contrast, this relationship is not apparent in the Young Lakes domain, where the nearest
contact (2-3 km away) is NW-SE striking, yet planar schlieren and troughs are oriented at high angle to the
NW-SE foliation, and instead show a preference for NE-SW striking orientations.
The relationship between schlieren orientations and the contact is ambiguous at Lyell Canyon,
potentially because the contact is a wide, gradational zone between units that hybridized (Žák and Paterson,
2005; Paterson et al., 2011, 2016; Oppenheim et al., in review) At Glen Aulin, the contacts between eHD
(probably near-solidus) and the pHD and CP (active magma mushes of varying crystallinity, Fig. 5.12b)
are proximal across a 100-200 m wide zone. Thermal and chemical gradients could be significant in driving
flow gradients, and perhaps play a larger role in defining the orientations of the structures at Glen Aulin
than in the other domains (Fig. 5.12b). As domains devoid of structures are interspersed with domains
concentrated in structures across zones both near and far from contacts, we suggest that nearby contacts are
not controlling the formation of schlieren, but potentially play an indirect role in their orientation at the
schlieren-bound structure scale.
5.2.3 Crystal clustering is widespread and independent of domains
The observation that K-feldspar megacryst clusters are found widely in the TIC, even in areas of sparse
to absent schlieren complicates the story of heterogeneous, isolated, schlieren-rich domains (e.g., Paterson
et al., 2005). In some cases they are associated with schlieren, but they also represent a similarly local
(meter-scale), selective mineral sorting and accumulation process that does not seem to be as spatially
restricted. The K-feldspar megacryst clusters could be mechanically sorted (Paterson et al., 2005; Vernon
and Paterson, 2006; Paterson, 2009), possibly representing ‘logjams’ in a less-dynamic flow field, brought
together in a cluster due to their large size (Weinberg et al., 2001).
5.3 Regional scale
5.3.1 Outward-younging patterns reflect re-mobilization of the mush by new pulses
At the regional scale, the TIC is defined by four compositionally distinct units, which are bounded by a
wide variety of external and internal contacts. Multiple magmatic fabrics are found in each of the units, and
cross several of these internal contacts. Compositionally defined magmatic structures occur across all TIC
units. Schlieren-bound structures are clustered, with a dominant pattern of structure migration or younging
outwards, towards internal contacts with older units. The regional patterns listed above may provide insight
into viable emplacement models for the TIC, including dike emplacement, sill/laccolith emplacement, or a
combination of sheet and diapir emplacement. Emplacement models are described below, specifically
focusing on what these models predict for schlieren-bound structure patterns at regional scales and how our
165
regional-scale results (contacts, fabrics, schlieren-bound structures) compare. A broader overview of TIC
emplacement is provided in section 5.5.
5.3.1.1 Incremental growth models for the TIC
Some studies have proposed that the TIC was emplaced by a series of stacked sills, or laccoliths that
were subsequently folded and eroded to produce the nested map pattern of major units (Bartley et al., 2006;
Coleman et al., 2012). Individual sills freeze before younger sheets or sills are intruded, and so deeper
source regions are responsible for observed compositional/textural heterogeneity in a closed system. In this
model, outward-younging patterns of schlieren-bound magmatic structures represent an artifact of the
folding of individual sills, assuming that the structures were initially horizontal (layer-parallel) and upward
younging, as suggested by Bartley et al. (2018). The model also predicts that schlieren-bound structure dips
should systematically shift from the center of the TIC to the margin, with initially vertical and upward-
younging orientations in the center, towards rotated, horizontal, and outward-younging orientations at the
margins. This model cannot explain why schlieren-bound structures are clustered. We have found no
evidence for systematically changing structural dip across transects of the TIC, in any of the mapped
structures. Further, magmatic fabrics show no evidence for being axial planar to a TIC-scale fold, nor is
there a systematic change in dip, as magmatic lineations are largely sub-vertical throughout the TIC. The
observation that fabrics can overprint unit contacts (e.g., Žák et al., 2007), and that crystal mixing occurred
between units (Chambers et al., 2020) precludes the idea that older units were completely solidified before
emplacement of new batches.
A similar model proposes that the pulses are vertical dikes instead of sills, which also freeze before
subsequent dikes are emplaced (Bartley et al., 2018). In this case, older dikes are displaced outwards by
younger, producing a strict temporal sequence. This model is also problematic for the TIC, as it shares
several of the same issues as the sill model above, with some additions: Firstly, sheeted contacts are
restricted to locally mapped zones (Memeti et al., 2010; Paterson et al., 2008; Žák et al., 2009). Secondly,
the schlieren-bound structures have a wide range of types, and do not show any preferential orientations
that would be consistent with flow along dike walls. The observations of magmatic deformation, erosion,
and flow of crystals with mixed mineral populations is also inconsistent with a dike model (e.g., Solgadi
and Sawyer, 2008).
A third model suggests that the TIC grew initially by sheets, which amalgamated into larger magma
bodies and facilitated the emplacement of diapirs (e.g., Memeti et al., 2010; Paterson et al., 2011). In this
model, schlieren-bound structures form from magmatic flow as a response, at the regional scale, to the
intrusion of new magma batches. Older and younger batches can interact within the magma chamber to
166
generate emplacement-level, open system magma mixing. This model explains observations of internal
contacts that vary along and across strike as regions of differential magmatic erosion and recycling, and the
regional extents of multiple magmatic fabrics as a record of regionally-extensive melt-present regions
(Paterson et al., 2016). Below we discuss our schlieren-bound structure results in the context of this model:
5.3.1.2 Return flow model
Outward younging directions in the Cathedral Peak unit are consistent with a model where new pulses,
intruded at the center of the complex, drives displacement of the mush and causes it to flow outwards and
downwards (see also Bateman, 1992; Paterson et al., 2016; Wiebe et al., 2017; Fig. 5.12a). Physical
gradients (e.g., flow velocity, viscosity, density) and thermal gradients help drive convection of the mush
(determining the relative size of rising and descending cells), as well as define the regions in which structure
formation and preservation are favorable (Fig. 5.12a). At the margins of the mush, cooler, more crystalline
portions sink, and flow downwards. The intruding hotter pulses re-heat the mush to create melt-rich regions,
as well as draw lower levels of the mush pile to higher levels of the chamber (Fig. 5.12a). This region of
mobile mush is referred to as the mixing bowl, after Bergantz et al. (2015, 2017). We propose that
movement of the mush occurred within a vertically extensive magma mush body, accommodated by
downward flow that creates space for new magma batches (Fig. 5.12a). This is supported by sub-vertical
external and internal contacts, a steep magmatic lineation, moderate to steeply-dipping schlieren and
vertical tube axes (Oliver, 1977; Paterson, 2009; Paterson et al., 2011). The return flow model implies that
towards the base of the magma body, the structures should start to young inwards, influenced by the
downward flow of the mush and the intruding pulse(s) (Fig. 5.12a). Inward younging structural patterns are
reported in other plutons in the central Sierra Nevada (e.g., Alasino et al., 2019), however this model would
optimally be tested in a tilted pluton, where multiple levels are exposed.
5.3.1.3 Alternative models
There are likely multiple causes of younging, or migration directions measured in schlieren-bound
structures. Flow in magma sheets related to cracking of the magma mush is one mechanism that could
create a broader-scale younging direction (Paterson et al. 2008; Žák et al. 2009). This process is likely to
have dominated where sheeted zones are documented.
Other scenarios that could cause structures to migrate and young outwards, such as regional extension,
a thin, laterally spreading mush, or a static, inwardly cooling mush, are not supported by our datasets. In
the former, the regional dextral-transpressive tectonic setting recorded in the surrounding host rock, as well
as rapid crustal thickening, regional inward magmatic focusing patterns, and steep magmatic lineations in
167
the TIC, all preclude any extensional emplacement model (e.g., Žák et al., 2005; Cao et al., 2015; Hartman
et al., 2018; Cao et al., 2016; Ardill et al., 2018). We note that extension could occur locally in the return
flow model. In a static, cooling mush, where cooling at the margins outpaces intrusion by new pulses, the
structures are predicted to young inwards, not outwards, due to collapse of the newly solidifying mush at
the margins.
5.3.1.4 Significance of return flow: mapping mixing bowls
As the outward younging pattern is found in each of the TIC units, this process occurred repeatedly over
the 10 m.y. lifespan of the TIC. In addition, the outward-younging structures across four mapped domains
of the CP unit are consistent with the possibility that it was as a single large mixing bowl (see below section
5.6; Fig. 5.12). In contrast, inward-younging structures in the eastern porphyritic Half Dome unit observed
in the Lyell Canyon area represent an exception to the outward-younging pattern found elsewhere in the
CP and pHD. Although the dataset is limited in size, one possibility is that the pHD mixing bowls were
smaller than the CP mixing bowl, and that the Lyell Canyon structures and the Tenaya Peak structures
record events from two different pHD mixing bowls. Although much of the pHD unit has been removed by
the CP, mapping larger areas containing pHD structures can test this hypothesis.
5.3.2 External processes are not a significant control on schlieren-bound structure orientation
Tectonism has been proposed to cause magma mush instability on the timescales of seconds to minutes
(e.g., earthquakes; Davis et al., 2007), and at the million-year time-scale can form fabrics and extract melts
(e.g, Paterson et al., 1998; Garibaldi et al., 2018; Bachmann and Huber 2019). Regional tectonism during
emplacement of the TIC was dextral transpressive, with metavolcanic and metasedimentary host-rock fabric
orientations consistent with the regional NW-SE magmatic foliation found in TIC domains (e.g., Žák et al.,
2007; Cao et al., 2015; Hartman et al., 2018). At the scale of the intrusive complex, regional E-W magmatic
foliations and vertical lineation are pervasive. Schlieren do not show a clear angular relationship to either
regional foliation orientations within any of the mapped domains. This supports the conclusion that strain
caused by local magmatic flow, forming schlieren-bound structures, dominated over regional-scale strains
that formed fabrics, and suggest this is due to higher strain rates in structures formed by magmatic flow
(Žák et al., 2007; Paterson et al., 2008). Magmatic fabrics are time-transgressive throughout the 10 Myr
lifespan of the TIC, indicating that the regional strain fields were long-lived (Žák et al., 2007). Schlieren-
bound structures, found in all TIC units, are thus also a time transgressive record of internal magmatic
processes.
168
5.4 Timing of schlieren-bound structure formation
Examination of the field and geochemical datasets indicate that schlieren-bound structure formation
occurred in a crystal-rich mush, where crystal frameworks were important to form, deform, and also
preserve schlieren-bound structures. In the broader context of the TIC, schlieren-bound structures formed
pre, syn, and post formation of major unit contacts, and after crystallization of minerals contained in
abundance in the schlieren-bound structure (e.g., zircon; Memeti et al., 2014; hornblende). Schlieren
formation was widespread, occurring repeatedly over the 10 Myr lifespan of the TIC. In this way, schlieren
are time-transgressive, reflecting a record of magmatic conditions that occurred at some point throughout
the entire TIC.
Magmatic erosion, re-deposition and deformation occurred in a magma that had a yield strength (see
also Hodge et al., 2012). Since these processes occurred in quick succession at a single outcrop, it suggests
that there were complex interactions between crystallinity and effective viscosity, particle-particle
interactions, pore-pressure, and crystal lubrication effects (Weinberg et al., 2001; Paterson, 2009; Bergantz
et al., 2017; Carrera et al., 2019). Discordance between local schlieren-defined fabric and the regional
fabrics and the overprinting relationships within felsic components indicates that the strain fields forming
regional fabrics and local schlieren fabrics were active at the same time, but at different rates. The schlieren
fabric likely formed at faster strain rates than the regional fabrics, producing discordance. When the rate of
magmatic flow and deposition slowed, the regional fabric was able to re-orient minerals in melt-rich felsic
components, producing the overprinting relationships (e.g., Paterson et al., 2008).
Mineralogically, quartz abundance varies little in abundance between schlieren and host samples, which,
together with petrographic textures, suggests that it was the main mineral to crystallize late in the magmatic
history (after schlieren formation) from trapped, interstitial melts that were evolved. The bulk rock
compositions and equilibrium melt compositions suggest that schlieren formed from an evolved magma
mush, containing 65-79 wt.% SiO
2.
5.5 Implications for emplacement models and the minimum sizes of magma chambers
The TIC is a composite magmatic complex, containing gabbro to leucogranite with granodiorite and
granite the most abundant rock types. The complex spans a 10 Myr zircon crystallization history, with four
major units bounded by sharp to gradational contacts that internally range from homogenous (at outcrop
scale) plutons to sheeted zones (Memeti et al., 2010; Žák et al., 2009). To explain these observations,
different incremental emplacement models have been proposed, and are outlined above (section 5.3).
Regionally extensive magmatic fabrics (e.g., Fig. 5.1a), recycling of older units into younger, compositional
169
evidence for widespread magma mixing and recycling in bulk rocks and in minerals, geochronologic
studies, and thermal models all suggest that TIC magma chambers (interconnected melt regions) in the
Cathedral Peak unit, and in each of the older units, were volumetrically extensive (~1000 km
3
) and long-
lived (up to 1.5 Myr) (e.g., Žák et al., 2007; Matzel et al., 2005, 2006b; Solgadi and Sawyer, 2008; Paterson
et al., 2011; Memeti et al., 2010, 2014; Barnes et al., 2016c; Paterson et al., 2016). Together with the above-
listed evidence, outcrop-scale flow features, domain-scale clustering and regional-scale outward-younging
schlieren-bound structure patterns presented here require a magma emplacement model where
incrementally emplaced new magma pulses can interact with older pulses dynamically by magmatic flow,
transferring mass and energy to the larger-scale interconnected-melt region, or magma mush. This has been
proposed to occur by the interaction of dike- or diapir- shaped pulses that amalgamate in time and space to
form long-lived, dynamic magma bodies. (e.g., Miller and Paterson, 2001; Memeti et al., 2010; Cao et al.,
2016).
The multi-scale observations of magmatic structures also allow us to place additional constraints on the
minimum sizes of likely magma chambers that existed in the Cathedral Peak unit. As the younging patterns
in each study area of the CP unit are all consistently outward (Fig. 5.12a), then the southern portion of the
CP unit, encompassing all four domains, represents the minimum size of one mixing bowl, or magma
chamber, in the return flow model described in section 5.3. This results in an area of about 150-175 km
2
,
or a volume between 750-1750 km
3
(using the vertical extents estimated by Karlstrom et al., 2017). The
zircon isochron map of Memeti et al. (2014) suggests that this mixing bowl formed across ~1.5 m.y. At
shorter timescales, the domainal fluidized zones could each represent their own active magma chambers,
capable of magma mixing, mingling and fractionation. Their sizes, between approximately 0.5-2 km
2
in
area (vertical extents unknown), suggest they could be locally important for producing structural and
compositional diversity in plutons, but are unlikely to explain the outcrop-scale to map-scale homogeneous
appearance of the well-mixed CP unit.
5.6 Implications for the behavior of magma mush systems
The above observations indicate crystal-rich mushes, even in upper crustal magma chambers, can be
highly dynamic environments. Numerical modeling of mushy plutonic systems has demonstrated the
potential for complex crystal-melt and crystal-crystal interactions that schlieren may represent just one
(visible) example of (e.g., Bergantz et al., 2017). The ‘mixing bowl’ model of Bergantz et al. (2015) and
Schleicher et al. (2016) displays a number of features that match field-based structural observations in the
TIC. These include magmatic faulting, slumping, and folding of schlieren, and erosion of earlier formed
layers. In addition, the spatial heterogeneity of the deforming mush, with locally ‘active’ and ‘static’ areas,
170
in the model is consistent with our interpretation of the spatial clustering of magmatic structures. Finally,
the outward younging patterns predicted by the upper parts of the model concur with our field dataset. The
widespread distribution of the structures across the TIC suggest that dynamic hypersolidus conditions, as
exemplified by the ‘mixing bowl’ model, existed in different units and areas of the TIC repeatedly
throughout its entire ~10 m.y. lifespan. The schlieren-bound structures represent sites of compositional
differentiation occurring at the emplacement level, and sites of local magmatic recycling (e.g., Paterson et
al., 2016).
No clear relationship yet exists between the orientation of regional magmatic fabrics (Types 3-5) and
schlieren-bound structure orientations (Type 1). This suggests that internal magmatic processes dominate
in the formation and orientation of schlieren-bound structures. Garibaldi et al., (2018) proposed that
magmatic fabrics and magmatic structures (e.g., miarolitic cavities) in the Huemul pluton, in the Andes,
record evidence for a crystal-rich mush mobilizing interstitial rhyolitic melts by tectonic filter pressing.
Other studies have also considered tectonism as a mechanism to mobilize crystal-mushes and re-orient
earlier-formed magmatic structures, including fabrics (e.g., Žák and Paterson, 2010; Alasino et al., 2019).
To reconcile these findings in the context of schlieren-bound structures in the TIC, we suggest that the
strain rates producing regional fabrics (of which some may be tectonic) are slower, but longer-lived than
strain rates producing local schlieren-bound structure fabrics.
Despite large variations in composition, size and emplacement depth, schlieren-bound structures are
ubiquitous in plutons, although less well developed in many. Some notable localities include the Tavares
pluton, Brazil (Weinberg et al., 2001), the Skaergaard intrusion, Greenland (Wager and Brown, 1968;
Holness et al., 2017), and the Coastal Maine batholith (Wiebe and Collins 1998), which, in addition to the
TIC, contain a high abundance and diversity of schlieren-bound structures. It suggests that the physical
flow-sorting mechanisms proposed here are an intrinsic feature of magma mushes, and are not primarily
dependent on magma compositions, pressure, or temperature.
5.7 Future directions
Further quantifying the physical and chemical magma properties at the time schlieren-bound structures
formed is an important step in reconstructing magma mush dynamics. In addition, exploring why some
plutons are rich in schlieren-bound structures, while others are poorer, may provide definition on the most
important magma properties or mechanisms to form schlieren-bound structures. Ongoing and future studies
must explore these themes at multiple scales.
171
At the mineral scale, this includes determining which minerals grew in-situ within a structure, and which
were physically sorted into schlieren. Studies have proposed that some structures may represent volatile
pathways (Clarke et al., 2013; Paterson 2009). The role of any exsolved fluid, or gas, that aids or impedes
the formation and preservation of structures also remains unclear and would benefit from quantitative
structural and compositional analysis. Recent studies indicate that magmatic deformation of mushes leaves
a microstructural signature (Holness et al., 2017; Holness, 2018), which could be illuminating in terms of
schlieren formation. For example, do the minerals in schlieren show evidence of crystal-plastic deformation
indicating that they were compacted (e.g., Holness, 2018), or is crystal-repacking a suitable alternative
(Bachmann and Huber, 2019)? What was the porosity of the schlieren cumulate when melts were extracted,
and did they experience secondary grain-boundary adjustments (e.g., Holness, 2018)? Is there a
microstructural record of regional-scale deformation in schlieren-bound structures, and can we quantify the
strain rates forming schlieren fabrics and regional fabrics?
At the outcrop scale, the diversity of structures in the TIC suggests that different magmatic conditions
and/or mechanisms operate within a magma mush at any one time or place. Future directions could
investigate the mechanical and fluid-dynamical relationships between different types of structures, as well
as quantitatively considering the role of the surrounding magma and it’s physical and chemical properties
in determining which type of structure forms.
Across the entire TIC, constraining the different mechanisms of fabric formation is needed to place a
regional context to the internal magmatic schlieren-forming process. Additional statistical analysis of the
regional compilation could resolve why schlieren-bound structures are clustered at this scale, and whether
certain structures dominate certain clusters (e.g., by comparing the structures that are the target structures
nearest neighbor) and resolve the distinction between the distribution of compositionally defined magmatic
structures like K-feldspar megacrysts structures, and schlieren-bound structures.
6. Conclusions
Schlieren layers are locally-sourced from nearby magmas and provide evidence for local magma flow,
mineral-melt separation, and compositional differentiation. Although the diversity of schlieren-bound
structures in the TIC suggests there are multiple mechanisms of formation, schlieren are a common link
between them. Thus, schlieren-forming processes such as hydrogranular flow-sorting, melt migration and
crystal accumulation are considered both widespread and significant in magma mush systems.
Analysis of magmatic structure orientations from outcrop, domainal and regional scales suggests that
magma mushes are highly dynamic environments across multiple length scales, and field observations are
172
remarkably similar to hydrogranular behavior described by the ‘mixing bowl’ models (Schliecher et al.,
2016; Bergantz et al., 2015, 2017). Domainal clustering of structures indicates that mush fluidization is
spatially and temporally heterogeneous.
Outward-younging patterns are likely driven by regional-scale convection and return flow of the magma
mush, driven by the intrusion of new pulses at the center of the magma mush chamber. This pattern,
recorded in all types of schlieren-bound structures studied here, can provide clues to the extent of magma
mixing bowls, and/or the sizes of active domains within them. Schlieren-bound structure orientations are
sensitive to some boundary types, such as local rheological boundaries and nearby internal contacts, but not
to regional stress fields. Thus, the interplay of internal and external forces operating within magma
chambers remains an exciting avenue of future research.
7. Acknowledgements
We acknowledge support from National Science Foundation grants EAR-1624847 and EAR-1019636
awarded to S. Paterson, and a GSA Graduate Student Research Grant Lipman Research Award (2015) to
K. Ardill that funded valuable fieldwork and bulk-rock geochemical analyses. We thank V. Memeti for
providing samples of the KC schlieren and host and for providing field data from the Tuolumne Intrusive
Complex with contributions from R. Miller and J. Žák. We thank G. Aldiss, L. Ardill, M. Cuevas, J. McColl,
K. O’ Rourke, and L. Teruya for assistance in fieldwork, and Yosemite National Park rangers for their
continued support. D. De La Cruz is thanked for assistance in organizing the TIC field data. We thank C.
Barnes for helpful comments during manuscript preparation.
173
Chapter 6: Conclusions
Different aspects of arc behavior explored in the previous chapters, from arc-wide records of evolving
magma plumbing systems, regional magma focusing, magma ascent and eruption, to internal magma
chamber processes, all demonstrate that the Sierra Nevada arc, and the Cretaceous arc in particular, was
dynamic across several orders of space and time. Each study investigated a different aspect of magmatism
in the central Sierra Nevada, to explore the interplay of the above processes, and to leverage the large
datasets that exist for this region. The goal of this study was to describe these processes from a field,
structural, and geochemical approach, investigate their causes, and evaluate their impacts on construction
and evolution of the arc. How these dynamic processes link together in other arc sections with different
attributes (e.g., basement, crustal thickness, lifespan, volume of magmatism, tectonic regime), and the use
of mineral-scale datasets to examine intensive variables in magmatic systems, are both rapidly growing
research fields that provide future directions for this study. Progress on each of the primary research themes
and questions is discussed below:
1. A dynamic arc crustal column
Chapters 2 and 3 concluded that integrating arc spatial patterns, temporal history, and dynamic processes
is a necessary step when using geochemical patterns to identify spatiotemporal arc behaviors. This process,
in turn, provides clues to the types of magma sources that are involved in forming upper-crustal plutons,
hypabyssal intrusions and volcanic rocks. Identifying signals of dynamic processes such as arc flare-ups,
migration, focusing or crustal thickening, required subtraction of spatial ‘static’ patterns, which persist for
the arc’s lifetime. In the case of the Sierra Nevada, this component was the laterally varying oceanic-
continental basement (Kistler, 1990). During the Cretaceous, migration and focusing had distinct spatial-
temporal patterns, making them possible to identify geochemically, and there are multiple lines of evidence
to support rapid crustal thickening (e.g., Saleeby et al., 2003; Profeta et al., 2015; Cao et al., 2016).
However, flare-up geochemical signals remain challenging to identify and interpret, in part because each
Sierran flare-up produced different geochemical patterns. This may suggest that flare-up magmatism may
not be entirely explained by one end-member model, or that the volumes of magmatism, or the composition
of the sources (i.e., mantle) changes through time and modifies the overall arc signal. Recent studies indicate
that the lithospheric or asthenospheric mantle represents a dominant magma source during flare-ups,
between 70-90%, and evolves through time (Martinez-Ardila et al., 2019; Attia et al., 2020). This is
consistent with CSN whole-rock isotope compositions of all ages falling along the mantle array, and with
174
Cretaceous samples overlapping with Cretaceous peridotite and pyroxenite xenolith compositions (e.g.,
Ducea and Saleeby, 1998).
Although the driving mechanisms for flare-ups, migration and focusing remain the focus of intense study
and are debated, this study supports ideas that flare-ups and migration were driven by deep mantle-lower
crustal interactions, while magma focusing in the CSN occurred in the upper-crust, although potentially
spans multiple levels of the crustal column. In contrast to flare-ups and migration, magma focusing has an
important emplacement-level control, possibly related to pluton stress fields (e.g., Karlstrom et al., 2009),
as both un-focused and focused plutons are exposed in the Cretaceous Sierra Nevada. Crustal thickening is
also a transcrustal process, consistent with driving mechanisms of tectonic shortening and magma addition
(Cao et al., 2016). Since each of these processes is expressed in the upper-crustal CSN section, they
represent elements of a transcrustal magma plumbing system, supported further by the finding that each
process operated at ~1-10 km/m.y. rates. The overall effects of these processes were transformative to the
crustal column, replacing 70-90% of the pre-existing crust (including older arc materials) with plutonic and
volcanic material by the end of the Cretaceous flare-up (Ratschbacher et al., 2019).
The compilation revealed that the largest transition in magma geochemistry occurred during Cretaceous
arc magmatism, starting with the end of the Late Jurassic-Early Cretaceous lull period, and continuing to
evolve markedly during the Cretaceous flare-up period. This culminated in the formation of large, upper-
crustal magma chambers, which experienced extensive emplacement-level mixing and outcrop-scale
homogenization (at the mineral scale the mixtures can be quite complex; Solgadi and Sawyer, 2008; Barnes
et al., 2016c) (ch.2, 5). This result was likely made possible by progressive thermal and mechanical
maturation of the entire crustal column (e.g., de Silva et al., 2006) that did not significantly advance during
Triassic or Jurassic arc magmatism. This was potentially a consequence of fewer dynamic processes
operating, such as a thinner crust, or lower magma addition rates during flare-ups (e.g., Paterson and Ducea,
2015). However, estimates of the earlier magma addition rates in the central Sierra Nevada are affected by
the overprinting relationships of each arc, they may increase with additional observations (S. Attia pers.
comm).
Maturation gradually started with the initiation of crustal thickening at ~140 Ma and was sustained and
continually fed by the higher rates of magma addition during the flare-up at 120 Ma. This is expressed in
the CSN by the intrusion of the ~2000 km
2
Fine Gold Intrusive Suite (e.g., Lackey et al., 2012). The
initiation of magma focusing at ~105 Ma, plus increased rates of crustal thickening, marked a transition
from ‘external’ factors playing a significant role in magma composition, such as the crustal/lithospheric
basement, to dominantly ‘internal’ factors where mixing, assimilation, storage and homogenization was
occurring in upper-crustal magma chambers of the TIC (e.g., Kistler et al., 1986; Memeti et al., 2010, 2014;
175
Paterson et al., 2016). Magma-chamber stress fields, ductile response of the surrounding host rock, and
carefully balanced volatile exsolution and magma recharge can promote the growth of magma chambers
over eruption (e.g., Karlstrom et al., 2009; Jellinek and DePaolo, 2003; Huber et al., 2019).
Some studies propose that the geochemical variation of TIC magmas originates, and is largely sourced,
at deep levels in the crust, suggesting that emplacement-level processing was limited (e.g., Gray et al.,
2008; Coleman et al., 2012). Chapter 5 provides evidence that TIC magmas were physically well-mixed
and stirred at the emplacement level, and also supports numerous other lines of evidence for emplacement-
level geochemical mixing and open-system magma processing, such as recycling, contact variations by
magma erosion, and mixing and accumulation of multiple mineral populations, including zircons (e.g.,
Memeti et al., 2010, 2014; Paterson et al., 2011, 2016; Chambers et al., 2020; Werts et al., 2020). These
observations do not preclude the idea that new batches intruding into the TIC, particularly during the later
stages of emplacement (e.g., CP) experienced some processing within magma chambers residing below the
TIC to arrive at emplacement levels already representing a multi-source magma mixture (Paterson et al.,
2016; Oppenheim et al., in review). In contrast, during the early stages of TIC formation, magma batches
were distinct (Memeti et al., 2010; Barnes et al., 2016c).
An additional line of evidence supporting upper-crustal MASH zones, and polybaric MASH zones in
general, is that during the Cretaceous flare-up, the median value of Sr/Y in intermediate compositions
(between 55-68 wt.% SiO
2 and MgO <4 wt.%) not only increases (used for crustal thickness calibration),
but the range dramatically increases as well, and low Sr/Y magmas are produced throughout the duration
of the flare-up. This suggests that both deep (garnet stable) and shallow (plagioclase stable) magma
differentiation was possible (e.g., Mamani et al., 2010). The decreasing Dy/Dy* signal during the
Cretaceous flare-up suggests that amphibole fractionation was significant, which could feasibly occur in a
mid-crustal reservoir (Davidson et al., 2007). Thermally, if these MASH-like processes were operating at
upper-crustal levels (as interpreted here), it is reasonable to interpret that they could be sustained at deeper
levels throughout the crustal column, where the geothermal gradients increase. Thus a multi-MASH
framework provides ample opportunities to create highly complex magma compositions, involving various
contaminants and multiple sources, at multiple stages in the crustal column. This can explain why it remains
challenging to quantify the upper-crustal vs. lower crustal geochemical contributions and the origin of
isotopic signals in the rock record in systems like this.
Other examples where lower-crustal MASH processes have been documented and distinguished using
geochemistry, within a framework of transcrustal magmatism, include the upper-crustal English Peak
plutonic complex, Klamath Mountain province, California (Barnes et al., 2016) and the Valle Fértil section,
Argentina (Walker et al., 2015). In the Altiplano-Puna Volcanic Complex, MASH zone processes are
176
thought to occur in both lower and upper-mid crustal locations due to thermal-mechanical maturation of
the arc crust enabling significant incorporation of lower and upper-crustal host rocks (de Silva et al., 2006;
Walker et al., 2013; Kern et al., 2016). A recent study links geochemical patterns of magma differentiation
to geophysically defined lower-crustal features in active volcanic arcs (Pu et al., 2017).
2. Plutonic-hypabyssal-volcanic magma pathways
The characteristics of volcanic eruptions are closely related to the size and connectivity of the underlying
magma plumbing system (e.g., Kiser et al., 2016). Likewise, information from upper-crustal plumbing
systems, by applying a ‘bottom-up’ approach, can also inform us about volcanism that is no longer exposed.
Chapter 4 investigated hypabyssal intrusions, emplaced in the uppermost levels of the Cretaceous CSN, as
a way to study the volcanic-plutonic connection in an arc where very little of the (once extensive)
Cretaceous volcanic cover remains. Study of CSN hypabyssal intrusions enabled the characterization of
both the physical (shape, size, textures, structures) and chemical components (major, trace elements and
isotopes) of the system. The porphyry intrusive unit was compositionally stratified from dacite to rhyolite,
and this broadly was associated with complex textural layering that was observed in the field. The isotopic
characteristics of the intrusions provided information on the magma sources; crustal contamination was
limited, and consequently the peraluminous rhyolite compositions could have been produced by extreme
fractionation. Thus, the Tioga Pass example is interpreted to represent a steep-sided magma feeder system
where magmas are distilled at shallow levels, promoting differentiation by fractionation. If the intrusive
complex had lost melt to feed an eruption, it must have been dacitic or rhyolitic in composition, matching
or more evolved than the compositions of the plutonic rocks. Sampling of remnant Cretaceous volcanic
rocks in the CSN compilation (ch.3) shows a preference for rhyolites (n=169, median=70.7 wt.% SiO
2,
IQR: 66.6-74.1 wt.% SiO 2). This by no means must represent a model for all the hypabyssal examples that
were identified in the CSN. Future work aims to test which hypabyssal intrusions lost melt to feed eruptions
(i.e., are cumulates), and which likely represent failed eruptions.
3. Evolution of upper-crustal magma chambers
Chapters 2 and 5 illustrated several different geochemical and structural records of the formation of
voluminous, long-lived, mobile magma chambers in the upper crust. As described above, the conclusion
that schlieren-bound structures formed from crystal sorting within a hydrogranular crystal-rich mush has
broader implications for the behavior of magmas within upper-crustal storage regions; namely that they are
able to flow dynamically, mix, erode/recycle material, and deform to form the structures. Furthermore, the
organization of structures into domains, and regional outward-younging patterns, places additional
constraints on the size of the regions which were mobile crystal mushes. In this example, within the
177
Cathedral Peak (CP) unit of the TIC, the minimum size of a mobile magma mush ‘mixing bowl’ (of
Bergantz et al., 2015, 2017), was estimated at 150 km
2
(~1500 km
3
) and may have extended farther (this is
testable with additional domain-scale mapping in the northern CP). The findings of this study, in line with
other field, geochemical, and structural evidence, placed key constraints on possible emplacement
mechanisms of the TIC, precluding small-batch dike or sill models, and supporting a temporal evolution
from sheets to larger pulses such as diapirs (e.g., Miller and Paterson, 2001; Memeti et al., 2010; Cao et al.,
2016).
Another finding that this research supports is that magmatic deformation, recorded in flow banding
(ch.4), as well as magmatic fabrics, and schlieren-bound structures (ch.5), can be related to internal
magmatic processing, with or without influence of external or tectonic forces, and that this distinction is
important (e.g., Paterson et al., 2018). Crustal thickening is another example that demonstrates that
deformation is intimately linked with arc magmatic processes and must be considered when building a
unifying framework for the evolution of transcrustal magma plumbing systems (e.g., Cao et al., 2015).
4. Implications for arc systems
How do Sierran arc dynamic processes compare to other arcs?
The Sierra Nevada arc section represents a well-studied example of the spatiotemporal patterns of
continental arc evolution. Studies of the Andean arc are comparative in scale, resolution, and intensity, with
an emphasis on volcanic records as the arc remains active today (e.g., Hildreth and Moorbath, 1988;
Allmendinger et al., 1997; Kay et al., 2005; de Silva et al., 2006; Mamani et al., 2010; Ward et al., 2017).
The Andean arc during its protracted magmatic history records evidence for all of the same dynamic
processes that occurred in the Mesozoic Sierra Nevada, including flare-ups (e.g. Kirsch et al., 2016;
DeCelles et al., 2015; Martinez-Ardila et al., 2019), multiple episodes of arc migration (e.g., Kay et al.,
2005), magmatic focusing (e.g., Grunder et al., 2008; de Silva et al., 2006; Longo et al., 2010), and extensive
crustal thickening (to >60-70 km in the present day Altiplano) (e.g., Allmendinger et al., 1997; Beck et al.,
1996). Thus, the Andean arc may represent a suitable modern-day analog to Sierran arc activity.
It remains an exciting open question if all these processes similarly occur in oceanic arcs, and at what
rates and timescales. In oceanic arcs, the duration of magmatism is generally much shorter, the crustal
column is thinner (~20-30 km maximum), and there is minor involvement of crustal materials, beyond what
is generated by the arc itself, or brought in via the subducting slab (e.g., Plank and Langmuir, 1993). In
contrast to continental arcs, oceanic arc magmatic episodicity has been proposed to be limited (Jicha and
Jagoutz, 2015). In addition, oceanic (or transitional) arc magma addition rates were estimated to be
considerably higher than long-term continental arc magma addition rates (Jicha and Jagoutz, 2015; Ducea
178
et al., 2017). A recent study of crustal-scale continental arc magma addition rates during flare-up periods
in contrast suggests that the rates are comparable to oceanic arcs, further suggesting that lulls may be unique
to continental arcs (Ratschbacher et al., 2019). A paucity of information on magmatic focusing in oceanic
arc systems (in part due to the requirements for high-density, spatially extensive sampling) precludes
detailed comparison with known continental examples, however, geophysical imaging of the Izu-Bonin arc
section hints that the distribution of magmatism along the arc is focused at the crustal-scale (Kodaira et al.,
2007). In contrast, arc migration is well documented in some island arc systems (e.g., Lesser Antilles;
Germa et al., 2011; Tonga-Kermadec; Yan and Kroenke, 1993; central Japan; Nakamura et al., 2014). Thus
flare-ups and lulls, and potentially magma focusing, may be key to the formation of mature continental arc
systems.
How common are TIC-like magma chambers in the upper-crust?
Each of the studies show that the various magmatic behaviors, whether at the crustal-scale (e.g., flare-
ups, focusing), within the upper-crust (e.g., hypabyssal plumbing systems) or within individual types of
magmatic structures, facilitated or supported (ch.2, 3), required (ch.5), or inferred (ch.4) the formation of
melt-present magma chambers in the upper-crust. In the Late Cretaceous CSN, this is represented by the
1,100 km
2
, nested Tuolumne Intrusive Complex (TIC), which formed particularly large (~100-500 km
2
),
long lived (0.5-1.5 m.y.) magma mush reservoirs during its evolution.
Compared to other intrusive complexes in the central Sierra Nevada, and around the world, the TIC may
be an unusual type of intrusive complex. It was incrementally grown over a protracted crystallization
duration of 10 m.y. (e.g., Coleman et al., 2004; Memeti et al., 2010; Paterson et al., 2011), and a
geochemical history consistent with mixing, fractionation, and recycling at the emplacement level and at
deeper levels of the plumbing system (Kistler et al., 1986; Bateman, 1992; Burgess and Miller, 2008;
Memeti et al., 2014; Paterson et al., 2016; Barnes et al., 2016; Werts et al., 2020; see above). Structurally,
the TIC is highly complex; contacts are highly variable along and across strike, there are five types of
magmatic fabrics (often more than one fabric at a single outcrop) and it contains a unique diversity of
schlieren-bound magmatic structures (Žák et al., 2007; Paterson and Ardill, 2019; ch. 5). Total volumes
may extend to >11,000 km
3
(Paterson et al., 2011). Including the effects of magma recycling would double
this estimate (Paterson et al., 2016).
For these reasons, it may represent an end-member type of magma body, or ‘super-pluton’, which is
complementary to the largest volcanic eruptions on Earth (i.e., super-eruptions) (Paterson et al., 2019;
Paterson et al., forthcoming). It may be comparable (albeit on a smaller scale) to the geophysically-defined
batholith beneath the Southern Rocky Mountain Volcanic Field, the site of the Oligocene Fish Canyon Tuff
179
super-eruption (Lipman, 2007). The batholith has an interpreted volume of 100,000 km
3
(Lipman, 2007).
The TIC may also share similarities to the melt-present, mid-crustal Altiplano Puna Magma Body, in the
Andean arc, with an estimated total volume of 520,000 km
3
(Ward et al., 2017). This designation of the
TIC as a ‘super-pluton’ also relates thermal modeling results of the complex by Paterson et al. (2011) that
produced melt-present magma chambers with lifespans >1 m.y., to crustal-scale thermal-maturation
modeling results described by Karakas et al. (2017), where large-volume mush zones can be sustained in
the upper-crust, given a long (> 1.m.y) duration of magma addition to the lower crust (Karakas et al., 2017
Fig. 4). With crystal accumulation evident in many TIC bulk rock samples, it is evident that the pluton has
lost melt, possibly to higher levels of the magma plumbing system, and could have erupted (Werts et al.,
2020; Barnes et al., 2020). Bentonite beds with K-Ar in biotite ages between 88-90 Ma in the western
interior hint that the erupted products of magma reservoirs, contemporaneous with the age of the TIC, could
have been volumetrically significant (Elder, 1988). Thus, the ability for super-sized magma chambers to
accumulate and segregate melt-rich portions has impacts on the size, shape, composition and longevity of
hypabyssal and volcanic sections of the plumbing system.
While the above characteristics suggests that TIC-like bodies may be unusual in terms of size, or
longevity, these do not appear to be a limiting factor in the formation of melt-present magma reservoirs in
smaller, or more rapidly assembled plutonic complexes (e.g., Coint et al., 2013; Eddy et al., 2016;
Ratschbacher et al., 2018). Nor does this preclude the formation of schlieren-bound magmatic structures in
these types of systems (e.g.,Skaergaard intrusion; Wager and Brown, 1968; Tavares pluton; Weinberg et
al., 2001; Ploumanac’h granite; Barrière, 1981) Some of the unusual characteristics of the TIC may be
related to its position in a focusing zone, forming after the peak of a magmatic flare-up, within a mature
and thick arc crustal column (ch.2, 3).
5. Future work
Dynamic, non-steady state behavior occurs at all scales in the Sierra Nevada arc. There are still many
avenues to explore regarding the driving mechanisms of each process, and the magmatic conditions that
dynamic behavior operates in. The next steps of this study are to provide a mineral-scale perspective on the
physical and chemical volcanic-hypabyssal-plutonic connections and the environment(s) in which
schlieren-bound structure formation is favorable:
Mineral stratigraphy tracing pluton-hypabyssal-volcanic connections
Werts et al. (2020) presented a comprehensive analysis of volcanic and plutonic hornblendes,
concluding that plutonic hornblendes record a melt composition that is no longer representative of the bulk
composition of the fully-solidified rock (i.e., the plutons are cumulates; Barnes et al., 2016a, c, 2020). The
180
volcanic hornblendes that were not in equilibrium with the glass or bulk compositions were interpreted to
have crystallized from less evolved melts, and/or mixed magmas (Werts et al., 2020). This supports the idea
that the crystal cargo within magmas preserves a record that spans multiple levels and multiple magmatic
events/conditions (see also Chambefort et al., 2013; Walker et al., 2013; Watts et al., 2016; Ganne et al.,
2018). In light of these results, mineral studies of hypabyssal intrusions have the potential to capture
transitions between the shallow-level magmatic conditions that generate plutonic cumulates and their
erupted products.
Mineral investigation of magmatic structure formation
Field, structural, and whole-rock geochemical analysis in Chapter 5 concluded that schlieren-bound
structures formed from dynamic magmatic flow, where crystals were sorted by size and density. Several
questions remain regarding schlieren formation, that involve constraining the intensive parameters of the
magmatic system at the time the structures formed. For example, what was the temperature, melt
composition, crystallinity, and effective viscosity of the mush? Is there a significant role of a fluid phase in
this process? Which minerals were physically sorted, and which grew in-situ within the structure? Analysis
of the minerals within magmatic structures, within the context of the overall magmatic evolution of the TIC
(e.g., Werts et al., 2020), is in progress to quantify and further constrain the magmatic conditions of
structure formation.
181
References
Ague, J.J. and Brimhall, G.H., 1988. Regional variations in bulk chemistry, mineralogy, and the
compositions of mafic and accessory minerals in the batholiths of California. Geological Society of
America Bulletin, 100(6), pp.891-911, doi: 10.1130/0016-7606(1988)100<0891:RVIBCM>2.3.CO;2
Ague, J.J., and Brimhall, G.H., 1988b, Magmatic arc asymmetry and distribution of anomalous plutonic
belts in the batholiths of California: Effects of assimilation, crustal thickness, and depth of
crystallization: Geological Society of America Bulletin, v. 100(6), p. 912-927, doi:10.1130/0016
7606(1988)100<0912:MAAADO>2.3.CO;2
Alasino, P.H., Casquet, C., Pankhurst, R.J., Rapela, C.W., Dahlquist, J.A., Galindo, C., Larrovere, M.A.,
Recio, C., Paterson, S.R., Colombo, F. and Baldo, E.G., 2016. Mafic rocks of the Ordovician
Famatinian magmatic arc (NW Argentina): New insights into the mantle contribution. Bulletin, 128(7-
8), pp.1105-1120, doi: 10.1130/B31417.1.
Alasino, P.H., Ardill, K., Stanback, J., Paterson, S.R., Galindo, C. and Leopold, M., 2019. Magmatically
folded and faulted schlieren zones formed by magma avalanching in the Sonora Pass Intrusive Suite,
Sierra Nevada, California. Geosphere, 15(5), pp.1677-1702, doi: 10.1130/GES02070.1.
Albarède, F., 1996. Introduction to geochemical modeling. Cambridge University Press.
Aldiss, G. P.F.B., 2017. The geology of Iron Mountain and Tenaya Peak areas, Sierra National Forest and
Yosemite National Park, Sierra Nevada, California. [BSc thesis], Durham University pp. 75.
Allmendinger, R.W., Jordan, T.E., Kay, S.M. and Isacks, B.L., 1997. The evolution of the Altiplano-Puna
plateau of the Central Andes. Annual review of earth and planetary sciences, 25(1), pp.139-174, doi:
10.1146/annurev.earth.25.1.139.
Allmendinger, R. W., Cardozo, N. C., and Fisher, D., 2013, Structural Geology Algorithms: Vectors &
Tensors: Cambridge, England, Cambridge University Press, 289 pp.
Anders, E. and Grevesse, N., 1989. Abundances of the elements: Meteoritic and solar. Geochimica et
Cosmochimica acta, 53(1), pp.197-214, doi: 10.1016/0016-7037(89)90286-X.
Anderson, J.L., Barth, A.P., Wooden, J.L., and Mazdab, F., 2008, Thermometers and thermobarometers
in granitic systems: Reviews in Mineralogy and Geochemistry, v. 69(1), p.121-142, doi:
10.2138/rmg.2008.69.4
Annen, C., Blundy, J.D. and Sparks, R.S.J., 2006. The genesis of intermediate and silicic magmas in deep
crustal hot zones. Journal of Petrology, 47(3), pp.505-539, doi: 10.1093/petrology/egi084.
Annen, C., Blundy, J.D., Leuthold, J. and Sparks, R.S.J., 2015. Construction and evolution of igneous
bodies: Towards an integrated perspective of crustal magmatism. Lithos, 230, pp.206-221, doi:
10.1016/j.lithos.2015.05.008.
Ardill, K., Paterson, S. and Memeti, V., 2018. Spatiotemporal magmatic focusing in upper-mid crustal
plutons of the Sierra Nevada arc. Earth and Planetary Science Letters, 498, pp.88-100, doi:
10.1016/j.epsl.2018.06.023.
182
Ardill, K.E., Memeti, V., and Paterson, S.R., (in press). Reconstructing the physical and chemical
development of a pluton-porphyry complex in a tectonically re-organized arc crustal section, Tioga
Pass, Sierra Nevada. Lithosphere
Armstrong, R.L., 1988. Mesozoic and early Cenozoic magmatic evolution of the Canadian Cordillera in
Clark, S.P., Burchfiel, B.C., and Suppe, J., eds., Processes in Continental Lithospheric Deformation.
Geological Society of America Special Paper, 218, pp.55-91.
Armstrong, R.L., and Ward, P.L., 1993, Late Triassic to earliest Eocene magmatism in the North
American Cordillera: Implications for the Western Interior Basin, in Caldwell, W.G.E., and Kauffman,
E.G., eds., Evolution of the Western Interior Basin: Geological Association of Canada Special Paper
39, p. 49–72.
Attia, S., 2020. Evolution of lithospheric architecture in arc orogens . [Ph.D. Thesis], University of
Southern California
Attia, S., Paterson, S.R., Wenrong, C., Chapman, A.D., Saleeby, J., Dunne, G., Stevens, C., Memeti, V.,
2017, Late Paleozoic tectonic assembly of the Sierra Nevada prebatholithic framework and Western
Laurentian provenance links based on synthesized detrital zircon geochronology: GSA Special Paper
540, p. 267, doi: 10.1130/2018.2540(12)
Attia, S., Cottle, J.M. and Paterson, S.R., 2020. Erupted zircon record of continental crust formation
during mantle driven arc flare-ups. Geology, 48, doi: 10.1130/G46991.1
Bachmann, O. and Bergantz, G.W., 2004. On the origin of crystal-poor rhyolites: extracted from
batholithic crystal mushes. Journal of Petrology, 45(8), pp.1565-1582, doi: 10.1093/petrology/egh019.
Bachmann, O. and Bergantz, G.W., 2006. Gas percolation in upper-crustal silicic crystal mushes as a
mechanism for upward heat advection and rejuvenation of near-solidus magma bodies. Journal of
Volcanology and Geothermal research, 149(1-2), pp.85-102, doi: 10.1016/j.jvolgeores.2005.06.002.
Bachmann, O., and Huber, C., 2016, Silicic magma reservoirs in the Earth’s crust. American
Mineralogist, v. 101(11), p.2377-2404, doi: 10.2138/am-2016-5675.
Bachmann, O. and Huber, C., 2019. The inner workings of crustal distillation columns; the physical
mechanisms and rates controlling phase separation in silicic magma reservoirs. Journal of Petrology,
60(1), pp.3-18, doi: 10.1093/petrology/egy103.
Bachmann, O., Dungan, M.A. and Lipman, P.W., 2002. The Fish Canyon magma body, San Juan
volcanic field, Colorado: rejuvenation and eruption of an upper-crustal batholith. Journal of Petrology,
43(8), pp.1469-1503, doi: 10.1093/petrology/43.8.1469.
Bachmann, O., Miller, C.F. and De Silva, S.L., 2007, The volcanic–plutonic connection as a stage for
understanding crustal magmatism: Journal of Volcanology and Geothermal Research, v. 167(1), p.1-
23, doi: 10.1016/j.jvolgeores.2007.08.002
Bacon, C.R. and Lanphere, M.A., 2006. Eruptive history and geochronology of Mount Mazama and the
Crater Lake region, Oregon. Geological Society of America Bulletin, 118(11-12), pp.1331-1359,
doi:10.1130/B25906.1
183
Barbarin, B., Dodge, F.C.W., Kistler, R.W. and Bateman, P.C., 1989. Mafic inclusions, aggregates, and
dikes in granitoid rocks, central Sierra Nevada Batholith, California; analytic data. USGS Professional
Paper No. 1899. Doi: 10.3133/b1899.
Barbey, P., 2009. Layering and schlieren in granitoids: A record of interactions between magma
emplacement, crystallization and deformation in growing plutons (The André Dumont medallist
lecture). Geologica Belgica,
Barbey, P., Gasquet, D., Pin, C. and Bourgeix, A.L., 2008. Igneous banding, schlieren and mafic enclaves
in calc-alkaline granites: The Budduso pluton (Sardinia). Lithos, 104(1-4), pp.147-163, doi:
10.1016/j.lithos.2007.12.004.
Barnes, C.G., Coint, N. and Yoshinobu, A., 2016a. Crystal accumulation in a tilted arc batholith.
American Mineralogist, 101(8), pp.1719-1734, doi: 10.2138/am-2016-5404
Barnes, C.G., Ernst, W.G., Berry, R. and Tsujimori, T., 2016b. Petrology and Geochemistry of an Upper
Crustal Pluton: a view into Crustal-scale Magmatism during Arc to Retro-arc Transition. Journal of
Petrology, 57(7), pp.1361-1388, doi: 10.1093/petrology/egw043.
Barnes, C.G., Memeti, V., and Coint, N., 2016c. Deciphering magmatic processes in calc-alkaline plutons
using trace element zoning in hornblende. American Mineralogist, 101, pp. 328–342, doi: 10.2138/am-
2016-5383.
Barnes, C.G., Werts, K., Memeti, V. and Ardill, K., 2020. Most granitoid rocks are cumulates: deductions
from hornblende compositions and zircon saturation. Journal of Petrology, doi:
10.1093/petrology/egaa008.
Barrière M.1981. On curved laminae, graded layers, convection currents and dynamic crystal sorting in
the Ploumanac'h (Brittany) subalkaline granite. Contributions to Mineralogy and Petrology, 77, p.
214–224, doi: 10.1007/BF00373537.
Barth, A.P., Walker, J.D., Wooden, J.L., Riggs, N.R. and Schweickert, R.A., 2011, Birth of the Sierra
Nevada magmatic arc: Early Mesozoic plutonism and volcanism in the east-central Sierra Nevada of
California: Geosphere, v. 7(4), p. 877-897, doi:10.1130/GES00661.1
Barth, A.P., Feilen, A.D.G., Yager, S.L., Douglas, S.R., Wooden, J.L., Riggs, N.R. and Walker, J.D.,
2012. Petrogenetic connections between ash-flow tuffs and a granodioritic to granitic intrusive suite in
the Sierra Nevada arc, California. Geosphere, 8(2), pp.250-264, doi: 10.1130/GES00737.1.
Barth, A.P., Wooden, J.L., Riggs, N.R., Walker, J.D., Tani, K., Penniston‐Dorland, S.C., Jacobson,
C.E., Laughlin, J.A. and Hiramatsu, R., 2018. Marine volcaniclastic record of early arc evolution in
the eastern Ritter Range pendant, central Sierra Nevada, California. Geochemistry, Geophysics,
Geosystems, 19(8), pp.2543-2559, doi: 10.1029/2018GC007456
Bartley, J.M., Coleman, D.S. and Glazner, A.F., 2006. Incremental pluton emplacement by magmatic
crack-seal. Earth and Environmental Science Transactions of the Royal Society of Edinburgh, 97(4),
pp.383-396, doi: 10.1017/S0263593300001528.
Bartley, J.M., Glazner, A.F., Coleman, D.S., and Law, B., 2013. Geometry and emplacement of ladder
dikes in the Cathedral Peak granodiorite, Yosemite National Park. Geological Society of America
Abstracts with Programs, 45 (6), p. 13.
184
Bartley, J.M., Glazner, A.F. and Coleman, D.S., 2018. Dike intrusion and deformation during growth of
the Half Dome pluton, Yosemite National Park, California. Geosphere, 14(3), pp.1283-1297, doi:
10.1130/GES01458.1.
Bartley, J.M, Glazner, A.F., and Law, B., 2019. Ladder structures in the Cathedral Peak granodiorite,
Yosemite: Two-phase flow in fingered dikes? Geological Society of America Abstracts with
Programs, 51 (5), doi: 10.1130/abs/2019AM-338571
Barton, M.D., 1996. Granitic magmatism and metallogeny of southwestern North America. Earth and
Environmental Science Transactions of the Royal Society of Edinburgh, 87(1-2), pp.261-280, doi: doi:
10.1017/S0263593300006672.
Barton, M.D. and Hanson, R.B., 1989. Magmatism and the development of low-pressure metamorphic
belts: Implications from the western United States and thermal modeling. Geological Society of
America Bulletin, 101(8), pp.1051-1065, doi: 10.1130/0016-
7606(1989)101<1051:MATDOL>2.3.CO;2.
Bateman, P.C., 1983. Geologic map of the Tuolumne Meadows quadrangle, Yosemite National Park,
California (No. 1570), doi: 10.3133/gq1570.
Bateman, P.C., 1988. Pre-Tertiary bedrock geologic map of the Mariposa 1 degree x 2 degrees
Quadrangle (No. 87-670). US Geological Survey.
Bateman, P.C., 1992. Plutonism in the central part of the Sierra Nevada batholith, California. USGS
Professional Paper 1483, doi: 10.3133/pp1483.
Bateman, P.C. and Chappell, B.W., 1979. Crystallization, fractionation, and solidification of the
Tuolumne intrusive series, Yosemite National Park, California. Geological Society of America
Bulletin, 90(5), pp.465-482, doi: 10.1130/0016-7606(1979)90<465:CFASOT>2.0.CO;2.
Bateman, P.C. and Dodge, F.W., 1970. Variations of major chemical constituents across the central Sierra
Nevada batholith. Geological Society of America Bulletin, 81(2), pp.409-420.
Bateman, P.C., Chappell, B.W., Kistler, R.W., Peck, D.L., and Busacca, A., 1988, Tuolumne Meadows
Quadrangle, California; Analytic Data: No. 1819, USGS, doi: 10.3133/b1819.
Beane, R. and Wiebe, R.A., 2012, Origin of quartz clusters in Vinalhaven granite and porphyry, coastal
Maine: Contributions to Mineralogy and Petrology, v. 163(6), p.1069-1082, doi: 10.1007/s00410-011-
0717-1
Beard, J.S., 2008. Crystal–melt separation and the development of isotopic heterogeneities in hybrid
magmas. Journal of Petrology, 49(5), pp.1027-1041, doi: 10.1093/petrology/egn015.
Beck, S.L., Zandt, G., Myers, S.C., Wallace, T.C., Silver, P.G. and Drake, L., 1996. Crustal-thickness
variations in the central Andes. Geology, 24(5), pp.407-410, doi: 10.1130/0091-
7613(1996)024<0407:CTVITC>2.3.CO;2.
Bergantz, G.W., 2000, On the dynamics of magma mixing by reintrusion: Implications for pluton
assembly processes: Journal of Structural Geology, v. 22, p. 1297–1309, doi: 10.1016/S0191-
8141(00)00053-5.
185
Bergantz, G.W. and Ni, J., 1999. A numerical study of sedimentation by dripping instabilities in viscous
fluids. International Journal of Multiphase Flow, 25(2), pp.307-320, doi: 10.1016/S0301-
9322(98)00050-0.
Bergantz, G.W., Schleicher, J.M. and Burgisser, A., 2015. Open-system dynamics and mixing in magma
mushes. Nature Geoscience, 8(10), pp.793-796, doi: 10.1038/ngeo2534.
Bergantz, G.W., Schleicher, J.M. and Burgisser, A., 2017. On the kinematics and dynamics of crystal‐rich
systems. Journal of Geophysical Research: Solid Earth, 122(8), pp.6131-6159, doi:
10.1002/2017JB014218.
Best, M.G., Christiansen, E.H., de Silva, S. and Lipman, P.W., 2016. Slab-rollback ignimbrite flareups in
the southern Great Basin and other Cenozoic American arcs: A distinct style of arc volcanism.
Geosphere, 12(4), pp.1097-1135, doi: 10.1130/GES01285.1.
Bhattacharji, S. and Smith, C.H., 1964. Flowage differentiation. Science, 145(3628), pp.150-153, doi:
10.1126/science.145.3628.150.
Blum, J.D. and Erel, Y., 1997. Rb-Sr isotope systematics of a granitic soil chronosequence: The
importance of biotite weathering. Geochimica et Cosmochimica Acta, 61(15), pp.3193-3204, doi:
10.1016/S0016-7037(97)00148-8.
Bonnichsen, B., and Kauffman, D.F., 1987, Physical features of rhyolite lava flows in the Snake River
Plain volcanic province, southwestern Idaho: Geological Society of America Special Papers, v. 212,
p.119-145, doi: 10.1130/SPE212-p119
Boudreau, A., 2011. The evolution of texture and layering in layered intrusions. International Geology
Review, 53(3-4), pp.330-353, doi: 10.1080/00206814.2010.496163.
Boynton, W. V. (1984). Cosmochemistry of the rare earth elements: meteorite studies. In: Henderson, P.
(ed.) Rare Earth Element Geochemistry. Developments in Geochemistry 2. Amsterdam: Elsevier, pp.
63–114.
Bracciali, L., Paterson, S.R., Memeti, V., Rocchi, S., Matzel, J. and Mundil, R., 2008. Build-up of the
Tuolumne Batholith, California: the Johnson Granite Porphyry. In LASI III conference abstracts. pp.
17-18.
Bremond d'Ars, J., Jaupart, C. and Sparks, R.S.J., 1995. Distribution of volcanoes in active margins.
Journal of Geophysical Research: Solid Earth, 100(B10), pp.20421-20432, doi: 10.1029/95JB02153
Brook, C.A., 1977, Stratigraphy and structure of the Saddlebag Lake roof pendant, Sierra Nevada,
California: Geological Society of America Bulletin, v. 88(3), p. 321-334, doi: 10.1130/0016-
7606(1977)88<321:SASOTS>2.0.CO;2
Brounce, M.N., Kelley, K.A. and Cottrell, E., 2014. Variations in Fe3+/∑ Fe of Mariana Arc basalts and
mantle wedge f O2. Journal of Petrology, 55(12), pp.2513-2536, doi: 10.1093/petrology/egu065.
Brown, M., 2007. Crustal melting and melt extraction, ascent and emplacement in orogens: mechanisms
and consequences. Journal of the Geological Society, 164(4), pp.709-730, doi: 10.1144/0016-
76492006-171
186
Burgess, S.D. and Miller, J.S., 2008. Construction, solidification and internal differentiation of a large
felsic arc pluton: Cathedral Peak granodiorite, Sierra Nevada Batholith. Geological Society, London,
Special Publications, 304(1), pp.203-233, doi: 10.1144/SP304.11.
Burgisser, A. and Bergantz, G.W., 2011. A rapid mechanism to remobilize and homogenize highly
crystalline magma bodies. Nature, 471(7337), pp.212-215, doi: 10.1038/nature09799.
Calkins, F. C., 1930, The granitic rocks of the Yosemite region, in Matthes, F. E., ed., Geologic history of
the Yosemite Valley: USGS Professional Paper 160, p. 120—129, doi: 10.3133/pp160.
Candela, P.A., 1991, Physics of aqueous phase evolution in plutonic environments: American
Mineralogist, v. 76, p. 1081-1091
Candela, P.A., 1997. A review of shallow, ore-related granites: textures, volatiles, and ore metals. Journal
of Petrology, v. 38, p.1619-1633, doi: 10.1093/petroj/38.12.1619
Cady, J.W., 1975. Magnetic and gravity anomalies in the Great Valley and western Sierra Nevada
metamorphic belt, California (Vol. 168). Geological Society of America, doi: 10.1130/SPE168-p1.
Cao, W., 2015. Links, tempos and mass balances of cyclic deformation and magmatism in arcs: a case
study on the Mesozoic Sierra Nevada arc integrating geological mapping, geochronology,
geobarometry, strain analyses and numerical simulations. [Ph.D. Thesis], University of Southern
California, 413 p.
Cao, W. and Paterson, S., 2016. A mass balance and isostasy model: Exploring the interplay between
magmatism, deformation and surface erosion in continental arcs using central Sierra Nevada as a case
study. Geochemistry, Geophysics, Geosystems, 17(6), pp.2194-2212, doi: 10.1002/2015GC006229.
Cao, W., Paterson, S.R., Memeti, V., Mundil, R., Anderson, J.L., and Schmidt, K., 2015, Tracking
paleodeformation fields in the Mesozoic central Sierra Nevada arc: Implications for intra-arc cyclic
deformation and arc tempos: Lithosphere, v. 7, p. 296–320, doi: 10.1130/L389.1
Cao, W., Kaus, B.J. and Paterson, S., 2016. Intrusion of granitic magma into the continental crust
facilitated by magma pulsing and dike‐diapir interactions: Numerical simulations. Tectonics, 35(6),
pp.1575-1594, doi: 10.1002/2015TC004076.
Cao, W.., Paterson, S., Saleeby, J. and Zalunardo, S., 2016. Bulk arc strain, crustal thickening, magma
emplacement, and mass balances in the Mesozoic Sierra Nevada arc. Journal of Structural Geology,
84, pp.14-30, doi: 10.1016/j.jsg.2015.11.002.
Cao, W., Lee, C.T.A. and Lackey, J.S., 2017. Episodic nature of continental arc activity since 750 Ma: A
global compilation. Earth and Planetary Science Letters, 461, pp.85-95, doi:
10.1016/j.epsl.2016.12.044.
Carrara, A., Burgisser, A. and Bergantz, G.W., 2019. Lubrication effects on magmatic mush dynamics.
Journal of Volcanology and Geothermal Research, 380, pp.19-30, doi:
10.1016/j.jvolgeores.2019.05.008.
Cashman, K.V., Sparks, R.S.J. and Blundy, J.D., 2017. Vertically extensive and unstable magmatic
systems: a unified view of igneous processes. Science, v. 355, doi: 10.1126/science.aag3055.
187
Cecil, M.R., Rotberg, G.L., Ducea, M.N., Saleeby, J.B. and Gehrels, G.E., 2012. Magmatic growth and
batholithic root development in the northern Sierra Nevada, California. Geosphere, 8(3), pp.592-606,
doi: 10.1130/GES00729.1.
Cecil, M.R., Rusmore, M.E., Gehrels, G.E., Woodsworth, G.J., Stowell, H.H., Yokelson, I.N., Chisom,
C., Trautman, M. and Homan, E., 2018. Along‐strike variation in the magmatic tempo of the Coast
Mountains batholith, British Columbia, and implications for processes controlling episodicity in arcs.
Geochemistry, Geophysics, Geosystems, 19(11), pp.4274-4289, doi: 10.1029/2018GC007874.
Cecil, M.R., Ferrer, M.A., Riggs, N.R., Marsaglia, K., Kylander-Clark, A., Ducea, M.N. and Stone, P.,
2019. Early arc development recorded in Permian–Triassic plutons of the northern Mojave Desert
region, California, USA. Bulletin, 131(5-6), pp.749-765, doi: 10.1130/B31963.1.
Chambefort, I., Dilles, J.H. and Longo, A.A., 2013. Amphibole geochemistry of the Yanacocha
Volcanics, Peru: Evidence for diverse sources of magmatic volatiles related to gold ores. Journal of
Petrology, 54(5), pp.1017-1046, doi: 10.1093/petrology/egt004.
Chambers, M., Memeti, V., Eddy, M.P. and Schoene, B., 2020. Half a million years of magmatic history
recorded in a K-feldspar megacryst of the Tuolumne Intrusive Complex, California, USA. Geology,
48(4), pp.400-404, doi: 10.1130/G46873.1.
Chapman, A.D., Saleeby, J.B., Wood, D.J., Piasecki, A., Kidder, S., Ducea, M.N., and Farley, K.A.,
2012, Late Cretaceous gravitational collapse of the southern Sierra Nevada batholith, California:
Geosphere, v. 8, p. 314-341, doi: 10.1130/GES00740.1
Chapman, J.B. and Ducea, M.N., 2019. The role of arc migration in Cordilleran orogenic cyclicity.
Geology, 47(7), pp.627-631, doi: 10.1130/G46117.1.
Chapman, J.B., Ducea, M.N., DeCelles, P.G. and Profeta, L., 2015. Tracking changes in crustal thickness
during orogenic evolution with Sr/Y: An example from the North American Cordillera. Geology,
43(10), pp.919-922, doi: 10.1130/G36996.1.
Chapman, J.B., Ducea, M.N., Kapp, P., Gehrels, G.E. and DeCelles, P.G., 2017. Spatial and temporal
radiogenic isotopic trends of magmatism in Cordilleran orogens. Gondwana Research, 48, pp.189-204,
doi: 10.1016/j.gr.2017.04.019.
Chen, J.H. and Moore, J.G., 1982. Uranium‐lead isotopic ages from the Sierra Nevada Batholith,
California. Journal of Geophysical Research: Solid Earth, 87(B6), pp.4761-4784, doi:
10.1029/JB087iB06p04761
Chen, J.H. and Tilton, G.R., 1991. Applications of lead and strontium isotopic relationships to the
petrogenesis of granitoid rocks, central Sierra Nevada batholith, California. Geological Society of
America Bulletin, 103(4), pp.439-447, doi: 10.1130/0016-7606(1991)103<0439:AOLASI>2.3.CO;2.
Chesterman, C.W., 1975. Geology of the Matterhorn Peak 15-minute Quadrangle, Mono and Tuolumne
Counties, California. California Division of Mines and Geology. Map sheet 22.
Chin, E.J., Lee, C.T.A., Luffi, P. and Tice, M., 2012. Deep lithospheric thickening and refertilization
beneath continental arcs: Case study of the P, T and compositional evolution of peridotite xenoliths
from the Sierra Nevada, California. Journal of Petrology, 53(3), pp.477-511, doi:
10.1093/petrology/egr069.
188
Chin, E.J., Lee, C.T.A. and Barnes, J.D., 2014. Thickening, refertilization, and the deep lithosphere filter
in continental arcs: Constraints from major and trace elements and oxygen isotopes. Earth and
Planetary Science Letters, 397, pp.184-200, doi: 10.1016/j.epsl.2014.04.022.
Church, T. M., 1979, Marine Barite, in Burns, R. G. ed., Marine minerals. Reviews in mineralogy, v6,
Mineralogical Society of America, p. 170-210.
Clarke, D.B., 2003. Exploded Xenoliths, layered Granodiorites, and Chaotic Schlieren associated with the
eastern contact of the South Mountain Batholith. Atlantic Geoscience Society.
Clarke D.B., Clarke G.K.C., 1998. Layered granodiorites at Chebutco Head, South Mountain batholith,
Nova Scotia. Journal of Structural Geology, 20, pp. 1305-1324, doi:
Clarke, D.B., Grujic, D., McCuish, K.L., Sykes, J.C. and Tweedale, F.M., 2013. Ring schlieren:
description and interpretation of field relations in the Halifax Pluton, South Mountain Batholith, Nova
Scotia. Journal of Structural Geology, 51, pp.193-205, doi: 10.1016/j.jsg.2013.01.009.
Clemens, J.D., Stevens, G., and Bryan, S.E., 2020, Conditions during the formation of granitic magmas
by crustal melting – Hot or cold; drenched, damp or dry? Earth Science Reviews, v. 200, doi:
10.1016/j.earscirev.2019.102982.
Clift, P.D., Pavlis, T., DeBari, S.M., Draut, A.E., Rioux, M. and Kelemen, P.B., 2005. Subduction erosion
of the Jurassic Talkeetna-Bonanza arc and the Mesozoic accretionary tectonics of western North
America. Geology, 33(11), pp.881-884, doi: 10.1130/G21822.1.
Cloos, E., 1936. Der Sierra Nevada Pluton in Californien. Neues Jahrbuch für Mineralogie. Geol.
Paleontol. 76, pp. 355–450.
Coats, R.R., 1936. Primary banding in basic plutonic rocks. The Journal of Geology, 44(3), pp.407-419,
doi: 10.1086/624432.
Coint, N., Barnes, C.G., Yoshinobu, A.S., Chamberlain, K.R. and Barnes, M.A., 2013. Batch-wise
assembly and zoning of a tilted calc-alkaline batholith: Field relations, timing, and compositional
variation. Geosphere, 9(6), pp.1729-1746, doi: 10.1130/GES00930.1.
Coint, N., Barnes, C.G., Yoshinobu, A.S., Barnes, M.A. and Buck, S., 2013. Use of trace element
abundances in augite and hornblende to determine the size, connectivity, timing, and evolution of
magma batches in a tilted batholith. Geosphere, 9(6), pp.1747-1765, doi: 10.1130/GES00931.1
Coleman, D.S. and Glazner, A.F., 1997. The Sierra Crest magmatic event: Rapid formation of juvenile
crust during the Late Cretaceous in California. International Geology Review, 39(9), pp.768-787, doi:
10.1080/00206819709465302.
Coleman, D.S., Glazner, A.F. and Frost, T.P., 1992. Evidence from the Lamarck Granodiorite for rapid
Late Cretaceous crust formation in California. Science, 258(5090), pp.1924-1926, doi:
10.1126/science.258.5090.1924.
Coleman, D.S., Gray, W. and Glazner, A.F., 2004. Rethinking the emplacement and evolution of zoned
plutons: Geochronologic evidence for incremental assembly of the Tuolumne Intrusive Suite,
California. Geology, 32(5), pp.433-436, doi: 10.1130/G20220.1.
189
Coleman, D.S., Bartley, J.M., Glazner, A.F. and Pardue, M.J., 2012. Is chemical zonation in plutonic
rocks driven by changes in source magma composition or shallow-crustal differentiation? Geosphere,
8(6), pp.1568-1587, doi: 10.1130/GES00798.1.
Colgan, J.P., John, D.A., Henry, C.D. and Watts, K.E., 2018. Insights into the emplacement of upper-
crustal plutons and their relationship to large silicic calderas, from field relationships, geochronology,
and zircon trace element geochemistry in the Stillwater–Clan Alpine caldera complex, western
Nevada, USA. Journal of Volcanology and Geothermal Research, v. 349, p.163-176, doi:
10.1016/j.jvolgeores.2017.10.015.
Coney, P.J. and Reynolds, S.J., 1977. Cordilleran benioff zones. Nature, 270(5636), p.403, doi:
10.1038/270403a0.
Corry, C.E., 1988. Laccoliths: mechanics of emplacement and growth, Geological Society of America
Special Papers, v. 220.
Davidson, J., Turner, S., Handley, H., Macpherson, C. and Dosseto, A., 2007. Amphibole “sponge” in arc
crust?. Geology, 35(9), pp.787-790, doi: 10.1130/G23637A.1.
Davidson, J., Turner, S. and Plank, T., 2013. Dy/Dy*: variations arising from mantle sources and
petrogenetic processes. Journal of Petrology, 54(3), pp.525-537, doi: 10.1093/petrology/egs076.
Davis, M., Koenders, M.A., and Petford, N., 2007, Vibro-agitation of chambered magma: Journal of
Volcanology and Geothermal Research, v. 167, p. 24–36, doi: 10.1016/j.jvolgeores.2007.07.012.
DeCelles, P.G., Ducea, M.N., Kapp, P. and Zandt, G., 2009. Cyclicity in Cordilleran orogenic systems.
Nature Geoscience, 2(4), p.251, doi: 10.1038/ngeo469.
DeCelles, P.G., Zandt, G., Beck, S.L., Currie, C.A., Ducea, M.N., Kapp, P., Gehrels, G.E., Carrapa, B.,
Quade, J., and Schoenbohm, L.M., 2015, Cyclical orogenic processes in the Cenozoic central Andes,
in DeCelles, P.G., Ducea, M.N., Carrapa, B., and Kapp, P.A., eds., Geodynamics of a Cordilleran
Orogenic System: The Central Andes of Argentina and Northern Chile: Geological Society of
America Memoir 212, p. 459–490, doi:10.1130/2015.1212(22).
Decker, M., Schwartz, J.J., Stowell, H.H., Klepeis, K.A., Tulloch, A.J., Kitajima, K., Valley, J.W. and
Kylander-Clark, A.R.C., 2017. Slab-triggered arc flare-up in the Cretaceous Median Batholith and the
growth of lower arc crust, Fiordland, New Zealand. Journal of Petrology, 58(6), pp.1145-1171, doi:
10.1093/petrology/egx049.
Deering, C.D. and Bachmann, O., 2010. Trace element indicators of crystal accumulation in silicic
igneous rocks. Earth and Planetary Science Letters, 297(1-2), pp.324-331, doi:
10.1016/j.epsl.2010.06.034.
Deering, C.D., Keller, B., Schoene, B., Bachmann, O., Beane, R. and Ovtcharova, M., 2016. Zircon
record of the plutonic-volcanic connection and protracted rhyolite melt evolution. Geology, 44(4),
pp.267-270, doi: 10.1130/G37539.1.
Degruyter, W., Parmigiani, A., Huber, C. and Bachmann, O., 2019. How do volatiles escape their shallow
magmatic hearth? Philosophical Transactions of the Royal Society A, 377(2139), doi:
10.1098/rsta.2018.0017.
190
DePaolo, D.J., 1981. A neodymium and strontium isotopic study of the Mesozoic calc‐alkaline granitic
batholiths of the Sierra Nevada and Peninsular Ranges, California. Journal of Geophysical Research:
Solid Earth, 86(B11), pp.10470-10488, doi: 10.1029/JB086iB11p10470.
de Silva, S.L. and Gosnold, W.D., 2007. Episodic construction of batholiths: insights from the
spatiotemporal development of an ignimbrite flare-up. Journal of Volcanology and Geothermal
Research, 167(1), pp.320-335, doi: 10.1016/j.jvolgeores.2007.07.015.
de Silva, S.L., and Gregg, P.M., 2014, Thermomechanical feedbacks in magmatic systems: Implications
for growth, longevity, and evolution of large caldera-forming magma reservoirs and their
supereruptions. Journal of Volcanology and Geothermal Research v. 282 pp.77-91, doi:
10.1016/j.jvolgeores.2014.06.001
de Silva, S.L., Zandt, G., Trumbull, R., Viramonte, J.G., Salas, G. and Jiménez, N., 2006. Large
ignimbrite eruptions and volcano-tectonic depressions in the Central Andes: a thermomechanical
perspective. Geological society, London, special publications, 269(1), pp.47-63, doi:
10.1144/GSL.SP.2006.269.01.04
de Silva, S.L., Riggs, N.R. and Barth, A.P., 2015. Quickening the pulse: Fractal tempos in continental arc
magmatism. Elements, 11(2), pp.113-118, doi: 10.2113/gselements.11.2.113.
Dickinson, W.R., 1975, Potash-depth (K-h) relations in continental margin and intra-oceanic magmatic
arcs: Geology, v. 3, p. 53–56, doi: 10.1130/0091-7613(1975)3<53:PKRICM>2.0.CO;2.
Dickinson, W.R. and Gehrels, G.E., 2009. Use of U–Pb ages of detrital zircons to infer maximum
depositional ages of strata: a test against a Colorado Plateau Mesozoic database. Earth and Planetary
Science Letters, 288(1-2), pp.115-125, doi: 10.1016/j.epsl.2009.09.013
Dietl, C., De Wall, H. and Finger, F., 2010. Tube-like schlieren structures in the Fürstenstein Intrusive
Complex (Bavarian Forest, Germany): Evidence for melt segregation and magma flow at intraplutonic
contacts. Lithos, 116(3-4), pp.321-339, doi: 10.1016/j.lithos.2009.11.011.
Dodge, F.C.W., and Calk, L.C., 1986, Lake Eleanor quadrangle, central Sierra Nevada, California
Analytic data: U.S. Geological Survey Bulletin 1565, 20 p, doi: 10.3133/gq1639.
Doe, B.R. and Delevaux, M.H., 1973. Variations in lead-isotopic compositions in Mesozoic granitic rocks
of California: A preliminary investigation. Geological Society of America Bulletin, 84(11), pp.3513-
3526, doi: 10.1130/0016-7606(1973)84<3513:VILCIM>2.0.CO;2.
Ducea, M. N., 1998. A petrologic investigation of deep-crustal and upper-mantle xenoliths from the
Sierra Nevada, California: constraints on lithospheric composition beneath continental arcs and the
origin of Cordilleran batholiths. [Ph.D. thesis], California Institute of Technology, Pasadena
Ducea, M.N., 2001. The California arc: Thick granitic batholiths, eclogitic residues, lithospheric-scale
thrusting, and magmatic flare-ups. GSA today, 11(11), pp.4-10, doi: 10.1130/1052-
5173(2001)011<0004:TCATGB>2.0.CO;2
Ducea, M.N, and Barton, M.D., 2007, Igniting flare-up events in Cordilleran arcs. Geology, 35, p. 1047-
1050, doi: 10.1130/G23898A.1.
191
Ducea, M. and Saleeby, J., 1998. Crustal recycling beneath continental arcs: silica-rich glass inclusions in
ultramafic xenoliths from the Sierra Nevada, California. Earth and Planetary Science Letters, 156(1-2),
pp.101-116, doi: 10.1016/S0012-821X(98)00021-1.
Ducea MN, Paterson SR, DeCelles P.G., 2015. High-volume magmatic events in subduction systems.
Elements, 11, p. 99-104, doi: 10.2113/gselements.11.2.99
Ducea, M.N., Saleeby, J.B. and Bergantz, G., 2015. The architecture, chemistry, and evolution of
continental magmatic arcs. Annual Review of Earth and Planetary Sciences, 43, pp.299-331, doi:
10.1146/annurev-earth-060614-105049.
Ducea, M.N., Bergantz, G.W., Crowley, J.L. and Otamendi, J., 2017. Ultrafast magmatic buildup and
diversification to produce continental crust during subduction. Geology, 45(3), pp.235-238, doi:
10.1130/G38726.1.
Dumitru, T.A., Gans, P.B., Foster, D.A. and Miller, E.L., 1991. Refrigeration of the western Cordilleran
lithosphere during Laramide shallow-angle subduction. Geology, 19(11), pp.1145-1148, doi:
Economos, R.C., Memeti, V., Paterson, S.R., Miller, J.S., Erdmann, S. and Žák, J., 2009. Causes of
compositional diversity in a lobe of the Half Dome granodiorite, Tuolumne Batholith, Central Sierra
Nevada, California. Earth and Environmental Science Transactions of the Royal Society of Edinburgh,
100(1-2), pp.173-183, doi: 10.1017/S1755691009016065.
Eddy M. P., Bowring S. A., Miller R. B., Tepper J. H., 2016. Rapid assembly and crystallization of a
fossil large-volume silicic magma chamber. Geology 44, 331–334, doi: 10.1130/G37631.1.
Eichelberger, J.C., Carrigan, C.R., Westrich, H.R. and Price, R.H., 1986. Non-explosive silicic volcanism.
Nature, 323(6089), p.598, doi: 10.1038/323598a0
Elder, W.P., 1988. Geometry of Upper Cretaceous bentonite beds: Implications about volcanic source
areas and paleowind patterns, western interior, United States. Geology, 16, pp.835–838, doi:
10.1130/0091-7613(1988)016<0835:GOUCBB>2.3.CO;2
Estep, J. and Dufek, J., 2012. Substrate effects from force chain dynamics in dense granular flows.
Journal of Geophysical Research: Earth Surface, 117(F1), doi: 10.1029/2011JF002125.
Evernden, J. F.; and Kistler, R. W., 1970. Chronology of emplacement of Mesozoic batholithic
complexes in California and western Nevada: U.S. Geol. Surv., Prof. Pap. 623, 42 p., doi:
10.3133/pp623.
Farner, M.J., Lee, C.T.A. and Mikus, M.L., 2018. Geochemical signals of mafic-felsic mixing: Case study
of enclave swarms in the Bernasconi Hills pluton, California. GSA Bulletin, 130(3-4), pp.649-660,
doi: 10.1130/B31760.1.
Fiske, R.S. and Tobisch, O.T., 1978. Paleogeographic significance of volcanic rocks of the Ritter Range
pendant, central Sierra Nevada, California. Pacific Coast Paleogeography Symposium 2, Mesozoic
Paleogeography of the Western United States, pp. 209-221
Fiske, R. S., and Tobisch, O.T., 1994. Middle Cretaceous ash-flow tuff and caldera-collapse deposit in the
Minarets Caldera, east-central Sierra Nevada, California: Geological Society of America Bulletin,
106(5), pp.582-593, doi: 10.1130/0016-7606(1994)106<0582:MCAFTA>2.3.CO;2.
192
Foley, B.J., Ball, E.N., Fischer, G.C., Thompson, J.M., Memeti, V., Pignotta, G.S., Anderson, J.L.,
Paterson, S.R., Matzel, J., Mundil, R., 2007, Downward ductile displacement of volcanic crust during
pluton emplacement in the central Sierra Nevada: Undergraduate Team Research at USC: Geological
Society of America Cordilleran Section meeting, No. 31-23
Frost T.P., 1987, Sample localities, descriptions, major- and trace-element abundances from the Lamarck
Granodiorite and associated mafic rocks, eastern Sierra Nevada, California: U.S. Geological Survey
Open File Report p. 87–193, doi: 10.3133/ofr87193.
Ganne, J., Bachmann, O. and Feng, X., 2018. Deep into magma plumbing systems: Interrogating the
crystal cargo of volcanic deposits. Geology, 46(5), pp.415-418, doi: 10.1130/G39857.1.
Garibaldi, N., Tikoff, B., Schaen, A.J. and Singer, B.S., 2018. Interpreting granitic fabrics in terms of
rhyolitic melt segregation, accumulation, and escape via tectonic filter pressing in the Huemul pluton,
Chile. Journal of Geophysical Research: Solid Earth, 123(10), pp.8548-8567, doi:
10.1029/2018JB016282.
Gehrels, G., 2014, Detrital zircon U-Pb geochronology applied to tectonics: Annual Review of Earth and
Planetary Sciences, v. 42, p.127-149, doi: 10.1146/annurev-earth-050212-124012
Gehrels, G.E., Valencia, V.A. and Ruiz, J., 2008. Enhanced precision, accuracy, efficiency, and spatial
resolution of U‐Pb ages by laser ablation–multicollector–inductively coupled plasma–mass
spectrometry. Geochemistry, Geophysics, Geosystems, v. 9, doi: 10.1029/2007GC001805
Gehrels, G., Rusmore, M., Woodsworth, G., Crawford, M., Andronicos, C., Hollister, L., Patchett, J.,
Ducea, M., Butler, R., Klepeis, K. and Davidson, C., 2009. U-Th-Pb geochronology of the Coast
Mountains batholith in north-coastal British Columbia: Constraints on age and tectonic evolution.
Geological Society of America Bulletin, 121(9-10), pp.1341-1361, doi: 10.1130/B26404.1.
Gelman, S.E., Deering, C.D., Bachmann, O., Huber, C. and Gutierrez, F.J., 2014. Identifying the crystal
graveyards remaining after large silicic eruptions. Earth and Planetary Science Letters, 403, pp.299-
306, doi: 10.1016/j.epsl.2014.07.005.
Germa, A., Quidelleur, X., Labanieh, S., Chauvel, C. and Lahitte, P., 2011. The volcanic evolution of
Martinique Island: Insights from K–Ar dating into the Lesser Antilles arc migration since the
Oligocene. Journal of Volcanology and Geothermal Research, 208(3-4), pp.122-135, doi:
10.1016/j.jvolgeores.2011.09.007.
Glazner, A.F., 2014. Magmatic life at low Reynolds number. Geology, 42(11), pp.935-938, doi:
10.1130/G36078.1.
Glazner, A.F., Bartley, J.M., Law, B., and Coleman, D.S., 2012. Ladder dikes, crazy geochemistry, and
liquid immiscibility (?) in otherwise sane granites. Geological Society of America Abstracts with
Programs, 44 (3) (2012), p. 21
Glazner, A.F., Coleman, D.S. and Mills, R.D., 2015, The volcanic-plutonic connection: in Breitkreuz, C.
and Rocchi, S., eds., Physical geology of shallow magmatic systems, Advances in Volcanology, doi:
10.1007/11157_2015_11.
Gonnermann, H.M., and Manga, M., 2005, Flow banding in obsidian: A record of evolving textural
heterogeneity during magma deformation: Earth and Planetary Science Letters, v. 236(1), p.135-147,
doi: 10.1016/j.epsl.2005.04.031
193
Graham, J. 2012. Yosemite National Park: Geologic resources inventory report. Natural Resource Report
NPS/NRSS/GRD/NRR—2012/560. National Park Service, Fort Collins, Colorado.
Gray, W., Glazner, A.F., Coleman, D.S. and Bartley, J.M., 2008. Long-term geochemical variability of
the Late Cretaceous Tuolumne intrusive suite, central Sierra Nevada, California. Geological Society,
London, Special Publications, 304(1), pp.183-201, doi: 10.1144/SP304.10.
Greene, D.C., 1995, The stratigraphy, structure, and regional tectonic significance of the Northern Ritter
Range pendant, eastern Sierra Nevada. California [Ph.D. thesis]: Reno, University of Nevada, 161 p.
Greene, D. C., and Schweickert, R. A., 1995, The Gem Lake Shear Zone – Cretaceous dextral
transpression in the northern Ritter Range pendant, eastern Sierra-Nevada California: Tectonics, v.
14(4), p 945– 961, doi: 10.1029/95TC01509
Greene, D.C., Hoffman, C.F., Klemetti, E.W., Toth, C., and Worm, T.J., 2017, Pervasive folding during
Cretaceous deformation in the Mineral King pendant, southern Sierra Nevada, California, Geological
Society of America, Abstracts with Programs, v. 49, n. 6
Gregg, P.M., De Silva, S.L. and Grosfils, E.B., 2013. Thermomechanics of shallow magma chamber
pressurization: Implications for the assessment of ground deformation data at active volcanoes. Earth
and Planetary Science Letters, 384, pp.100-108, doi: 10.1016/j.epsl.2013.09.040
Griffiths, R.W., 1986. Thermals in extremely viscous fluids, including the effects of temperature-
dependent viscosity. Journal of Fluid Mechanics, 166, pp.115-138, doi:
10.1017/S002211208600006X.
Grunder, A.L., Klemetti, E.W., Feeley, T.C., and McKee, C.M., 2008. Eleven million years of arc
volcanism at the Aucanquilcha Volcanic Cluster, northern Chilean Andes: implications for the life
span and emplacement of plutons: Transactions of the Royal Society of Edinburgh: Earth Sciences, 97,
pp.415–436, doi: 10.1017/S0263593300001541
Hanson, R.B., Sorensen, S.S., Barton, M.D. and Fiske, R.S., 1993. Long-term evolution of fluid–rock
interactions in magmatic arcs: evidence from the Ritter Range pendant, Sierra Nevada, California, and
numerical modeling. Journal of Petrology, 34(1), pp.23-62, doi: 10.1093/petrology/34.1.23.
Hartman, S.M., Paterson, S.R., Holk, G.J. and Kirkpatrick, J.D., 2018. Structural and hydrothermal
evolution of a strike-slip shear zone during a ductile-brittle transition, Sierra Nevada, CA. Journal of
Structural Geology, 113, pp.134-154, doi: 10.1016/j.jsg.2018.05.010
Hildreth, W., 1981. Gradients in silicic magma chambers: implications for lithospheric magmatism.
Journal of Geophysical Research: solid earth, 86(B11), pp.10153-10192, doi:
10.1029/JB086iB11p10153.
Hildreth, W. and Moorbath, S., 1988. Crustal contributions to arc magmatism in the Andes of central
Chile. Contributions to mineralogy and petrology, 98(4), pp.455-489, doi: 10.1007/BF00372365.
Hillhouse, J.W. and Grommé, S., 2011. Updated paleomagnetic pole from Cretaceous plutonic rocks of
the Sierra Nevada, California: Tectonic displacement of the Sierra Nevada block. Lithosphere, 3(4),
pp.275-288, doi: 10.1130/L142.1.
194
Hochstaedter, A.G., Gill, J.B., Taylor, B., Ishizuka, O., Yuasa, M. and Monta, S., 2000. Across‐arc
geochemical trends in the Izu‐Bonin arc: Constraints on source composition and mantle melting.
Journal of Geophysical Research: Solid Earth, 105(B1), pp.495-512, doi: 10.1029/1999JB900125
Hochstaedter, A., Gill, J., Peters, R., Broughton, P., Holden, P. and Taylor, B., 2001. Across‐arc
geochemical trends in the Izu‐Bonin arc: Contributions from the subducting slab. Geochemistry,
Geophysics, Geosystems, 2(7), doi: 10.1029/2000GC000105.
Hodge, K.F., Carazzo, G., Montague, X. and Jellinek, A.M., 2012. Magmatic structures in the Tuolumne
Intrusive Suite, California: a new model for the formation and deformation of ladder dikes.
Contributions to Mineralogy and Petrology, 164(4), pp.587-600, doi: 10.1007/s00410-012-0760-6.
Holness, M.B., 2018. Melt segregation from silicic crystal mushes: a critical appraisal of possible
mechanisms and their microstructural record. Contributions to Mineralogy and Petrology, 173(6),
p.48, doi: 10.1007/s00410-018-1465-2.
Holness, M.B., Vukmanovic, Z. and Mariani, E., 2017. Assessing the role of compaction in the formation
of adcumulates: a microstructural perspective. Journal of Petrology, 58(4), pp.643-673, doi:
10.1093/petrology/egx037.
Holt, A.F., Becker, T.W. and Buffett, B.A., 2015. Trench migration and overriding plate stress in
dynamic subduction models. Geophysical Journal International, 201(1), pp.172-192, doi:
10.1093/gji/ggv011.
Huber, N.K., Bateman, P.C. and Wahrhaftig, C., 1989, Geologic map of Yosemite National Park and
vicinity, California: No. 1874, U.S. Geological Survey, scale 1: 125,000.
Huber, C., Townsend, M., Degruyter, W. and Bachmann, O., 2019. Optimal depth of subvolcanic magma
chamber growth controlled by volatiles and crust rheology. Nature Geoscience, 12(9), pp.762-768,
doi: 10.1038/s41561-019-0415-6.
Hughes, G.R. and Mahood, G.A., 2008. Tectonic controls on the nature of large silicic calderas in
volcanic arcs. Geology, 36(8), pp.627-630, doi: 10.1130/G24796A.1.
Hurai, V., Simon, K., Wiechert, U., Hoefs, J., Konečný, P., Huraiová, M., Pironon, J. and Lipka, J., 1998.
Immiscible separation of metalliferous Fe/Ti-oxide melts from fractionating alkali basalt: P-T-f O2
conditions and two-liquid elemental partitioning. Contributions to Mineralogy and Petrology, 133(1-
2), pp.12-29, doi: 10.1007/s004100050433.
Hutton, J. 1788. Theory of the Earth; or an investigation of the laws observable in the composition,
dissolution, and restoration of land upon the globe. Transactions of the Royal Society of Edinburgh, 1,
209–304.
Irvine TN, Andersen JCØ, Brooks CK (1998) Included blocks (and blocks within blocks) in the
Skaergaard intrusion: geologic relations and the origins of rhythmic modally graded layers. Geological
Society of America Bulletin, 110, pp. 1398–1447, doi: 10.1130/0016-
7606(1998)110<1398:IBABWB>2.3.CO;2.
Jackson, M.D., Blundy, J. and Sparks, R.S.J., 2018. Chemical differentiation, cold storage and
remobilization of magma in the Earth’s crust. Nature, 564(7736), pp.405-409, doi: 10.1038/s41586-
018-0746-2.
195
Jacques, G., Hoernle, K., Gill, J., Hauff, F., Wehrmann, H., Garbe-Schönberg, D., van den Bogaard, P.,
Bindeman, I. and Lara, L.E., 2013. Across-arc geochemical variations in the Southern Volcanic Zone,
Chile (34.5–38.0 S): constraints on mantle wedge and slab input compositions. Geochimica et
Cosmochimica Acta, 123, pp.218-243, doi: 10.1016/j.gca.2013.05.016.
Jellinek, A.M. and DePaolo, D.J., 2003. A model for the origin of large silicic magma chambers:
precursors of caldera-forming eruptions. Bulletin of Volcanology, 65(5), pp.363-381, doi:
10.1007/s00445-003-0277-y.
Jennings, C.W., 1977, Geologic Map of California: California Division of Mines and Geology Geologic
Data Map No. 2, scale 1:750,000
Jewell, P.W. and Stallard, R.F., 1991. Geochemistry and paleoceanographic setting of central Nevada
bedded barites. The Journal of Geology, 99(2), pp.151-170, doi: 10.1086/629482.
Ji, W., Wu, F., Liu, C. and Chung, S., 2009. Geochronology and petrogenesis of granitic rocks in
Gangdese batholith, southern Tibet. Science in China Series D: Earth Sciences, 52(9), pp.1240-1261,
doi: 10.1007/s11430-009-0131-y.
Jiang, D. and Bentley, C., 2012. A micromechanical approach for simulating multiscale fabrics in large‐
scale high‐strain zones: Theory and application. Journal of Geophysical Research: Solid Earth,
117(B12), doi: 10.1029/2012JB009327.
Jicha, B.R. and Jagoutz, O., 2015. Magma production rates for intraoceanic arcs. Elements, 11(2), pp.105-
111, doi: 10.2113/gselements.11.2.105
Johnson M C, Plank T., 1999. Dehydration and melting experiments constrain the fate of subducted
sediments, Geochemistry, Geophysics, Geosystems, 1, doi:10.1029/1999GC000014
Johnson, C.M., Czamanske, G.K. and Lipman, P.W., 1989. Geochemistry of intrusive rocks associated
with the Latir volcanic field, New Mexico, and contrasts between evolution of plutonic and volcanic
rocks. Contributions to Mineralogy and Petrology, 103(1), pp.90-109, doi: 10.1007/BF00371367
Kaiser, J.F., de Silva, S., Schmitt, A.K., Economos, R. and Sunagua, M., 2016. Million-year melt–
presence in monotonous intermediate magma for a volcanic–plutonic assemblage in the Central
Andes: Contrasting histories of crystal-rich and crystal-poor super-sized silicic magmas. Earth and
Planetary Science Letters, 457, pp.73-86, doi: 10.1016/j.epsl.2016.09.048
Karakas, O., Degruyter, W., Bachmann, O. and Dufek, J., 2017. Lifetime and size of shallow magma
bodies controlled by crustal-scale magmatism. Nature Geoscience, 10(6), pp.446-450, doi:
10.1038/ngeo2959.
Karakas, O., Wotzlaw, J.F., Guillong, M., Ulmer, P., Brack, P., Economos, R., Bergantz, G.W., Sinigoi,
S. and Bachmann, O., 2019. The pace of crustal-scale magma accretion and differentiation beneath
silicic caldera volcanoes. Geology, 47(8), pp.719-723, doi: 10.1130/G46020.1.
Karlstrom, L., Dufek, J. and Manga, M., 2009. Organization of volcanic plumbing through magmatic
lensing by magma chambers and volcanic loads. Journal of Geophysical Research: Solid Earth, 114,
doi: 10.1029/2009JB006339.
196
Karlstrom, L., Lee, C.T. and Manga, M., 2014. The role of magmatically driven lithospheric thickening
on arc front migration. Geochemistry, Geophysics, Geosystems, 15(6), pp.2655-2675, doi:
10.1002/2014GC005355.
Karlstrom, L., Paterson, S.R. and Jellinek, A.M., 2017. A reverse energy cascade for crustal magma
transport. Nature Geoscience, 10(8), pp.604-608, doi: 10.1038/ngeo2982.
Kawabata H, Nishiura D, Sakaguchi H, and Tatsumi Y., 2013. Self-organized domain microstructures in a
plate-like particle suspension subjected to rapid simple shear. Rheol. Acta, 52, pp. 1–21,
doi:10.1007/s00397-012-0657-3.
Kay, S.M., Godoy, E. and Kurtz, A., 2005. Episodic arc migration, crustal thickening, subduction erosion,
and magmatism in the south-central Andes. Geological Society of America Bulletin, 117(1-2), pp.67-
88, doi: 10.1130/B25431.1.
Kelemen, P.B., Whitehead, J.A., Aharonov, E. and Jordahl, K.A., 1995. Experiments on flow focusing in
soluble porous media, with applications to melt extraction from the mantle. Journal of Geophysical
Research: Solid Earth, 100(B1), pp.475-496, doi: 10.1029/94JB02544.
Kendall, C., Sklash, M. G., and Bullen, T. D., 1995. Isotope Tracers of Water and Solute Sources in
Catchments, In: Solute Modelling in Catchment Systems, John Wiley and Sons, New York, pp. 261-
303.
Kern, J.M., de Silva, S.L., Schmitt, A.K., Kaiser, J.F., Iriarte, A.R. and Economos, R., 2016.
Geochronological imaging of an episodically constructed subvolcanic batholith: U-Pb in zircon
chronochemistry of the Altiplano-Puna Volcanic Complex of the Central Andes. Geosphere, 12(4),
pp.1054-1077, doi: 10.1130/GES01258.1.
Kerrick, D.M., 1970, Contact metamorphism in some areas of the Sierra Nevada, California: Geological
Society of America Bulletin, v. 81(10), p.2913-2938, doi: 10.1130/0016-
7606(1970)81[2913:CMISAO]2.0.CO;2
Kirsch, M., Paterson, S.R., Wobbe, F., Ardila, A.M.M., Clausen, B.L. and Alasino, P.H., 2016. Temporal
histories of Cordilleran continental arcs: Testing models for magmatic episodicity. American
Mineralogist, 101(10), pp.2133-2154, doi: 10.2138/am-2016-5718.
Kiser, E., Palomeras, I., Levander, A., Zelt, C., Harder, S., Schmandt, B., Hansen, S., Creager, K. and
Ulberg, C., 2016. Magma reservoirs from the upper crust to the Moho inferred from high-resolution
Vp and Vs models beneath Mount St. Helens, Washington State, USA. Geology, 44(6), pp.411-414,
doi: 10.1130/G37591.1.
Kistler, R.W., 1966, Geologic map of the Mono Craters quadrangle, Mono and Tuolumne Counties,
California: No. 462, U.S. Geological Survey, scale 1: 62,500
Kistler, R.W. 1990. Two different lithosphere types in the Sierra Nevada, California, in Anderson, J.L.,
ed., The nature and origin of Cordilleran magmatism: Geological Society of America Memoir, 174,
pp.271-281.
Kistler, R.W. and Dodge, F.C.W., 1966. Potassium‐argon ages of coexisting minerals from pyroxene‐
bearing granitic rocks in the Sierra Nevada, California. Journal of geophysical research, 71(8),
pp.2157-2161, doi: 10.1029/JZ071i008p02157
197
Kistler, R. W., and Fleck, R. J., 1994, Field guide for a transect of the Central Sierra Nevada, California:
Geochronology and isotope geology: No. 94-267, United States Geological Survey Open-File Report,
50 p, doi: 10.3133/ofr94267.
Kistler, R.W. and Peterman, Z.E., 1973. Variations in Sr, Rb, K, Na, and initial Sr87/Sr86 in Mesozoic
granitic rocks and intruded wall rocks in central California. Geological Society of America Bulletin,
84(11), pp.3489-3512, doi: 10.1130/0016-7606(1973)84<3489:VISRKN>2.0.CO;2.
Kistler, R.W. and Peterman, Z.E., 1978. Reconstruction of crustal blocks of California on the basis of
initial strontium isotopic compositions of Mesozoic granite rocks, USGS Professional Paper, 1071.
Kistler, R.W. and Swanson, S.E., 1981, Petrology and geochronology of metamorphosed volcanic rocks
and a Middle Cretaceous volcanic neck in the east‐central Sierra Nevada, California: Journal of
Geophysical Research: Solid Earth, v. 86(B11), p.10489-10501, doi: 10.1029/JB086iB11p10489
Kistler, R.W., Chappell, B.W., Peck, D.L. and Bateman, P.C., 1986. Isotopic variation in the Tuolumne
intrusive suite, central Sierra Nevada, California. Contributions to Mineralogy and Petrology, 94(2),
pp.205-220, doi: 10.1007/BF00592937.
Klemetti, E.W., Lackey, J.S. and Starnes, J., 2014. Magmatic lulls in the Sierra Nevada captured in zircon
from rhyolite of the Mineral King pendant, California. Geosphere, 10(1), pp.66-79, doi:
10.1130/GES00920.1.
Kodaira, S., Sato, T., Takahashi, N., Miura, S., Tamura, Y., Tatsumi, Y. and Kaneda, Y., 2007. New
seismological constraints on growth of continental crust in the Izu-Bonin intra-oceanic arc. Geology,
35(11), pp.1031-1034, doi: 10.1130/G23901A.1.
Lackey, J. S.,Valley, J.W. & Saleeby, J. B., 2005. Supracrustal input to magmas in the deep crust of
Sierra Nevada batholith: evidence from high-d18O zircon. Earth and Planetary Science Letters 235,
315-330, doi: 10.1016/j.epsl.2005.04.003.
Lackey, J.S., Valley, J.W., Chen, J.H. and Stockli, D.F., 2008. Dynamic magma systems, crustal
recycling, and alteration in the central Sierra Nevada batholith: the oxygen isotope record. Journal of
Petrology, 49(7), pp.1397-1426, doi: 10.1093/petrology/egn030
Lackey, J.S., Cecil, M.R., Windham, C.J., Frazer, R.E., Bindeman, I.N., and Gehrels, G.E., 2012, The
Fine Gold Intrusive Suite: The roles of basement terranes and magma source development in the Early
Cretaceous Sierra Nevada batholith: Geosphere, v. 8(2), p.292-313, doi: 10.1130/GES00745.1
Lake, E.T. and Farmer, G.L., 2015. Oligo-Miocene mafic intrusions of the San Juan Volcanic Field,
southwestern Colorado, and their relationship to voluminous, caldera-forming magmas. Geochimica et
Cosmochimica Acta, 157, pp.86-108, doi: /10.1016/j.gca.2015.02.020.
Le Maitre, R.W., Streckeisen, A., Zanettin, B., Le Bas, M.J., Bonin, B., Bateman, P., Bellieni, G., Dudek,
A., Efremova, A., Keller, J. and Lameyre, J., Sabine, P.A., Schmid, R., Sørensen, H., and Wooley,
A.R., 2004. Igneous rocks. A classification and glossary of terms. Recommendations of the IUGS
Subcommission on the Systematics of Igneous Rocks. Cambridge University Press
Lee, C.T.A., and Lackey J.S., 2015. Global continental arc flare-ups and their relation to long-term
greenhouse conditions. Elements, 11(2), pp. 125-130, doi: 10.2113/gselements.11.2.125.
198
Lee, C.T.A. and Morton, D.M., 2015. High silica granites: Terminal porosity and crystal settling in
shallow magma chambers. Earth and Planetary Science Letters, 409, pp.23-31, doi:
10.1016/j.epsl.2014.10.040.
Lee, C.T.A., Thurner, S., Paterson, S. and Cao, W., 2015. The rise and fall of continental arcs: Interplays
between magmatism, uplift, weathering, and climate. Earth and Planetary Science Letters, 425,
pp.105-119, doi: 10.1016/j.epsl.2015.05.045.
Li, X.C., Zhou, M.F., Yang, Y.H., Zhao, X.F. and Gao, J.F., 2018. Disturbance of the Sm-Nd isotopic
system by metasomatic alteration: A case study of fluorapatite from the Sin Quyen Cu-LREE-Au
deposit, Vietnam. American Mineralogist, 103 (9), pp. 1487-1496, doi: 10.2138/am-2018-6501.
Lipman, P.W., 1984, The roots of ash flow calderas in western North America: windows into the tops of
granitic batholiths: Journal of Geophysical Research: Solid Earth, v. 89(B10), p.8801-8841, doi:
10.1029/JB089iB10p08801
Lipman, P.W., 2007. Incremental assembly and prolonged consolidation of Cordilleran magma chambers:
Evidence from the Southern Rocky Mountain volcanic field. Geosphere, 3, pp.42–70, doi:
10.1130/GES00061.1
Lipman, P.W. and Bachmann, O., 2015. Ignimbrites to batholiths: Integrating perspectives from
geological, geophysical, and geochronological data. Geosphere, 11(3), pp.705-743, doi:
10.1130/GES01091.1.
Loetterle, J., 2004. Sedimentary structures in a layered granodiorite: a window into physical conditions
present during the development of the Tuolumne Intrusive Suite, Sierra Nevada, California. MSc
thesis, University of Washington.
Loetterle J., and Bergantz G.W., 2003. Sedimentary structures in a layered granodiorite: an example of
magma multiphase dynamics from the Tuolumne Intrusive Suite, Sierra Nevada, California.
Geological Society of America Abstracts with Programs, 35, p. 555
Lofgren, G.E., 1971, Experimentally produced devitrification textures in natural rhyolitic glass.
Geological Society of America Bulletin, v. 82, p. 111-124, doi: 10.1130/0016-
7606(1971)82[111:EPDTIN]2.0.CO;2
Lofgren G.E ., 1974, An experimental study of plagioclase crystal morphology: Isothermal crystallization,
American Journal of Science, v. 274, p. 243 -273, doi: 10.2475/ajs.274.3.243.
Longo, A.A., Dilles, J.H., Grunder, A.L. and Duncan, R., 2010. Evolution of calc-alkaline volcanism and
associated hydrothermal gold deposits at Yanacocha, Peru. Economic Geology, 105(7), pp.1191-1241,
doi: 10.2113/econgeo.105.7.1191.
Lowe, T.K., 1995. Petrogenesis of the Minarets and Merced Peak Volcanic-plutonic Complexes, Sierra
Nevada, California. [PhD Thesis], Stanford University.
Ludwig, K.R., 2003, User's manual for Isoplot 3.00, a geochronological toolkit for Microsoft Excel.
Berkeley Geochronology Center Spec. Publication, v.4, pp.25-32.
Lyell, C.S. (1838) Elements of Geology. John Murray, London.
199
Macias, S.E., 1996, The Sonora Intrusive Suite: Constraints on the assembly of a Late Cretaceous,
concentrically-zoned granitic pluton of the Sierra Nevada batholith [M.S. thesis]: Seattle, University
of Washington, 66 p.
Mamani, M., Wörner, G. and Sempere, T., 2010. Geochemical variations in igneous rocks of the Central
Andean orocline (13 S to 18 S): Tracing crustal thickening and magma generation through time and
space. Bulletin, 122(1-2), pp.162-182, doi: 10.1130/B26538.1.
Marsh B.D., 1981. On the crystallinity, probability of occurrence, and rheology of lava and magma.
Contributions to Mineralogy and Petrology, 78, pp. 85-98, doi: 10.1007/BF00371146.
Marsh, B.D., 1982. On the mechanics of igneous diapirism, stoping, and zone melting. American Journal
of Science, 282(6), pp.808-855, doi: 10.2475/ajs.282.6.808.
Marsh, B.D., 1996, Solidification fronts and magmatic evolution: Mineralogical Magazine, v. 60, p. 5–40,
doi: 10.1180/minmag.1996.060.398.03.
Marsh, B., 2004. A magmatic mush column rosetta stone: the McMurdo Dry Valleys of Antarctica. Eos,
Transactions American Geophysical Union, 85(47), pp.497-502, doi:
Martinez-Ardila, A.M., Paterson, S.R., Memeti, V., Parada, M.A. and Molina, P.G., 2019. Mantle driven
Cretaceous flare-ups in Cordilleran arcs. Lithos, 326, pp.19-27, doi: 10.1016/j.lithos.2018.12.007.
Masi, U., O'Neil, J.R. and Kistler, R.W., 1981. Stable isotope systematics in Mesozoic granites of central
and northern California and southwestern Oregon. Contributions to Mineralogy and Petrology, 76(1),
pp.116-126, doi: 10.1007/BF00373691.
Matzel J. Mundil R. Paterson S. Renne P. Nomade S., 2005. Evaluating pluton growth models using high
resolution geochronology: Tuolumne intrusive suite, Sierra Nevada, CA: Geological Society of
America Abstracts with Programs, v. 37, no. 7, p. 131.
Matzel J. Miller J.S. Mundil R. Paterson S.R., 2006b. Zircon saturation and the growth of the Cathedral
Peak pluton, CA. Geochimica et Cosmochimica Acta, 70(18), p. 403, doi: 10.1016/j.gca.2006.06.813.
McBirney, A.R. and Noyes, R.M., 1979. Crystallization and layering of the Skaergaard intrusion. Journal
of Petrology, 20(3), pp.487-554, doi: 10.1093/petrology/20.3.487.
McColl, J, 2017. Geological Mapping in Yosemite National Park, Sierra Nevada, California: Geologic
History of the Mount Dana area and Magmatic Structure Kinematics in the Tuolumne Intrusive Suite
[BSc thesis]: Durham University, 74p.
McIntire, M.Z., Bergantz, G.W. and Schleicher, J.M., 2019. On the hydrodynamics of crystal clustering.
Philosophical Transactions of the Royal Society A, 377(2139), doi: 10.1098/rsta.2018.0015.
McNulty, B.A., 1995, Shear zone development during magmatic arc construction: The Bench Canyon
shear zone, central Sierra Nevada, California: Geological Society of America Bulletin , v. 107, p.
1094–1107, doi:10.1130/0016-7606(1995)107<1094:SZDDMA>2.3.CO;2.
Memeti, V., 2009, Growth of the Cretaceous Tuolumne batholith and synchronous regional tectonics,
Sierra Nevada, CA: A coupled system in a continental margin arc setting, [Ph.D. thesis], University of
Southern California, 300 p.
200
Memeti, V., Paterson, S., Matzel, J., Mundil, R. and Okaya, D., 2010a. Magmatic lobes as “snapshots” of
magma chamber growth and evolution in large, composite batholiths: An example from the Tuolumne
intrusion, Sierra Nevada, California. Geological Society of America Bulletin, 122(11-12), pp.1912-
1931, doi: 10.1130/B30004.1.
Memeti, V., Gehrels, G.E., Paterson, S.R., Thompson, J.M., Mueller, R.M. and Pignotta, G.S., 2010b,
Evaluating the Mojave–Snow Lake fault hypothesis and origins of central Sierran metasedimentary
pendant strata using detrital zircon provenance analyses: Lithosphere, v. 2(5), p. 341-360, doi:
10.1130/L58.1
Memeti, V., Paterson, S., Mundil, R., Paterson, S.R. and Putirka, K.D., 2014. Day 4: magmatic evolution
of the Tuolumne Intrusive Complex, in Memeti, V., Paterson, S.R., and Putirka, K.D., eds., Formation
of the Sierra Nevada Batholith: Magmatic and Tectonic Processes and Their Tempos: Geological
Society of America Field Guide, 34, pp.43-74, doi: 10.1130/2014.0034(04).
Miller, J.S., Matzel, J.E., Miller, C.F., Burgess, S.D. and Miller, R.B., 2007. Zircon growth and recycling
during the assembly of large, composite arc plutons. Journal of Volcanology and Geothermal
Research, 167(1-4), pp.282-299, doi: 10.1016/j.jvolgeores.2007.04.019.
Miller R.B. Paterson S.R., 2001. Construction of mid-crustal sheeted plutons: Examples from the North
Cascades, Washington: Geological Society of America Bulletin, 113(11), p.1423–1442, doi:
10.1130/0016-7606(2001)113<1423:COMCSP>2.0.CO;2.
Miller, R.B., Paterson, S.R., Matzel J.P., and Snoke, A.W., 2009. Plutonism at different crustal levels:
Insights from the~ 5–40 km (paleodepth) North Cascades crustal section, Washington, in Miller, R.B.,
and Snoke, A.W., eds. Crustal cross sections from the western North American Cordillera and
elsewhere: Implications for tectonic and petrologic processes. Geological Society of America Special
Paper, 456, p. 125-149, doi: 10.1130/2009.2456(05).
Moore, J.G., 1959. The quartz diorite boundary line in the western United States. The Journal of Geology,
67(2), pp.198-210, doi: 10.1086/626573.
Morton, D.M., Miller, F.K., Kistler, R.W., Premo, W.R., Lee, C.T.A., Langenheim, V.E., Wooden, J.L.,
Snee, L.W., Clausen, B.L. and Cossette, P., 2014. Framework and petrogenesis of the northern
Peninsular Ranges batholith, southern California. Peninsular Ranges Batholith, Baja California and
Southern California: Geological Society of America Memoir, 211, pp.61-143, doi:
10.1130/2014.1211(03).
Mottana, A., 1986. Blueschist-facies metamorphism of manganiferous cherts: A review of the alpine
occurrences. In Evans, B.E., and Brown, E.H., (eds), Blueschists and eclogites, Geological Society of
America, Memoir 164, doi: 10.1130/MEM164-p267.
Mukhopadhyay, B. and Manton, W.I., 1994. Upper-mantle fragments from beneath the Sierra Nevada
Batholith: partial fusion, fractional crystallization, and metasomatism in a subduction related ancient
lithosphere. Journal of Petrology, 35(5), pp.1417-1450, doi: 10.1093/petrology/35.5.1417.
Murase, T., McBirney, A.R. and Melson, W.G., 1985. Viscosity of the dome of Mount St. Helens. Journal
of volcanology and geothermal research, 24(1-2), pp.193-204, doi: 10.1016/0377-0273(85)90033-2.
Nadin, E.S., Saleeby, J. and Wong, M., 2016. Thermal evolution of the Sierra Nevada batholith,
California, and implications for strain localization. Geosphere, 12(2), pp.377-399, doi:
10.1130/GES01224.1
201
Nakamura, H., Oikawa, T., Geshi, N. and Matsumoto, A., 2014. Migration of a volcanic front inferred
from K–A r ages of late Miocene to Pliocene volcanic rocks in central Japan. Island Arc, 23(3),
pp.236-250, doi: 10.1111/iar.12073.
Namur, O., Abily, B., Boudreau, A.E., Blanchette, F., Bush, J.W., Ceuleneer, G., Charlier, B., Donaldson,
C.H., Duchesne, J.C., Higgins, M.D., Morata, D., Nielsen T.F.D., O’Driscoll, B., Pang, K. N.,
Peacock, T., Spandler, C.J., Toramaru, A., and Veksler, I.V., 2015. Igneous layering in basaltic
magma chambers. In Layered intrusions. pp. 75-152). Springer, Dordrecht.
Nandedkar, R.H., Ulmer, P. and Müntener, O., 2014. Fractional crystallization of primitive, hydrous arc
magmas: an experimental study at 0.7 GPa. Contributions to Mineralogy and Petrology, 167(6),
p.1015, doi: 10.1007/s00410-014-1015-5
O'Driscoll, B. and VanTongeren, J.A., 2017. Layered intrusions: from petrological paradigms to precious
metal repositories. Elements: An International Magazine of Mineralogy, Geochemistry, and Petrology,
13(6), pp.383-389, doi: 10.2138/gselements.13.6.383.
O'Driscoll, L.J., Humphreys, E.D. and Saucier, F., 2009. Subduction adjacent to deep continental roots:
Enhanced negative pressure in the mantle wedge, mountain building and continental motion. Earth and
Planetary Science Letters, 280(1-4), pp.61-70, doi: 10.1016/j.epsl.2009.01.020.
Oliver, H.W., 1977. Gravity and magnetic investigations of the Sierra Nevada batholith, California.
Geological Society of America Bulletin, 88(3), pp.445-461, doi: 10.1130/0016-
7606(1977)88<445:GAMIOT>2.0.CO;2.
Oppenheim, L.F., Memeti, V., Barnes, C.G., Chambers, M., Werts, K., and Esposito, R., (in review).
Feldspar recycling across magma mush bodies during the voluminous Half Dome and Cathedral Peak
stages of the Tuolumne Intrusive Complex, Yosemite National Park, California. Submitted to
Geosphere.
Otamendi, J.E., Ducea, M.N., Tibaldi, A.M., Bergantz, G.W., Jesús, D. and Vujovich, G.I., 2009,
Generation of tonalitic and dioritic magmas by coupled partial melting of gabbroic and
metasedimentary rocks within the deep crust of the Famatinian magmatic arc, Argentina: Journal of
Petrology, p. 841-873, doi: 10.1093/petrology/egp022
Paterson, S.R., 2009. Magmatic tubes, pipes, troughs, diapirs, and plumes: Late-stage convective
instabilities resulting in compositional diversity and permeable networks in crystal-rich magmas of the
Tuolumne batholith, Sierra Nevada, California. Geosphere, 5(6), pp.496-527, doi:
10.1130/GES00214.1.
Paterson, S.R., and Ardill, K.E., 2019. Implications of multiple magmatic fabrics reflecting strains caused
by local flow and by regional stresses of crystal mushes in the Tuolumne Intrusive Complex, Sierra
Nevada, CA. AGU Fall Meeting Abstracts, V51H-0150
Paterson, S. R. and Ducea, M. N., 2015, Arc Magmatic Tempos: Gathering the Evidence. Elements, 11,
pp.91-98, doi: 10.2113/gselements.11.2.91
Paterson, S.R., and Farris, D.W., 2008. Downward host rock transport and the formation of rim
monoclines during the emplacement of Cordilleran batholiths: Transactions of the Royal Society of
Edinburgh: Earth Sciences; Special Issue Plutons and Batholiths (The Wallace Pitcher Memorial
Volume), v. 97(4), p. 397-413, doi: https://doi.org/10.1017/S026359330000153X.
202
Paterson, S.R., and Memeti, V., 2014, Day 5: Mesozoic volcanic rocks of the central Sierra Nevada Arc,
in Memeti, V., Paterson, S.R., and Putirka, K.D., eds., Formation of the Sierra Nevada Batholith:
Magmatic and Tectonic Processes and Their Tempos: Geological Society of America Field Guide 34,
p.75-85, doi: 10.1130/2014.0034(05)
Paterson, S.R., Fowler Jr, T.K., Schmidt, K.L., Yoshinobu, A.S., Yuan, E.S. and Miller, R.B., 1998.
Interpreting magmatic fabric patterns in plutons. Lithos, 44(1-2), pp.53-82, doi: 10.1016/S0024-
4937(98)00022-X.
Paterson, S.R., Vernon, R.H. and Žák, J., 2005. Mechanical instabilities and physical accumulation of K-
feldspar megacrysts in granitic magma, Tuolumne Batholith, California, USA. Journal of the Virtual
Explorer, 18(1), doi: 10.3809/jvirtex.2005.00114.
Paterson, S.R., Žák, J. and Janoušek, V., 2008. Growth of complex sheeted zones during recycling of
older magmatic units into younger: Sawmill Canyon area, Tuolumne batholith, Sierra Nevada,
California. Journal of Volcanology and Geothermal Research, 177(2), pp.457-484, doi:
10.1016/j.jvolgeores.2008.06.024.
Paterson, S.R., Okaya, D., Memeti, V., Economos, R. and Miller, R.B., 2011. Magma addition and flux
calculations of incrementally constructed magma chambers in continental margin arcs: Combined
field, geochronologic, and thermal modeling studies. Geosphere, 7(6), pp.1439-1468, doi:
10.1130/GES00696.1.
Paterson, S.R., Memeti, V., Pignotta, G., Erdmann, S., Žák, J., Chambers, J. and Ianno, A., 2012.
Formation and transfer of stoped blocks into magma chambers: The high-temperature interplay
between focused porous flow, cracking, channel flow, host-rock anisotropy, and regional deformation.
Geosphere, 8(2), pp.443-469, doi: 10.1130/GES00680.1.
Paterson, S.R., Memeti, V., Anderson, L., Cao, W., Lackey, J.S., Putirka, K.D., Miller, R.B., Miller, J.S.
and Mundil, R., 2014. Day 6: Overview of arc processes and tempos, in, Memeti, V., Paterson, S.R.,
and Putirka, K., eds, Formation of the Sierra Nevada Batholith: Magmatic and Tectonic Processes and
their Tempos: Geological Society of America Field Guide, 34, pp.87-116, doi:
10.1130/2014.0034(06).
Paterson, S., Memeti, V., Mundil, R. and Žák, J., 2016. Repeated, multiscale, magmatic erosion and
recycling in an upper-crustal pluton: Implications for magma chamber dynamics and magma volume
estimates. American Mineralogist, 101(10), pp.2176-2198, doi: 10.2138/am-2016-5576.
Paterson, S.R., Ardill, K., Vernon, R. and Žák, J., 2018. A review of mesoscopic magmatic structures and
their potential for evaluating the hypersolidus evolution of intrusive complexes. Journal of Structural
Geology, 125, pp. 134-147, doi: /10.1016/j.jsg.2018.04.022.
Paterson S.R., Ardill K, Memeti V, Werts K & Barnes C, 2019, Forming super-plutons: priming arcs and
producing active, long-lived magma mushes. Goldschmidt Abstracts, 2019 #2590
Peck, D.L., 1980, Geologic map of the Merced Peak quadrangle, central Sierra Nevada, California: No.
1531, U.S. Geological Survey, scale 1: 62,500
Peck, D.L., 2002, Geologic map of the Yosemite quadrangle, central Sierra Nevada, California: No. 2751,
U.S. Geological Survey, scale 1: 62,500
203
Peck, D.L., and Van Kooten, G.K., 1983, Merced Peak Quadrangle, Central Sierra Nevada, California:
Analytic Data, No. 1170, US Geological Survey
Petford, N., 2003. Rheology of granitic magmas during ascent and emplacement. Annual Review of Earth
and Planetary Sciences, 31(1), pp.399-427, doi: 10.1146/annurev.earth.31.100901.141352.
Petford, N., 2009. Which effective viscosity? Mineralogical Magazine, 73(2), pp.167-191, doi:
10.1180/minmag.2009.073.2.167.
Piccoli, P. and Candela, P., 1994. Apatite in felsic rocks; a model for the estimation of initial halogen
concentrations in the Bishop Tuff (Long Valley) and Tuolumne Intrusive Suite (Sierra Nevada
Batholith) magmas. American Journal of Science, 294(1), pp.92-135, doi: 10.2475/ajs.294.1.92.
Pickett, D.A. and Saleeby, J.B., 1993. Thermobarometric constraints on the depth of exposure and
conditions of plutonism and metamorphism at deep levels of the Sierra Nevada batholith, Tehachapi
Mountains, California. Journal of Geophysical Research: Solid Earth, 98(B1), pp.609-629, doi:
10.1029/92JB01889.
Pinotti, L.P., D'Eramo, F.J., Weinberg, R.F., Demartis, M., Tubía, J.M., Coniglio, J.E., Radice, S.,
Maffini, M.N. and Aragón, E., 2016. Contrasting magmatic structures between small plutons and
batholiths emplaced at shallow crustal level (Sierras de Córdoba, Argentina). Journal of Structural
Geology, 92, pp.46-58, doi: 10.1016/j.jsg.2016.09.009.
Pitcher, W.S., 1997. The nature and origin of granite. Springer Science & Business Media.
Plank T., 2005. Constraints from thorium/lanthanum on sediment recycling at subduction zones and the
evolution of the continents, Journal of Petrology, v.46, p. 921-944.
Plank, T., Langmuir, C., 1993. Tracing trace elements from sediment input to volcanic output at
subduction zones. Nature 362, p. 739–743, doi: 10.1038/362739a0
Pons, J., Barbey, P., Nachit, H. and Burg, J.P., 2006. Development of igneous layering during growth of
pluton: The Tarcouate Laccolith (Morocco). Tectonophysics, 413(3-4), pp.271-286, doi:
10.1016/j.tecto.2005.11.005.
Profeta, L., Ducea, M.N., Chapman, J.B., Paterson, S.R., Gonzales, S.M.H., Kirsch, M., Petrescu, L. and
DeCelles, P.G., 2015. Quantifying crustal thickness over time in magmatic arcs. Scientific reports, 5,
p. 17786, doi: 10.1038/srep17786.
Pu, X., Delph, J.R., Shimizu, K., Rasmussen, D.J. and Ratschbacher, B.C., 2017. Where do arc magmas
differentiate? A seismic and geochemical search for active, deep crustal MASH zones. AGUFM, 2017,
pp.V11C-0359.
Putirka, K., 2016a. Amphibole thermometers and barometers for igneous systems and some implications
for eruption mechanisms of felsic magmas at arc volcanoes. American Mineralogist, 101, 841–858,
doi: 10.2138/am-2016-5506.
Putirka, K.D., Canchola, J., Rash, J., Smith, O., Torrez, G., Paterson, S.R. and Ducea, M.N., 2014. Pluton
assembly and the genesis of granitic magmas: Insights from the GIC pluton in cross section, Sierra
Nevada Batholith, California. American Mineralogist, 99(7), pp.1284-1303, doi:
10.2138/am.2014.4564.
204
Rapela, C.W., Pankhurst, R.J., Casquet, C., Dahlquist, J.A., Fanning, C.M., Baldo, E.G., Galindo, C.,
Alasino, P.H., Ramacciotti, C.D., Verdecchia, S.O. and Murra, J.A., 2018. A review of the Famatinian
Ordovician magmatism in southern South America: evidence of lithosphere reworking and continental
subduction in the early proto-Andean margin of Gondwana. Earth-science reviews, 187, pp.259-285,
doi: 10.1016/j.earscirev.2018.10.006.
Ratajeski, K., Glazner, A.F. and Miller, B.V., 2001. Geology and geochemistry of mafic to felsic plutonic
rocks in the Cretaceous intrusive suite of Yosemite Valley, Geological Society of America Bulletin,
113(11), pp.1486-1502, doi: 10.1130/0016-7606(2001)113<1486:GAGOMT>2.0.CO;2
Ratschbacher, B.C., Keller, C.B., Schoene, B., Paterson, S.R., Anderson, J.L., Okaya, D., Putirka, K. and
Lippoldt, R., 2018. A new workflow to assess emplacement duration and melt residence time of
compositionally diverse magmas emplaced in a sub-volcanic reservoir. Journal of Petrology, 59(9),
pp.1787-1809, doi: 10.1093/petrology/egy079.
Ratschbacher, B.C., Paterson, S.R. and Fischer, T.P., 2019. Spatial and depth‐dependent variations in
magma volume addition and addition rates to continental arcs: Application to global CO2 fluxes since
750 Ma. Geochemistry, Geophysics, Geosystems, 20(6), pp.2997-3018, doi: 10.1029/2018GC008031.
Reid Jr, J.B., Murray, D.P., Hermes, O.D. and Steig, E.J., 1993. Fractional crystallization in granites of
the Sierra Nevada: How important is it? Geology, 21(7), pp.587-590, doi: 10.1130/0091-
7613(1993)021<0587:FCIGOT>2.3.CO;2.
Reiners, P.W., Carlson, R.W., Renne, P.R., Cooper, K.M., Granger, D.E., McLean, N.M., Schoene, B.,
2018, Geochronology and Thermochronology, John Wiley & Sons Ltd., p. 493
Ribeiro, J.M., Ishizuka, O., Lee, C.T.A. and Girard, G., 2020. Evolution and maturation of the nascent
Mariana arc. Earth and Planetary Science Letters, 530, p.115912, doi: 10.1016/j.epsl.2019.115912.
Riciputi, L.R., Johnson, C.M., Sawyer, D.A. and Lipman, P.W., 1995. Crustal and magmatic evolution in
a large multicyclic caldera complex: isotopic evidence from the central San Juan volcanic field.
Journal of Volcanology and Geothermal Research, 67(1), pp.1-28, doi: 10.1016/0377-0273(94)00097-
Z.
Robinson, A.C., and Kistler, R.W., 1986, Maps showing isotopic dating in the Walker Lake 1° by 2°
quadrangle, California and Nevada: U.S. Geological Survey Miscellaneous Field Studies Map MF–
1382–N, scale 1:250,000.
Rocher, S., Alasino, P.H., Grande, M.M., Larrovere, M.A. and Paterson, S.R., 2018. K-feldspar
megacryst accumulations formed by mechanical instabilities in magma chamber margins, Asha pluton,
NW Argentina. Journal of Structural Geology, 112, pp.154-173, doi: 10.1016/j.jsg.2018.04.017.
Rosenberg, C.L. and Handy, M.R., 2005. Experimental deformation of partially melted granite revisited:
implications for the continental crust. Journal of Metamorphic Geology, 23(1), pp.19-28, doi:
10.1111/j.1525-1314.2005.00555.
Russell, S.J., 1976, Stratigraphy and structure of Mesozoic metavolcanic rocks in the vicinity of Mt.
Dana, Yosemite National Park, California [Master's thesis]: California State University, Fresno, 71 p.
205
Ryan-Davis, J., Lackey, J.S., Gevedon, M., Barnes, J.D., Lee, C.A., Kitajima, K. and Valley, J.W., 2019.
Andradite skarn garnet records of exceptionally low δ 18 O values within an Early Cretaceous
hydrothermal system, Sierra Nevada, CA. Contributions to Mineralogy and Petrology, 174(8), p.68,
doi: 10.1007/s00410-019-1602-6.
Ryerson, F.J., and Hess, P.C., 1978, Implications of liquid-liquid distribution coefficients to mineral-
liquid partitioning: Geochimica et Cosmochimica Acta, v. 42, p. 921-932, doi: 10.1016/0016-
7037(78)90103-5.
Saleeby J.B., 1990. Progress in Tectonic and Petrogenetic Studies in an Exposed Cross-Section of Young
(~100 Ma) Continental Crust, Southern Sierra Nevada, California. In: Salisbury M.H., Fountain D.M.
(eds) Exposed Cross-Sections of the Continental Crust. NATO ASI Series (Series C: Mathematical
and Physical Sciences), 317, Springer, Dordrecht, doi: 10.1007/978-94-009-0675-4_6.
Saleeby, J., 2007. The western extent of the Sierra Nevada batholith in the Great Valley basement and its
significance in underlying mantle dynamics. AGU Fall Meeting Abstracts.
Saleeby, J., 2011, Geochemical mapping of the Kings-Kaweah ophiolite belt, southwestern Sierra Nevada
Foothills—Evidence for progressive mélange formation in a large offset transform-subduction
initiation environment, in Wakabayashi, J., and Dilek, E., eds., Mélanges: Processes of Formation and
Societal Significance: Geological Society of America Special Paper 480, p. 31–73,
doi:10.1130/2011.2480(02).
Saleeby, J.B., Speed R., Blake M., Allmendinger R., Gans P.B., Kistler R.W., 1986. Continent-Ocean
Transect, Corridor C2, Monterey Bay Offshore to the Colorado Plateau: Geological Society of
America Map and Chart Series TRA C2, 2 sheets, scale 1:500,000, 87 p.
Saleeby, J.B., Shaw, H.F., Niemeyer, S., Edelman, S.H., and Moores, E.M., 1989. U/Pb, Sm/Nd and
Rb/Sr geochronological and isotopic study of northern Sierra Nevada ophiolitic assemblages,
California: Contributions to Mineralogy and Petrology, 102, 205–220, doi: 10.1007/BF00375341.
Saleeby, J. B., Kistler, R.W., Longiaru, S., Moore, J.G. and Nokleberg, W.J., 1990, Middle Cretaceous
silicic metavolcanic rocks in the Kings Canyon area, central Sierra Nevada, California: Geological
Society of America Memoirs, v. 174, p. 251-271
Saleeby, J., Ducea, M. and Clemens‐Knott, D., 2003. Production and loss of high‐density batholithic root,
southern Sierra Nevada, California. Tectonics, 22(6), doi: 10.1029/2002TC001374.
Sampson, D.E., 1987, Textural heterogeneities and vent area structures in the 600-year-old lavas of the
Inyo volcanic chain, eastern California: Geological Society of America Special Papers, v. 212, p.89-
102, doi: 10.1130/SPE212-p89
Schaen, A.J., Singer, B.S., Cottle, J.M., Garibaldi, N., Schoene, B., Satkoski, A.M. and Fournelle, J.,
2018. Textural and mineralogical record of low-pressure melt extraction and silicic cumulate
formation in the late Miocene Risco Bayo–Huemul plutonic complex, southern Andes. Journal of
Petrology, 59(10), pp.1991-2016, doi: 10.1093/petrology/egy087.
Scheland, C., 2019. The physical and petrologic evolution of the Jack Main canyon intrusive suite- a
migrating intrusion in the central Sierra Nevada, California [MSc Thesis], California State University
Fullerton
206
Schleicher, J.M., Bergantz, G.W., Breidenthal, R.E. and Burgisser, A., 2016. Time scales of crystal
mixing in magma mushes. Geophysical Research Letters, 43(4), pp.1543-1550, doi:
10.1002/2015GL067372.
Schwartz, J.J., Klepeis, K.A., Sadorski, J.F., Stowell, H.H., Tulloch, A.J. and Coble, M.A., 2017. The
tempo of continental arc construction in the Mesozoic Median Batholith, Fiordland, New Zealand.
Lithosphere, 9(3), pp.343-365, doi: 10.1130/L610.1.
Schweickert, R.A. and Lahren, M.M., 1990. Speculative reconstruction of the Mojave‐Snow Lake fault:
Implications for Paleozoic and Mesozoic orogenesis in the western United States. Tectonics, 9(6),
pp.1609-1629, doi: 10.1029/TC009i006p01609.
Schweickert, R.A., and Lahren, M.M., 1999, Triassic caldera at Tioga Pass, Yosemite National park,
California: structural relationships and significance: Geological Society of America Bulletin, 111(11),
p.1714-1722, doi: 10.1130/0016-7606(1999)111<1714:TCATPY>2.3.CO;2
Schweickert, R.A., and Lahren, M.M., 2006, Geologic evolution of Saddlebag Lake pendant, eastern
Sierra Nevada, California: Tectonic implications, in Girty, G.H., and Cooper, J.D., eds., Using
stratigraphy, sedimentology, and geochemistry to unravel the geologic history of the southwestern
Cordillera: Pacific Section, SEPM (Society for Sedimentary Geology) Publication 101, p. 27–56.
Seaman, S.J., 2000, Crystal clusters, feldspar glomerocrysts, and magma envelopes in the Atascosa
Lookout lava flow, southern Arizona, USA: recorders of magmatic events: Journal of Petrology, v.
41(5), p.693-716, doi: 10.1093/petrology/41.5.693
Sharp, W.D., Tobisch, O.T., Renne, P.R., 2000, Development of Cretaceous transpressional cleavage
synchronous with batholith emplacement, central Sierra Nevada, California: Geological Society of
America Bulletin , v. 112, p. 1059–1066, doi:10.1130/0016-
7606(2000)112<1059:DOCTCS>2.0.CO;2.
Sisson, T.W. and Moore, J.G., 2013. Geologic map of southwestern Sequoia National Park, Tulare
County, California (No. 2013-1096). US Geological Survey.
Skinner, S.M. and Clayton, R.W., 2011. An evaluation of proposed mechanisms of slab flattening in
central Mexico. Pure and Applied geophysics, 168(8-9), pp.1461-1474, doi: 10.1007/s00024-010-
0200-3.
Smith, R.L., 1979. Ash-flow magmatism. Ash-flow tuffs: Geological Society of America Special Paper,
180, pp.5-27, doi: 10.1130/SPE180.
Smith, T.E., 1974. The geochemistry of the granitic rocks of Halifax County, Nova Scotia. Canadian
Journal of Earth Sciences, 11(5), pp.650-657, doi: 10.1139/e74-062.
Snow, C.A., 2007. Petrotectonic evolution and melt modeling of the Peñon Blanco arc, central Sierra
Nevada foothills, California. Geological Society of America Bulletin, 119(7-8), pp.1014-1024, doi:
10.1130/B25972.1.
Solgadi, F. and Sawyer, E.W., 2008. Formation of igneous layering in granodiorite by gravity flow: A
field, microstructure and geochemical study of the Tuolumne Intrusive Suite at Sawmill Canyon,
California. Journal of Petrology, 49(11), pp.2009-2042, doi: 10.1093/petrology/egn056.
207
Sorensen, S.S., Dunne, G.C., Hanson, R.B., Barton, M.D., Becker, J., Tobisch, O.T. and Fiske, R.S.,
1998, From Jurassic shores to Cretaceous plutons: Geochemical evidence for paleoalteration
environments of metavolcanic rocks, eastern California: Geological Society of America Bulletin, v.
110(3), p.326-343, doi: 10.1130/0016-7606(1998)110<0326:FJSTCP>2.3.CO;2
Sparks, R.S.J., Annen, C., Blundy, J.D., Cashman, K.V., Rust, A.C. and Jackson, M.D., 2019. Formation
and dynamics of magma reservoirs. Philosophical Transactions of the Royal society A, 377(2139),
doi: 10.1098/rsta.2018.0019.
Spera, F.J. and Bohrson, W.A., 2018. Rejuvenation of crustal magma mush: a tale of multiply nested
processes and timescales. American Journal of Science, 318(1), pp.90-140, doi: 10.2475/01.2018.05.
Stanback, J., 2018. Structural, emplacement and tectonic history of spatially and temporally linked
volcanic, hypabyssal, and plutonic units in the Beartrap Lake area, Eastern Sierra Nevada, CA [MSc
Thesis], University of Southern California, 101 p,
Stern, R.J. and Scholl, D.W., 2010. Yin and yang of continental crust creation and destruction by plate
tectonic processes. International Geology Review, 52(1), pp.1-31, doi: 10.1080/00206810903332322.
Stern, R.J., Jackson, M.C., Fryer, P. and Ito, E., 1993. O, Sr, Nd and Pb isotopic composition of the
Kasuga cross-chain in the Mariana Arc: A new perspective on the Kh relationship. Earth and Planetary
Science Letters, 119(4), pp.459-475, doi: 10.1016/0012-821X(93)90056-F.
Steven, T.A., Luedke, R.G. and Lipman, P.W., 1974. Relation of mineralization to calderas in the San
Juan volcanic field, southwestern Colorado. J. Res. US geol. Surv., 2, pp.405-409.
Stevens, C. H., Stone, P., Dunne, G. C., Greene, D. C., Walker, J. D., and Swanson, B. J., 1997, Paleozoic
and Mesozoic evolution of east-central California: International Geology Review, v. 39, p. 788–829,
doi: 10.1080/00206819709465303.
Storck, J.C., Wotzlaw, J.F., Karakas, Ö., Brack, P., Gerdes, A. and Ulmer, P., 2020. Hafnium isotopic
record of mantle-crust interaction in an evolving continental magmatic system. Earth and planetary
science letters, 535, p.116100, doi: 10.1016/j.epsl.2020.116100.
Sun, S.S. and McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts:
implications for mantle composition and processes. Geological Society, London, Special
Publications, 42(1), pp.313-345, doi: 10.1144/GSL.SP.1989.042.01.19.
Swanson, S.E., Naney, M.T., Westrich, H.R. and Eichelberger, J.C., 1989, Crystallization history of
Obsidian dome, Inyo domes, California. Bulletin of Volcanology, v. 51(3), pp.161-176, doi:
10.1007/BF01067953.
Tierney, C.R., Schmitt, A.K., Lovera, O.M. and de Silva, S.L., 2016. Voluminous plutonism during
volcanic quiescence revealed by thermochemical modeling of zircon. Geology, 44(8), pp.683-686, doi:
10.1130/G37968.1.
Tikoff, B., Davis, M.R., Teyssier, C., Blanquat, M.D.S., Habert, G. and Morgan, S., 2005. Fabric studies
within the Cascade Lake shear zone, Sierra Nevada, California. Tectonophysics, 400(1-4), pp.209-
226, doi: 10.1016/j.tecto.2005.03.003
208
Till, C.B., Kent, A.J.R., Abers, G.A., Janiszewski, H.A., Gaherty, J.B. and Pitcher, B.W., 2019. The
causes of spatiotemporal variations in erupted fluxes and compositions along a volcanic arc. Nature
communications, 10(1), pp.1-12, doi: 10.1038/s41467-019-09113-0.
Titus, S.J., Clark, R. and Tikoff, B., 2005. Geologic and geophysical investigation of two fine-grained
granites, Sierra Nevada Batholith, California: Evidence for structural controls on emplacement and
volcanism. Geological Society of America Bulletin, 117(9-10), pp.1256-1271., doi:
10.1130/B25689.1.
Tobisch, O.T., Barton, M.D., Vernon, R.H. and Paterson, S.R., 1991. Fluid-enhanced deformation:
transformation of granitoids to banded mylonites, western Sierra Nevada, California, and southeastern
Australia. Journal of Structural Geology, 13(10), pp.1137-1156, doi:
Tobisch, O.T., Fiske, R.S., Saleeby, J.B., Holt, E., and Sorensen, S.S., 2000, Steep tilting of metavolcanic
rocks by multiple mechanisms, central Sierra Nevada, California: Geological Society of America
Bulletin, v. 112(7), p.1043-1058, doi: 10.1130/0016-7606(2000)112<1043:STOMRB>2.0.CO;2
Tomek, F., Žák, J., Verner, K., Holub, F.V., Sláma, J., Paterson, S.R. and Memeti, V., 2017. Mineral
fabrics in high-level intrusions recording crustal strain and volcano–tectonic interactions: the
Shellenbarger pluton, Sierra Nevada, California. Journal of the Geological Society, 174(2), pp.193-
208, doi: 10.1144/jgs2015-151.
Truschel, J. P., 1996. Petrogenesis of the Fine Gold intrusive suite, Sierra Nevada Batholith, California.
[M.S. thesis], California State University, Northridge, California.
Valley, J.W., Kinny, P.D., Schulze, D.J. and Spicuzza, M.J., 1998. Zircon megacrysts from kimberlite:
oxygen isotope variability among mantle melts. Contributions to mineralogy and petrology, 133(1-2),
pp.1-11, doi: 10.1007/s004100050432.
Van der Molen, I. and Paterson, M.S., 1979. Experimental deformation of partially-melted granite.
Contributions to Mineralogy and Petrology, 70(3), pp.299-318, doi: 10.1007/BF00375359.
Vance, J.A., 1969, On synneusis: Contributions to Mineralogy and Petrology, v. 24(1), p.7-29, doi:
10.1007/BF00398750
Vance, J.A., and Gilreath, J.P., 1967., Effect of synneusis on phenocryst distribution patterns in some
porphyritic igneous rocks: American Mineralogist, v. 52(3-4), p.529
Vernon, R.H. and Paterson, S.R., 2006. Mesoscopic structures resulting from crystal accumulation and
melt movement in granites. Earth and Environmental Science Transactions of the Royal Society of
Edinburgh, 97(4), pp.369-381, doi: 10.1017/S0263593300001516.
Vukmanovic, Z., Holness, M.B., Monks, K. and Andersen, J.Ø., 2018. The Skaergaard trough layering:
sedimentation in a convecting magma chamber. Contributions to Mineralogy and Petrology, 173(5),
p.43, doi: 10.1007/s00410-018-1466-1.
Wager, L.R. and Brown, G.M., 1968. Layered igneous rocks. Oliver and Boyd. Edinburgh and London,
pp. 588, doi: 10.1180/minmag.1968.036.284.25.
Wager L.R., and Deer W.A., 1939. Geological investigations in East Greenland. Part III. The petrology of
the Skaergaard intrusion, Kangerdlussuaq, East Greenland. Meddelelser om Grønland. 105(4):352.
209
Wagner, D.L., Bortugno, E.J. and McJunkin, R.D., 1991. Geologic map of the San Francisco-San Jose
quadrangle. California Department of Conservation, Division of Mines and Geology.
Wahrhaftig, C., 1979, Significance of asymmetric schlieren for crystallization of granites in the Sierra
Nevada batholith, California: Geological Society of America Abstracts with Programs , v. 11, p. 133.
Walker, B.A., Grunder, A.L. and Wooden, J.L., 2010. Organization and thermal maturation of long-lived
arc systems: Evidence from zircons at the Aucanquilcha volcanic cluster, northern Chile. Geology,
38(11), pp.1007-1010, doi: 10.1130/G31226.1
Walker, B.A., Klemetti, E.W., Grunder, A.L., Dilles, J.H., Tepley, F.J. and Giles, D., 2013. Crystal
reaming during the assembly, maturation, and waning of an eleven-million-year crustal magma cycle:
thermobarometry of the Aucanquilcha Volcanic Cluster. Contributions to Mineralogy and Petrology,
165(4), pp.663-682, 10.1007/s00410-012-0829-2.
Walker, B.A., Bergantz, G.W., Otamendi, J.E., Ducea, M.N., Cristofolini, E.A., 2015. A MASH Zone
Revealed: The Mafic Complex of the Sierra Valle Fértil, Journal of Petrology, 56 (9), pp.1863-1896,
doi: 10.1093/petrology/egv057.
Wallace, P.J., Plank, T., Edmonds, M. and Hauri, E.H., 2015. Volatiles in magmas. In Sigurdsson, H.,
eds., The Encyclopedia of Volcanoes, Academic Press, p. 163-183.
Ward, K.M., Delph, J.R., Zandt, G., Beck, S.L. and Ducea, M.N., 2017. Magmatic evolution of a
Cordilleran flare-up and its role in the creation of silicic crust. Scientific Reports, 7(1), pp.9047.
Watts, K.E., John, D.A., Colgan, J.P., Henry, C.D., Bindeman, I.N. and Schmitt, A.K., 2016. Probing the
Volcanic–Plutonic Connection and the Genesis of Crystal-rich Rhyolite in a Deeply Dissected
Supervolcano in the Nevada Great Basin: Source of the Late Eocene Caetano Tuff: Journal of
Petrology, v. 57(8), p.1599-1644, doi: 10.1093/petrology/egw051
Weinberg, R.F., Sial, A.N. and Pessoa, R.R., 2001. Magma flow within the Tavares pluton, northeastern
Brazil: Compositional and thermal convection. Geological Society of America Bulletin, 113(4),
pp.508-520, doi: 10.1130/0016-7606(2001)113<0508:MFWTTP>2.0.CO;2.
Wendt, I. and Carl, C., 1991, The statistical distribution of the mean squared weighted deviation:
Chemical Geology, Isotope Geoscience Section, v. 86(4), p. 275–285, doi: 10.1016/0168-
9622(91)90010-T.
Werts, K., Barnes, C.G., Memeti, V., Ratschbacher, B., Williams, D. and Paterson, S.R., 2020.
Hornblende as a tool for assessing mineral-melt equilibrium and recognition of crystal accumulation.
American Mineralogist: Journal of Earth and Planetary Materials, 105(1), pp.77-91, doi: 10.2138/am-
2020-6972.
Whitehead, J.A., Helfrich, K.R., 1991. Instability of flow with temperature-dependent viscosity: a model
of magma dynamics. Journal of Geophysical Research 96, pp.4145-4155, doi: 10.1029/90JB02342
Wiebe, R.A. and Collins, W.J., 1998. Depositional features and stratigraphic sections in granitic plutons:
implications for the emplacement and crystallization of granitic magma. Journal of Structural
Geology, 20(9-10), pp.1273-1289, doi: 10.1016/S0191-8141(98)00059-5.
210
Wiebe, R.A., Jellinek, M., Markley, M.J., Hawkins, D.P. and Snyder, D., 2007. Steep schlieren and
associated enclaves in the Vinalhaven granite, Maine: possible indicators for granite rheology.
Contributions to Mineralogy and Petrology, 153(2), p.121, doi: 10.1007/s00410-006-0142-z.
Wiebe, R.A., Jellinek, A.M. and Hodge, K.F., 2017. New insights into the origin of ladder dikes:
Implications for punctuated growth and crystal accumulation in the Cathedral Peak granodiorite.
Lithos, 277, pp.241-258, doi: 10.1016/j.lithos.2016.09.015.
Williams, D., 2018. From sheets to blobs: The initial emplacement and evolution of a large, long-lived
intrusion, the Tuolumne intrusive complex, Sierra Nevada, California. [MSc thesis], California State
University Fullerton.
Wooden, J.L., Kistler, R.W., and Tosdal, R.M., 1999, Strontium, lead, and oxygen isotopic data for
granitoids and volcanic rocks from the northern Great Basin and Sierra Nevada, California, Nevada,
and Utah: U.S. Geological Survey, Open-File Report 99-569, 20 p, doi: 10.3133/ofr99569.
Yan, C.Y., and Kroenke, L.W., 1993. A plate tectonic reconstruction of the southwest Pacific, 0-100 Ma,
in Berger, W.H., Kroenke, L.W., Mayer L.A., et al., eds., Proceedings of the Ocean Drilling Program,
Scientific Results, v. 30
Žák, J. and Klomínský, J., 2007. Magmatic structures in the Krkonoše–Jizera Plutonic Complex,
Bohemian Massif: evidence for localized multiphase flow and small-scale thermal–mechanical
instabilities in a granitic magma chamber. Journal of Volcanology and Geothermal Research, 164(4),
pp.254-267, doi: 10.1016/j.jvolgeores.2007.05.006.
Žák, J. and Paterson, S.R., 2005. Characteristics of internal contacts in the Tuolumne Batholith, central
Sierra Nevada, California (USA): Implications for episodic emplacement and physical processes in a
continental arc magma chamber. Geological Society of America Bulletin, 117(9-10), pp.1242-1255,
doi: 10.1130/B25558.1.
Žák, J., and Paterson, S.R., 2010, Magmatic erosion of the solidification front during reintrusion: The
eastern margin of the Tuolumne batholith, Sierra Nevada, California: International Journal of Earth
Sciences , v. 99, p. 801–812, doi: 10.1007/s00531-009-0423-7.
Žák, J., Paterson, S.R. and Memeti, V., 2007. Four magmatic fabrics in the Tuolumne batholith, central
Sierra Nevada, California (USA): Implications for interpreting fabric patterns in plutons and evolution
of magma chambers in the upper crust. Geological Society of America Bulletin, 119(1-2), pp.184-201,
doi: 10.1130/B25773.1.
Žák, J., Paterson, S.R., Janoušek, V. and Kabele, P., 2009. The Mammoth Peak sheeted complex,
Tuolumne batholith, Sierra Nevada, California: a record of initial growth or late thermal contraction in
a magma chamber? Contributions to Mineralogy and Petrology, 158(4), p.447, doi: 10.1007/s00410-
009-0391-8.
Zartman, R.E., 1974. Lead isotopic provinces in the Cordillera of the western United States and their
geologic significance. Economic Geology, 69(6), pp.792-805, doi: 10.2113/gsecongeo.69.6.792
Zellmer, G.F., Edmonds, M. and Straub, S.M., 2015. Volatiles in subduction zone magmatism.
Geological Society, London, Special Publications, 410(1), pp.1-17, doi: 10.1144/SP410.13.
211
Zen, E.A., 1986. Aluminum enrichment in silicate melts by fractional crystallization: some mineralogic
and petrographic constraints. Journal of Petrology, 27(5), pp.1095-1117, doi:
10.1093/petrology/27.5.1095.
Zen, E.A., 1988, Phase relations of peraluminous granitic rocks and their petrogenetic implications:
Annual Review of Earth and Planetary Sciences, v.16(1), p.21-51, doi:
10.1146/annurev.ea.16.050188.000321
Zhang, J., Humphreys, M.C., Cooper, G.F., Davidson, J.P., and Macpherson, C.G., 2017. Magma mush
chemistry at subduction zones, revealed by new melt major element inversion from calcic amphiboles.
American Mineralogist, 102, 1353–1367, doi: 10.2138/am-2017-5928.
Zimmerer, M.J. and McIntosh, W.C., 2013. Geochronologic evidence of upper-crustal in-situ
differentiation: Silicic magmatism at the Organ caldera complex, New Mexico. Geosphere, 9(1),
pp.155-169, doi: 10.1130/GES00841.1.
Zindler, A. and Hart, S., 1986. Chemical geodynamics. Annual review of earth and planetary sciences,
14(1), pp.493-571.
Appendix A: Co-authored papers
Appendix A contains published manuscripts that I am a co-author of, completed during my Ph.D. program.
The first manuscript (Paterson et al. 2018) is a review of the different types of magmatic structures in
plutons, and their utility in documenting the magmatic histories of plutons. My contribution to this paper
includes a summary of localized compositionally defined structures (section 3.2) and the modern
approaches for studying magmatic structures (section 4.3). I also contributed to figure drafting and editing
of the manuscript for publication. The second manuscript (Alasino et al. 2019) is a field-based study of
schlieren-bound structures in the Sonora Pass Intrusive Suite (SPIS), Sierra Nevada, California, which are
interpreted to form from magma avalanching events at a margin-parallel solidification front. In this study,
I participated in fieldwork within the SPIS, wrote sections on the geological setting of the SPIS, described
field relationships and timing of schlieren-structure formation, and examined thin sections. I helped put
together figures and edited the manuscript for publication. The third manuscript (Barnes et al. 2020)
presents a new method to estimate the amount of crystal accumulation in plutonic rocks using trace elements
in hornblende. I contributed hornblende trace element data from the Cathedral Peak unit and compiled
whole-rock geochemistry from volcanic samples in the central Sierra Nevada to compare to the plutonic
estimates. I edited a figure for the appendix and edited the manuscript in preparation for publication.
212
Contents lists available at ScienceDirect
Journal of Structural Geology
journal homepage: www.elsevier.com/locate/jsg
A review of mesoscopic magmatic structures and their potential for
evaluating the hypersolidus evolution of intrusive complexes
Scott R. Paterson
a,∗
, Katie Ardill
a
, Ron Vernon
b
,Jiří Žák
c
a
Department of Earth Sciences, University of Southern California, 3651 Trousdale Parkway, Los Angeles, CA 90089-0740, USA
b
Department of Earth and Planetary Sciences, Macquarie University, Sydney, NSW 2109, Australia
c
Institute of Geology and Paleontology, Faculty of Science, Charles University, Albertov 6, Prague, 12843, Czech Republic
ARTICLE INFO
Keywords:
Fabric
Granite
Intrusive complex
Magmatic structures
Crystal sorting
ABSTRACT
The perception that plutonic bodies are structurally simple is disappearing with the recognition of an array of
magmatic structures useful for constraining hypersolidus temporal histories, evolving rheologies, strain fields,
flow directions, growth and cooling patterns, tilting and syn-emplacement tectonism. These histories provide a
powerful meansoftesting anarray ofgrowth, emplacement, chamber evolution and tectonic models.Important
points include that: (1) many structures form in “hydrogranular” or congested magma slurries during magma
mush avalanching, local convection and late hypersolidus strain challenging the notion that magmas must have
≤55% crystals to convect/fractionate and form compositional diversity in upper crustal magma chambers; (2)
magmaticfabricsreflecttransientstrainintheseslurriesratherthanflowdirectly:thelattermustbeinferredby
linking geometries with temporal and kinematic information; (3) caution is needed when using anisotropy of
magnetic susceptibility and similar quantitative tools that characterize, through a single ellipsoid, preferred
orientations of mineralgrainsgiven theincreasingrecognition ofmultiple fabricsinplutons; and(4) thatfuture
studies are particularly needed in plutons focusing on the distribution and styles of compositionally defined
structures, magmatic folding, shear zones and faults, multiple fabrics and of the physical/chemical behaviors in
hydrogranular slurries.
1. Introduction
The earliest study of mesoscale structures in plutons is often traced
back to at least the seminal works by H. Cloos (1925),E. Cloos (1936,
1946) and Balk (1937), who emphasized the value of studying mag-
matic foliation, lineation, layering and joints in plutons. Subsequently,
the conspicuous layering, unconformities and erosional features, in-
cluding 'cross-bedding' and magmatic shears in mafic–ultramafic in-
trusions attracted much attention (e.g., Wager and Brown, 1968;
Jackson, 1971; Irvine, 1980, 1987). These were explained by crystal
settling, marginal crystallization combined with rhythmic super-
saturation (e.g., McBirney and Noyes, 1979), and deposition from
crystal-mush density currents. Similar features were noted in grani-
toids. For example, Gilbert (1906) and E. Cloos (1936) described ac-
cumulations of alkali-feldspar megacrysts in single-mineral aggregates
andmodally-sortedlayersofschlieren,commonlywithgraded-bedding
and cross-bedding.
By 1977 (the initiation of Journal of Structural Geology), structural
analyses of granitoids were beginning to expand (e.g., King, 1964;
WagerandBrown,1968;PitcherandBerger,1972;Pitcheretal.,1985;
Didier, 1973; Moore and Lockwood, 1973; Wiebe, 1974), but still
lagged behind the depth and understanding of structural work in me-
tamorphic rocks, possibly because of the perception that many grani-
toids lacked mesoscale internal structures, and certainly because the
complex nature of the hypersolidus to subsolidus behaviors of plutons
(Fig. 1) inhibited a full understanding of plutonic structures. Great
advances have been made over the last 40 years, as the 'granite com-
munity' has meshed detailed field studies of igneous systems with
technologically-advanced structural, chemical and modeling tools.
Belowweprovideasurveyoftherichdiversityofmesoscale(visible
at outcrop scale, sometimes aided with hand lens) magmatic structures
in plutons (Fig. 2) and discuss their value for unravelling hypersolidus
histories of granitoids. To evaluate structures in plutons, one mustfirst
distinguish magmatic versus solid state structures and then determine
whether they were formed by internal or tectonic processes (Fig. 3).
After a brief introduction to these steps, we survey five groups
(Figs. 2–5) of magmatic structures: (1) internal contacts; (2) composi-
tionally defined structures; (3) preferred orientations; (4) deformation
structures; and (5) local indicators of growth, younging, vertical and
kinematics, followed by mention of outstanding challenges and a few
https://doi.org/10.1016/j.jsg.2018.04.022
Received 3 January 2018; Received in revised form 14 April 2018; Accepted 16 April 2018
∗
Corresponding author.
E-mail address: paterson@usc.edu (S.R. Paterson).
Journal of Structural Geology 125 (2019) 134–147
Available online 12 May 2018
0191-8141/ © 2018 Elsevier Ltd. All rights reserved.
T
213
particularly useful modern approaches for addressing these challenges.
The equally rich world of microstructures in plutons (visible micro-
scopically in thin sections) and of meso- and microstructures in the
coupled host rocks (e.g., Vernon, 2004; Passchier and Trouw, 1996)
will not be emphasized, although all three should be integrated in
practice.Weconcludewithinterpretationsonhowmagmaticstructures
form (Fig. 6) and their potential for unravelling hypersolidus growth
histories of intrusive complexes.
2. Deciphering magmatic vs. solid-state and internal vs. tectonic
processes
Thefirst step in evaluating structures in plutons is to determine the
rheological state of the deforming material when the structures formed
(Fig. 1). This is challenging because as the crystal-to-melt proportion
increases, magmas change from a dilute crystal suspension in melt, to
“hydrogranular” or congested magma slurries with a transient crystal
framework and connected melt network, and finally to isolated melt-
filled pores just above the solidus (Arzi, 1978; Vigneresse, 2008; Costa
et al., 2009; Brown, 2013; Bergantz et al., 2017). Thus, the magma's
rheological state undergoes a continuous evolution (Vigneresse et al.,
1996; Scaillet et al., 2000; Petford, 2009) that is sometimes broadly
divided into magmatic, submagmatic and solid-state fields (see below
and references in Paterson et al., 1989; Vernon, 2004).
Themagmatic rheologicfieldinvolves displacement ofmagmawith
enough melt for objects (crystals, enclaves, rock fragments) to move
and rotate into alignment without significant interactions nor de-
forming internally. At the mesoscale, the most robust observations for
recognizing structures formed during magmatic flow are those asso-
ciatedwith:(1)igneousshapes(euhedralforselectedcrystals),internal
crystal features (concentric zoning, growth twinning; e.g., Streck,
2008), and crystal distributions (clustered and anticlustered; e.g.,
Jerrametal.,2003);(2)plutonicstructuresofidenticalagethatintrude
or crosscut them; (3) growth of magmatic features (comb layering,
orbicular structure; Moore and Lockwood, 1973); (4) evidence of high
strainwithnosolid-statemicrostructuresincrystals,implyingthatmelt
Fig. 1. Schematic diagram showing the complex transition from magmatic to solid state behavior, its dependence on magma crystallinity, strain rate, and changing
deformationmechanisms(colorcodedfromlowerlefttoupperrightofdiagram).Magmatic,plasticandbrittleprocessescanoccursimultaneouslyinthetransitional
'submagmatic'zone,particularlyatdifferentstrainrates.Commonstructuresineachtransitionalzonearedepictedasfollows:(a)euhedralcrystalswithnopreferred
orientation; (b) magmatic foliation; (c) magmatic shear; (d) brittle fracture of minerals; (e) mineral-mineral interactions forming synneusis or glomerocrysts; (f)
minerals with melt-filled fractures; (g) deformation twinning; (h) veins and fractures cross-cutting minerals; (i) chessboard subgrain formation in quartz; (j) de-
formation creep in phenocrysts. Three dashed lines show possible paths of tectonically driven magma deformation (slow strain rates), internal magmatic processes,
and processes during eruption (fast strain rates).
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
135 214
takes up the strain (Park and Means, 1996); and (5) contemporaneous
meltrelocationstructures(e.g.,meltpocketsinlowstresssites;Hibbard
and Watters, 1985). Microstructural criteria for magmatic flow are
summarized by Paterson et al. (1989) and Vernon (2000, 2004).
The submagmatic rheologic field involves flow/strain of hydro-
granular magma slurries with transient crystal frameworks and melt
networks by a complex combination of processes, such as melt-assisted
grain-boundary sliding, grain-boundary migration assisted by contact
melting, strain partitioning into melt-rich zones, melt-enhanced em-
brittlement, high-temperature intracrystalline plastic deformation, and
transfer of melt into sites of low mean pressure (Paterson et al., 1989;
Vigneresse et al., 1996; Vernon, 2004; Bergantz et al., 2017). Discrete-
element numerical modeling by Bergantz et al. (2017) and Schleicher
et al. (2016) elegantly display the complexfluid to visco-plastic beha-
vior of these flows, including the ephemeral development of 'crystal
force chains' (see also Philpotts and Dickson, 2000) that can transmit
stresses, drive intracrystalline deformation and change local melt
pressures. This field is the most challenging to recognize at the me-
soscale, since it involves a mixture of both magmatic and solid-state
features across short spatial and temporal scales. Perhaps the best
Fig. 2. Photos of mesoscale magmatic structures (scale: rulers = 15 cm, hammer∼22 cm, notebook∼15 cm): (a) cross-cutting contacts and internal layering; (b)
gradedschlieren;(c)differentstylesoflayeringincludingfromrighttoleftleucogranitedike,enclaveswarm,schlieren,troughs,schlieren+felsiclayers, “ridgeand
pillar”;(d)migratingtubestructureandreintrusionbyhostmagma;(e)diapir;(f)cross-cuttingschlierentroughs,locallyfolded(whitearrow)orfaulted(redarrow),
heightofoutcrop∼5m;(g)plumewithschlieren-richheadandgradationaltail,arrow=2Dmovementdirection;(h)maficellipsoid;(i)magmatic “dishandpillar”
structuresduetoupwardmeltmigration;(j)feldsparUSTstructure;(k)foldedplagioclase+amphibolemagmaticfoliationwithnewmagmatichornblendeparallel
toaxialplane;(l) foldedaplitedike(thindashedline)withmagmaticaxialplanarfoliation(thickline)definedbyplagioclase, biotite,hornblendeandenclaves;(m)
two magmatic preferred orientations at a single outcrop. Older enclave alignment = solid line. Younger plagioclase, hornblende, biotite alignment = dashed line.
Enclave tips bending from old to new; (n) magmatic mullion at margin of rhyolite plug intruding marine metasediments; (o) K-feldspar megacryst cluster; (p) pipe-
shaped K-feldspar megacryst cluster; (q) complex normal faulting, trough truncation, and folding of schlieren. Comb layering evident in some light-colored layers.
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
136 215
criteria are the evidence of patches of crystal plasticity or elastic phe-
nomena (subgrains, faulting) contemporaneous with melt relocation
phenomena. For example, the presence of primary mineral aggregates
(namely quartz, alkali-feldspar, and sodic plagioclase) in con-
temporaneous fractures(Bouchezetal.,1992;Vernon,2004),indicates
that late-stage melt travelled to low-pressure sites during brittle de-
formation of crystals.
At the mesoscale, solid-state deformation is characterized by: (1)
evidence of grain-shape change and grainsize reduction by internal
deformation and recrystallization; (2) lenticular recrystallized ag-
gregates and 'tails' of minerals; (3) structures passing through solid
xenoliths and/or continuous with structures in solidhost rock; (4)fine-
grained folia anastomosing around less deformed 'relic' aggregates; (5)
boudinageofstrongmineralsorclasts,withrecrystallizedaggregatesof
weaker minerals (e.g., quartz and mica) between the boudins; (6) het-
erogeneous strain with local mylonitic zones; and (7) fractures and
cataclasites.Microstructuralcriteriaforsolid-stateflowaresummarized
by Paterson et al. (1989), Passchier and Trouw (1996) and Vernon
(2000, 2004).
Once hypersolidus structures are identified and characterized,
consideration must be given to whether they reflect localized internal
processes within an active magma chamber or reflect regional tectonic
deformation imposed on a magma chamber. Magmatic and submag-
matic structures can form by either process, since regional deformation
can act on a crystallizing magma chamber. We view 'magma chambers'
as structural geologists, that is not just eruptible magma, but any mush
that contains enough interconnected melt to allow magmatic to sub-
magmatic structures, including late fabrics to form. Two reliable cri-
teria for recognizing tectonically-formed hypersolidus structures in-
clude: (1) coupling or continuity with regional tectonic structures in
host rock (Fig. 3c); and (2) continuous structures crosscutting or
overprinting internal contacts and/or xenoliths, indicating that magma
pulse juxtaposition predated the structure (e.g., Paterson et al., 1998).
Pervasive solid-state ductile structures in plutons are typically caused
by regional deformation due to the mechanical challenge of lower
viscosity magmas deforming higher viscosity solid rock (e.g., Ramsay,
1989; Johnson et al., 2003).
Complications can occur if regional deformation destabilizes
growingcrystalmushzonesinchambers,whichthencollapseandform
local magmatic structures, such as the magmatic folds, faults and
troughs. We know of no reliable criteria to easily identify such cases as
being of tectonic origin beyond the permissive observation of being
spatially linked to a syn-intrusive fault (Emelus and Troll, 2014).
3. Inventory of magmatic structures
Space restrictions only allow a brief introduction below of 5 pro-
minent groups of structures that are the most common and most useful
for unravelling hypersolidus granitoid histories (Fig. 2).
3.1. Structures defining regional internal contacts and layering
Incrementalemplacementofcompositionally-distinctmagmapulses
inintrusivecomplexesresultsinarangeofmagmaticstructures,suchas
dikesandgradationaltosharpunitcontacts(Fig.2a,b,c,o;Pitcherand
Berger, 1972; Hardee, 1982; Paterson et al., 2011). Flow, mixing and
mingling along these contacts commonly leads to hybrid zones and
layered intrusive sequences (e.g., Fernandez and Gasquet, 1994; Wiebe
and Collins, 1998; Barbey, 2009).
Internal contacts are also defined by a wide range of igneous
layering, particularly schlieren, in granitoids, ranging from kilometer-
scale layered zones (Fig. 2b; Lucus and St-Onge, 1995; Clarke and
Clarke, 1998; Žák et al., 2009; Burgess and Miller, 2008) in which
layering maintains subparallel orientations, mineral grading and local
cross-bedding, to more complexly-shaped, 10cm-to 10m-scale struc-
tures discussed below (Fig. 2d, e, f, q; e.g., Barrière, 1981; Parsons,
1986; Paterson et al., 2008; Pinotti et al., 2016). Mafic–ultramafic
complexes also are characterized by both internal intrusive contacts,
compositional layering of cumulates, and local 'unconformable' struc-
tures resembling cross-bedding (Fig. 2; e.g., Wager and Brown, 1968;
Irvine, 1980; Naslund and McBirney, 1996). These layers can be mod-
ally, or grain and density sorted, and have graded or sharp contacts.
Fig. 3. Cartoon showing three 'endmember' preferred orientation patterns in plutons (pink) and surrounding host rock modified from Paterson et al. (1998). (a)
Completely decoupled system (internally complex); (b) partially coupled system (margin parallel); (c) completely coupled system (rectilinear or folded). Black
lines=magmatic foliation, lineation, or folds. Grey lines=host rock structures. Filled dots=vertically plunging lineations; arrows=plunging lineations.
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
137 216
Fig. 4. Key elements used to define 'compositionally defined' structures (modified from Paterson, 2009). (a) Schlieren-bound troughs showing truncations, mineral
grading, local mineral clusters (Photos Fig. 2f, p); (b) stationary tubes; (c) migrating tubes (Fig. 2d); (e) mesoscale diapir (Photo Fig. 2e); (f) plumes with schlieren
'heads' and open 'tails' (Fig. 2g); (g) mafic ellipsoids with internal margin parallel foliation (Fig. 2h); (h) irregular mineral clusters; (i) pipes.
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
138 217
Fig. 5. Cartoon of local, mesoscale structures used to infer growth and younging directions, paleovertical, and kinematics of magma movement (the latter modified
fromPhilpottsandAsher,1994):(a)cross-cutting,schlieren-boundtroughsandmineralgrading(PhotosFig.2c,f,p);(b)asymmetric(chilledversushybrid)contacts
after Wiebe and Collins (1998). Enclaves are flattened near chilled margin and irregular near hybrid; (c) undeformed tube or pipe axes; (d) concave up enclave
channels with enclaves molding minerals at channel base after Wiebe and Collins (1998); (e) magmaticflame structures and load casts; (f) layering deflected by
sinking block; (g) (g) granophyre segregations attached to phenocrysts; (h) imbricate minerals; (i) crystal orientation patterns due toflow in dike; (j) broken and
sheared phenocrysts; (k) granophyre wisps and back-folding by differentialflow; (l) granophyre wisps; (m) magmatic S–C structures; (n) ramp structures; (o) shear
bands;(p)vesicleselongatedparalleltotheaxialsurfaceoffoldedflowlayers;(q)asymmetricmicro-foldsinflowlayeringformedduringrotationofaphenocrystor
glomerocryst; (r) Riedel shears; (s) asymmetric folds of magmatic foliation or layering.
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
139 218
Some have non-planar geometries, akin to schlieren-defined structures
in silicic plutons.
3.2. Localized compositionally defined structures
An increasingly diverse collection of compositionally defined, local
magmatic structures are recognized in granitoids (Figs. 2 and 4). Many
of these are partly characterized by schlieren, that is, centimeter-to
meter-scale layers typically with sharp basal contacts, diffuse upper
contacts and downward modal increases of dense major and accessory
minerals (e.g., Barrière, 1981; Reid et al., 1993; Burgess and Miller,
2008; Paterson, 2009). Minerals within schlieren layers are strongly
aligned parallel to the basal schlieren orientation (Fig. 2). Schlieren
form in widely ranging orientations, including vertical (Wiebe et al.,
2007; Žák and Klomínský, 2007; Paterson, 2009), indicating that hy-
drogranularflow sorting of crystals, and not just gravity, plays an im-
portant role in schlieren formation.
Schlieren-bound structures include the following:
(1) Troughs (Figs. 2f, q and 4a) are open-ended, schlieren-bound
channels comprised of multiple stacked and cross-cutting layers
(e.g.,WagerandBrown,1968;McBirneyandNoyes,1979;Barrière,
1981; Reid et al., 1993; Burgess and Miller, 2008). The perpendi-
cular direction to a layer truncation defines a younging or trough
growth direction (Fig. 2).
(2) Tubes (Figs. 2d and 4b, c) are closed, cylindrical, schlieren-bound
channels in 3D, with channels either nested inwards (stationary
tubes; Fig. 4b) or forming dike-like zones of crescent-shaped,
crosscuttingschlieren(migratingtubesorladderdikes;Fig.4c;Reid
et al., 1993; Weinberg et al., 2001; Žák and Klomínský, 2007;
Paterson, 2009; Dietl et al., 2010). Schlieren truncations provide
younging or migration directions of tube margins. Some tube
structures are re-intruded, rotated and broken during and after
formation(e.g.,Paterson,2009;Hodgeetal.,2012).Aswithtrough
structures, the younging direction of tubes may show systematic
patterns across plutons (Memeti et al., 2014; Ardill et al., 2017;
Wiebe et al., 2017).
(3) Mesoscale diapirs (Figs. 2e and 4d) are 10cm-to 10m-scale, ellip-
soidal to tear-drop shaped, compositionally variable and distinct
from the host magma, and with narrow tails and schlieren-bound
'heads'.Theyhavebeendescribedinbothmafic(LarsenandBrooks,
1994) and felsic (Paterson, 2009) plutons. The schlieren at the
diapir head is thought to form byfilter pressing and accumulation
of mafic and accessory minerals (Weinberg et al., 2001). The trend
of a line through the tail and symmetrically dividing the head
provides a 2D movement direction. Where examined in detail,
diapir'movementdirections'arefoundinallorientationsandsoare
not simply buoyancy driven (Weinberg et al., 2001; Paterson,
2009).
(4) Plumestructures(Figs.2gand4e)havemodallydistinct(from host
magma), schlieren-bound 'heads' and broad tails, sometimes mi-
micking mushroom or jellyfish shapes, that compositionally grade
Fig. 6. Cartoon depicting mesoscale mag-
matic structure formation in an in-
crementally grown (purple units) crystal-
lizing magma chamber intruded by new
magma pulses (red and orange); modified
from Paterson et al. (2016). Crystal-rich
mush zones form and collapse along mar-
gins of older pulses. Arriving new pulses
drive convection and 'returnflow' (grey ar-
rows) in chamber. Examples of how in-
dividualstructuresmayforminthissystem:
(a) cross-cutting troughs during margin
collapse; (b) migrating tubes during con-
vection; (c) diapirs and plumes (d) moving
sideways as new pulse enters chamber; (e)
mafic ellipsoids and cognate inclusions
formed from collapsing or mingling margin
material with new pulse; (f) mineral clus-
ters duringflow sorting; (g) pipe structures
by flow sorting during rising or sinking of
magmas; (h) magmatic faults cooling/
shrinking and collapsing margins; (i) dikes
and veins as late melts migrate through
marginal cracks.
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
140 219
into the host granitoid. Their broad, gradational tails make them
distinct fromdiapirs and indicate that the source of plume material
is the host magma. The trend of a line symmetrically splitting the
tailandtheheadprovidesa2Dmovementdirection,whichincases
where measured also have variable orientations.
(5) Mafic ellipsoids (Figs. 2h and 4f) are 1–3m long, elongate 3D el-
lipsoids of fine-grained, more mafic plutonic rock distinct from
theirhost,withaweakschlierenborder(Memetietal.,2014).They
canresemblemicrogranitoidenclaves, exceptfortheschlieren rims
and because mineral alignment within the mafic ellipsoids usually
defines a well-developed margin-parallel pattern.
(6) Mineral clusters (Fig. 2o and p, 4g) range from millimeter-scale
joined crystals (synneusis of Vance, 1969), centimeter-scale crystal
clusters and glomerocrysts (Seaman, 2000; Cashman et al., 2017),
to decimeter-scale clusters with highly varied shapes, from irre-
gular, dike-shaped, trough-shaped, to pipe-like, (Paterson et al.,
2005; Memeti et al., 2014). Mesoscale clusters are dominated by a
single mineral, but may include microgranitoid enclaves and xe-
noliths, and may or may not have bounding schlieren. Mineral
preferred orientations in clusters are highly variable.
(7) Pipes(Figs.2oand4h)areclosed,cylindrical,oftensub-vertical3D
structures with little internal layering. Pipe compositions are typi-
cally dominated by more evolved (buoyant) compositions (Wiebe
and Collins, 1998) or by one mineral and thus a type of mineral
cluster (e.g., Paterson, 2009). Their margins with the host magma
may or may not be marked with weak schlieren. Internal mineral
orientation occurs parallel to pipe margins, although these folia-
tions may rotate to 'regional' fabric orientations in pipe centers.
Rocheretal.(inpress)describemeter-scale,verticallyelongated K-
feldspar accumulations with inward-dipping concentric foliation
and steeply plunging lineations in their upper parts that change to
flat foliation with enclaves at their lower ends.
3.3. Structures defined by preferred orientation
Strain of crystal mushes (Benn, 1994) caused by gradients during
magmaflow or tectonic deformation, lead to the alignment of crystals,
microgranitoid enclaves and xenoliths (Fig. 2k, l, m), as well as trans-
position of older magmatic structures during folding and faulting,
leading to the development of one or more magmatic foliations and
lineations (Fig.3).Acommonmisconception isthatonlyasinglefabric
(foliation ± lineation) forms in intrusive complexes, whereas close
inspection often establishes the presence of two or more sets of fabrics
thatmayhavecontemporaneousoroverprintingrelationships(e.g., Žák
et al., 2007; Kratinová et al., 2010; Macchioli Grande et al., 2015;
Alasino et al., 2017). Different mineral populations may show different
intensities of alignment and/or different orientations (Blumenfeld and
Bouchez, 1988; Schulmann et al., 1997; Paterson et al., 1998). Fur-
thermore, microgranitoid enclaves may have internal mineral align-
ments parallel to, or distinct from, alignments in the host matrix and
parallel to, or distinct from, the enclave long axis (Paterson et al.,
2004). Xenoliths may preserve older magmatic or metamorphic struc-
tures.Locallayeringorothermagmaticstructuresdiscussedabovemay
have mineral alignments distinct from preferred orientations in nearby
host granitoids (Paterson, 2009; Alasino et al., 2017).
Magmatic fabric patterns may or may not show refraction across
compositional boundaries (Compton, 1955; Hines et al., in press), be
parallel to fold axial planes and fold axes (Pitcher and Berger, 1972)
andshowdeflectionsinmagmaticshearzones.Regionalpatternscanbe
quite varied, although Paterson et al. (1998) defined three simplified
'endmember' patterns: (1) internally complex, decoupled (Fig. 3a); (2)
margin parallel or partially coupled (Fig. 3b); and (3) rectilinear, in-
ternally folded and coupled with host rock (Fig. 3c. Type #1 and 2
foliationpatternscanlocallyshowsteepeninglineationpatterns,which
have been used to infer 'magma feeder zones' (e.g., Vigneresse and
Bouchez, 1997; Olivier et al., 1997; Talbot et al., 2004). Careful
observation of these relationships has the power to document complex
strain histories in magma chambers (e.g., Fernandez, 1988; Žák et al.,
2007; Cao et al., 2015).
3.4. Deformation-related structures
Once formed, all magmatic structures can be altered by continued
deformation, resulting in transposition of these older structures during
formationofavarietyofmagmaticfolds,faultsandshearzones(Fig.2).
Both brittle and ductile faults can develop in magma (Hibbard and
Watters,1985;Dingwell,1997;Katzetal.,2006;Bergantzetal.,2017).
Magmatic brittle faults and 'ductile' shears typically do not show high
strain, elongate crystal shapes and strong crystal preferred orientations
(Fig. 2f, q), since most displacement is accommodated by melt (Park
and Means, 1996; Paterson et al., 1998). Faults also may draw in ad-
ditional melt (Geshi, 2001; Pinotti et al., 2016), become sealed by
continuedcrystalgrowth(ParkandMeans,1996)andthereforemaybe
cryptic, unless offset markers are present.
Although magmatic folds are briefly described in several publica-
tions (Brown et al., 1986; Harm, 1991; Smith and Houston, 1994;
Paterson et al., 1998), detailed studies of magmatic folds are rare. Ex-
amples of folded layering, dikes, enclaves, and magmatic foliations are
common in bothvolcanic (e.g., Vernon, 1987; Iezziand Ventura, 2000;
Castro et el. 2002; Ventura, 2004) and plutonic rocks (e.g., Rosenberg
etal.,1995;Patersonetal.,1998;Yoshinobuetal.,2009).Variablefold
profiles, fold-related geometric and mineral lineations, presence or
absence ofanaxialplanar foliationand refoldingalloccur (Fig.2f,k, i,
q).
Mullions, another response to strain (Fletcher, 1982; Price and
Cosgrove, 1990; Urai et al., 2001; Schmalholz and Schmid, 2012), also
occurinmagmaticrocks(Patersonetal.,1998;Yoshinobuetal.,2009),
and would benefit from further detailed studies. Fig. 2n shows a mul-
lion formed at the contact between an Ordovician rhyolite plug and
marine metasediments in the Chaschuil region, NW Argentina (Lusk
et al., 2017). Magmatic mullions have been described along dike-host
rock margins and in some cases the tip of the mullion can tighten and
feed veins intruding subparallel to the axial planes of host rock folds
(Patersonetal.,1998;Yoshinobuetal.,2009).Mullionaxesareparallel
to regional fold axes and the shortening directions implied by the
mullions are consistent with regional shortening implied by the host
foliation. These observations document syn-magmatic regional de-
formation during which the magma mush had a higher effective visc-
osity than its host.
3.5. Magmatic indicators of local solidification, younging, paleovertical,
and kinematics
A variety of local magmatic structures provide additional informa-
tion about structures, such as their initial orientation, growth/crystal-
lization direction and kinematics of associated deformation.
Unidirectional solidification (or crystallization) structures (Fig. 2j)
(UST'sofShannonetal.,1982)include:(1)comblayering;(2)crenulate
layering; (3) dentritic growth patterns, similar to crescumulate and
spinifex microstructures; and (4) intergrowth layering. Besides cross-
cutting intrusive boundaries (Fig. 2a), mesoscale magmatic structures
available to provide local growth directions include: (1) local cross-
cutting boundaries in compositionally defined structures (Fig. 2d, f, q,
5a); (2) mineral grading (Fig. 2b, f, 5a); (3) magmatic load cast and
flame structures (Fig. 5e; Puziewicz and Wojewoda, 1984); (4) asym-
metric contact characteristics (Figs. 2b and 5b), such as the chilled
versus hybrid margins of Wiebe and Collins (1998); and (5) enclave
channels (Figs. 2c and 5d) with mineral molding at base (Wiebe and
Collins, 1998). Proposed 'way-up' or 'paleo-vertical' indicators include:
(1) load casts, flame structures (Puziewicz and Wojewoda, 1984;
Vaughan et al., 1995); (2) block sinking patterns in non-convecting
magma (Fig. 5f; Paterson et al., 1998; Emeleus and Troll, 2014); (3)
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
141 220
axes of undeformed pipes and tubes (Fig. 2d, p, 5c; Wiebe and Collins,
1998; Paterson, 2009), and (4) 'dish and pillar' structures during which
rising melts deflect layering upwards (Fig. 2i). Kinematics caused by
deviatoric strain during movement of magma, can be established from:
(1) crystal orientation patterns defining S–C structures, shear bands,
andhyperbolicpatternsindikes(Fig.5i,m,n,o;Komar,1972;Shelley,
1985; Vernon, 1987; Miller and Paterson, 1994; Correa-Gomes et al.,
2001; Geoffroy et al., 2002); (2) imbricate phenocrysts (Fig. 5h;
Blumenfeld and Bouchez, 1988); (3) broken, sheared and rotated phe-
nocrysts (Fig. 5g, q) (Vernon, 1987; Blumenfeld and Bouchez, 1988);
(4) granophyre wisps (Fig. 5i; Philpotts and Asher, 1994); (5) asym-
metricfolds(Fig.5s);(6)segregationsattachedtophenocrysts(Fig.5g);
(7) Riedel shears (Fig. 5r); and (8) ramp structures (Fig. 5n; Philpotts
and Asher, 1994).
4. Discussion
4.1. Hypersolidus structural histories
The wealth of magmatic structures described above emphasizes
that,comparabletometamorphicstructures,agrowingsetofstructural
'tools' exist for unravelling hypersolidus histories. For example, cross-
cutting or overprinting structures help to establish temporal histories.
Layering, foliation patterns, folds, and faults all provide information
about hypersolidus strain fields. And when combined with kinematic
informationtheycanbeusedtoinfermagmaflowdirections.Structures
indicating paleovertical in chambers can be used to evaluate local
magmatic tilting of layers or tectonic tilting of entire intrusive com-
plexes. There also is great potential for constraining changing rheolo-
giesandevolvinginternalgradientsaswebetterunderstandstructures,
particularly the non-vertical motions of structures like diapirs and
plumes and the mechanisms and kinematics of magmatic folding and
faulting. The increasing mesoscale evidence of structures younging to-
wards contacts with older units (unit ages established through geo-
chronology) and statistically preferred migration directions are likely
driven by internal motions of magma during incremental growth and
crystallization of intrusive complexes. Finally, in our experience, all
plutonshaveatleastlimitedmagmaticstructures(e.g.,Bouchez,1997),
although some intrusive complexes are much richer in magmatic
structures than others (e.g., Weinberg et al., 2001; Pinotti et al., 2016).
Why this is so will provide exciting information about the construction
and behavior of these systems. Thus, the results of magmatic structural
studies provide powerful constraints on models of magmatic systems
and need to be fully incorporated into geochemically dominated dis-
cussions of these systems.
4.2. Key issues about magmatic structures
4.2.1. Hypersolidus deformation mechanisms
Oneunique aspectofmagmatic structuresis thattheyformacrossa
wide range of temperatures, effective viscosities and deformation me-
chanisms associated with evolving physical and chemical processes
(e.g., Cruden, 1990; Rosenberg and Handy, 2005). It has been tradi-
tionaltoidentifykey,ratherabrupt,transitionsinthesebehaviors(e.g.,
van der Molen and Paterson, 1979; Vigneresse et al., 1996; Rosenberg
and Handy, 2005), and to conclude that magma with less than ca 55%
melt becomes 'static', which we believe typically means to others that
magma is hard to erupt or vigorously convect. However, many of the
preserved structures noted in this paper likely formed during this
transition from 55% to<20%. Nicolas and Ildefonse (1996), Park and
Means (1996) and Paterson etal. (1998) arguethat magma can deform
with even less than 20% melt through grain alignment and sliding on
melt films aided by dissolution or melting at grain impingements
(called contact melting by Park and Means, 1996). Benn (1994) es-
tablished that only small increments of strain are needed to form
magmatic fabrics. And Bergantz et al. (2017) and Schleicher et al.
(2016) show the complex simultaneous, viscous–plastic–brittle beha-
vior of deforming hydrogranular crystal mushes and argue against
abrupt rheological transitions. Therefore, a key issue is to recognize
that crystal mushes can continue to deform at low melt proportions
during which melt-aided, crystal-plastic, and even brittle deformation
mechanisms increasingly overlap (Bergantz et al., 2017; Holness et al.,
2017).
4.2.2. Layering
Attempts to categorize layering (e.g., Wager and Brown, 1968;
Irvine, 1980; Barbey, 2009; Pinotti et al., 2016) collectively suggest
that the following styles form by processes that often occur simulta-
neously in growing plutons: (1) intrusion of newly arriving magma
pulses or of locally derived magma from other parts of the chamber
(e.g., late melts forming dikes) resulting in internal contacts and/or
planar zones of mingling and mixing (e.g., layering in hybrid zones
along contacts; Fig. 2a; Hutton, 1992; Paterson et al., 2008; Barbey,
2009); (2) fractionation driven by hydrodynamic and gravity-driven
sorting and crystal–melt separation resulting in modally defined
layering (Fig. 2b; Irvine, 1980; Barrière, 1981). Petford (2009) also
noted that igneous layering may form during increasing shear rate and
shear thinning in congested magma slurries when crystals arrange
themselves into layers parallel with the mean flow direction, a phe-
nomenon observed both experimentally and in computer simulations
(Stickel and Powell, 2005); (3) thermally/chemically driven fractiona-
tion and flow forming diffuse layering (Boudreau, 1995); (4) de-
formation-assisted (magmatic or tectonic) strain (e.g., attenuated en-
claves forming schlieren), crystal-melt segregation, or collection of
crystals and rock fragments into planar zones (Fig. 2c, o, q; Tobisch
et al., 1997); and (5) chemical and textural mineral modifications of
evolving layers as bulk compositions and environmental conditions
shift (Boudreau, 1995).
Onlythefirstoftheprocesseslistedabovetoforminternalcontacts/
layers involves the arrival of a new pulse of magma. A challenge for
futuresstudiesistoapplycriteriaforestablishingwhichcontactsreflect
juxtaposition of new pulses versus internally formed boundaries (e.g.,
Miller and Paterson, 2001). Thus two critical issues regarding internal
contacts/layering are the following: (1) that many internal contacts
form by processes in already emplaced magma and are not contacts
between arriving pulses; and(2) thedegree towhich internal contacts/
layering can be removed.
Some recent publications examining incremental growth of in-
trusive complexes conclude that evidence for the numerous original
intrusive contacts have been removed by subsequent processes (e.g.,
Colemanetal.,2004,2012).Totestthisidea,itisimportanttoexamine
the processes proposed to remove contacts. Older internal contacts/
layering may be removed by: (1) being entirely transported out of the
map section or recycled by a newly arriving magma pulse; (2) wide-
spread convection and homogenization in a magma chamber; (3) re-
working of contacts during widespread subsolidus deformation and
metamorphism; and (4) pervasive annealing during widespread near-
solidus, thermally driven (possiblyfluid aided) recrystallization.
Evidenceofmagmaticerosionresultinginsharptruncationsofolder
structures along new intrusive contacts, the presence of cognate blocks
of older phases in younger, and the presence of antecrysts in younger,
particularly compositionally uniform units support #1 and #2
(Paterson et al., 2016; Gaschnig et al., 2017). Widespread convection
(#2) is supported by the extensive dispersal of foreign objects (zircon
antecrysts, microgranitoid enclaves, chemically distinct mineral popu-
lations) throughout new pulses and homogenization of the originally
distinct pulses. The above mechanisms require a large, active magma
chamber.
Removal of magmatic structures without an active chamber can
occur by #3 or #4. Extensive deformation and recrystallization at
amphibolitefaciesconditions(#3)resultsineasilyidentifiablegneisses
withstrongmetamorphicfabricand/orcompositionalbanding,andloss
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
142 221
of igneous microstructures, such as oscillatory zoning. Grains are
characterized by polygonal to irregularly xenoblastic quartz and feld-
spar, and rounded inclusions of quartz in feldspar and vice versa
(Vernon, 2000, 2004).
Conceivably, contacts could be obscured by #4. Grain boundary
adjustments and mineral reactions are recognized in mafic/ultramafic
magmas where temperatures are significantly higher (Boudreau, 1995;
Holness et al., 2017). If so, these processes should remove mineral
primary igneous microstructures such as oscillatory zoning, either by
grain-boundary migration or advance of a diffusion front. Although
limited mineral alteration has been documented in granitoids (Barnes
et al., 2017), considerable doubt exits that modifications extensive
enough to remove evidence of older structures occurs without wide-
spread subsolidus deformation (Vernon and Paterson, 2008b; Holness
et al., 2017). Furthermore, the common preservation of euhedral grain
shapes, euhedral inclusions, disequilibrium dihedral angles, and oscil-
latory zoning in the minerals in granitoids lacking solid-state de-
formation is strong evidence that extensive subsolidus grain-shape
changes did not occur during cooling of granitic magmas (Vernon and
Paterson, 2008b; Holness et al., 2017). Finally, in typical granitoid
plutons, internal structures, formed during both the construction of the
plutons and subsequent internal processes, are typically well preserved
(Fig. 2). This is true even for the granitoids emplaced at amphibolite
facies conditions. Consequently, magmatic structures are unlikely to be
obliterated by intergranular changes, unless pervasive subsolidus tec-
tonic deformation occurs.
4.2.3. Compositionally defined magmatic structures
It is our expectation that knowledge of compositionally defined
structures will continue to grow, both in the recognition of new types
and in a better understanding of formation. Mechanisms of formation
presently remain controversial (e.g., compare Weinberg et al., 2001;
Paterson, 2009; Hodge et al., 2012; and Wiebe et al., 2017). Although
crystallization, fractionation, and mixing certainly play a role in
formingcrystalsorthemeltsfromwhichcrystals grow,wesuggest that
existing evidence favors that physical processes dominate in the for-
mation of these structures: below we present some of the more popular
ideas for their formation.
Wiebe and Collins (1998), Weinberg et al. (2001), Paterson (2009)
suggestedthattubesandpipesrepresentcylindricalchannelsofmagma
flow through crystal mushes, and that the main flow direction is par-
alleltothetubewalls.Schlierenalongthetubewallsformbycombined
flow sorting andfilter pressing. Weinberg et al. (2001) also noted that
coarse K-feldspar aggregates occurred on the upstream side of 'dike
necks' (pipes), as inferred from crosscutting relationships in the dike.
They suggested that flow necking (a consequence of melt extraction,
crystal growth and flow sorting) led to a 'megacryst logjam’, which
continued to grow by thefiltering out of megacrysts upstream (see also
Clarke and Clarke, 1998; Rocher et al., in press). Magmaflow through
tubes and pipes, potentially either up or down, can be explained by
thermal or compositional buoyancy (e.g., Griffiths, 1986; Martin et al.,
1987; Weinberg et al., 2001), Raleigh–Taylor fingering (Wiebe and
Collins, 1998; Perugini and Poli, 2005; Schofield et al., 2010), rising
volatile-rich magmas (Clarke et al., 2013; Weinberg et al., 2001;
Memetietal.,2014)orfallingobjects(Castroetal.,2008;Rocheretal.,
in press).
Schlieren-bounded troughs are generally viewed as open (towards
more crystal poor magma) mush channels analogous to flow channels
witha'bed-load'andformedinaporousmedia(e.g.,Wahrhaftig,1979;
Dufek and Bergantz, 2007). Magmaflow is interpreted to be parallel to
trough axes, inferred by aligned hornblendes in basal schlieren in
granitoids(Paterson,2009)oraligned,elongateplagioclaseandolivine
parallel to the trough axes in mafic rocks (Holness et al., 2017). The
intra-chamber channel flow is potentially caused by the disturbance
andcollapseofgrowingmagmamushzonesalongmargins(Humphreys
and Holness, 2010; Marsh, 2015) or erosion of previous pulses
(Paterson et al., 2016), and avalanching of crystal piles down into re-
gionsofthechamberwithlowereffectiveviscosities(e.g.,Irvine,etal.,
1998; Bergantz, 2000; Žák and Paterson, 2009).
Mesoscale diapirs and plumes show all the characteristics of
Raleigh–Taylor instabilities: but one surprising recent observation has
been the wide variation in movement directions of these structures,
many of which show non-vertical directions (Paterson, 2009). Alter-
natively, they may record local flow gradients/displacement fields
(stress, strain, rheology) in crystal mush systems, potentially driven by
either regional deformation of magma chambers and/or convective
movement in reservoirs during arrival of new magma batches.
Maficellipsoidshaveonlyrecentlybeenrecognizedassuch(Memeti
et al., 2014) but may be comparable to microgranitoid enclaves and
largely formed by magma mingling. The greatest uncertainties are the
original source of magmas, the type of structure prior to mingling
(mafic dikes, pipes, sheets) and where initial mingling occurred (e.g.,
Barbey et al., 2008; Paterson et al., 2016).
Although individual types have been examined, the full range of
sizes and shapes of mineral clusters have not been systematically stu-
died. Many are thought to have formed by the combined physical
processes of nucleation and surface energy effects, flow sorting, and
filter pressing (e.g., Vance, 1969; Barrière, 1981; Jerram et al., 2003;
Vernon and Paterson, 2008a). This clustering also implies that stress
transfers between crystals, leading to crystal plasticity, may be more
common in these hypersolidus systems than previously recognized.
Petford (2009) and Bergantz et al. (2017) summarized processes
such as the following that occur in flowing hydrogranular magma
slurries because of their complex times-transgressive rheologic beha-
vior: (1) crystal disorder-order transition to form compositional
layering parallel to the flow direction; (2) either crystal jamming
(forming clumps) or dilatancy (drawing in melts); and (3) transient
particle pressures that fluidize crystal mushes resulting in an in-
stantaneous pressuredecreaseandinsitububbleformationinthemelt,
promoting chamber-wide instability and formation of melt pipes and
channels, disrupted layering, crystal clustering, synmagmatic faulting,
andcollapseandremixingofmushzones.Thesephenomenacouldform
a number of the magmatic structures listed above.
4.2.4. Preferred orientations of mineral grains
Two outstanding issues regarding the interpretation of preferred
orientations are the following: (1) the need to recognize that they re-
flecttransienthydrogranularstrainandareonlyparalleltoflowplanes/
directionsundercertainconditions;and(2)thattherearemanyplutons
with more than one foliation/lineation. Thefirst issue is discussed ex-
tensively by Paterson et al. (1998) who noted that foliation/lineation
patterns (Fig. 3) must be combined with kinematic and other in-
formation to establish flow patterns. And Bergantz et al. (2017) ex-
amine what fabrics record at the mineral population scale. The re-
cognition of multiple fabrics in granitoids is typically ignored in AMS
andotherquantitativefabricanalysesinwhichoneellipsoidperstation
is determined. This forces multiple preferred orientations to be aver-
aged and the loss of information about temporal histories. These
quantitative studies will be more valuable if combined with the re-
cognition of multiple fabrics.
4.2.5. Deformation structures
It is our interpretation that the 'deformation structures' listed above
forminhydrogranular “submagmatic”musheseitheratfaststrainrates
or as the magma is approaching its solidus and thus represent an im-
portant record of the rheological behavior of magma chambers under
these conditions.
Magmatic fold morphology depends on the rheologic state of ma-
terial being folded, particularly whether or not an effective viscosity
contract exists between different folded phases (Fernandez and
Gasquet, 1994), and the relative crystal/melt proportion. Folding in
crystal poor magmas may be more akin to soft sediment folding in
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
143 222
whichafluidphaseispresent.Magmaticfoldingcanobviouslyoccurat
very fast strain rates (see volcanic examples and Albertz et al., 2005).
Examplesbelowexemplifythepotentialofdetailedstudiesofmagmatic
folding. Vernon (1987) described magmatic folds in glassy rhyolitic
obsidian in which flow banding displays limb thinning (relative to
hinges) with crystallites and stretched vesicles defining an axial planar
foliation,reflectingstrainduringfolding.However,thesenseofrelative
rotation of the phenocrysts (Fig. 5p) remains constant around the folds
indicating that they record a pre-folding shear strain. Paterson et al.
(1998) describe a magmatic fold in an∼90 Ma pluton in the Cascades
core, Washington, where a magmatic foliation, defined by plagioclase
and hornblende, is folded and a new magmatic foliation defined by
poikilitic hornblende grew parallel to the axial plane (Fig. 2k). The
number of overgrown grains define a minimum crystal percent during
folding of ca 70%. Žák et al. (2007) describe a magmatically folded
leucogranitedikethathasparasiticfoldsinbothlimbsandhinge,hinge
thickening,andanaxialplanarmagmaticfoliation(Fig.2l).Linelength
method of 'unfolding' the dike indicates∼28% shortening.
We are not aware of any systematic study of different types of
magmatic faults and shear zones in plutons, indicating that this is an
importanttopicforfutureresearch.Publisheddescriptionsofindividual
fault types suggest that there are three main groups, namely those
caused by: (1) differential flow/strain gradients during local or
chamber wide convection (e.g., Hibbard and Watters, 1985; Neufeld
and Wettlaufer, 2008; Pinotti et al., 2016; Bergantz et al., 2017); (2)
marginal processes in growing crystal mush zones, particularly re-
sulting in normal faults during gravitational collapse of mushes
(Fig. 2q; e.g., Irvine et al., 1998; Solgadi and Sawyer, 2008; Paterson
et al., 2008) and cooling induced fractures (Geshi, 2001; John and
Stünitz,1997; Žáketal.,2009;HumphreysandHolness,2010);and(3)
synchronous tectonism during magma chamber crystallization (e.g.,
Blumenfeld and Bouchez, 1988; John and Blundy, 1993; Zibra, 2012).
Studies of magmatic faults are an underutilized tool to evaluate late
strainfieldsinplutons,thecollapseofcrystalrichmarginsandlocalized
convection in magma chambers, which are otherwise resisting further
flow or strain.
4.3. Modern approaches for studying magmatic structures
A combination of the following techniques provide a powerful
means to address the key issues outlined above: (1) GIS-based, geor-
eferenced grid-mapping provides key spatial information about the
type, orientation, and distribution of magmatic structures within plu-
tons (e.g., Olivier et al., 1997; Launeau and Cruden, 1998; Peternell
et al., 2011; Rocher et al., in press): (2) crystal size distribution data
provide a quantitative measure of cooling and crystallization histories
of plutonic bodies (e.g., Cashman and Marsh, 1988; Zieg and Marsh,
2002; Jerram et al., 2003): (3) Electron Back-Scatter Diffraction pro-
vides mineral lattice preferred orientations and when combined with
qualitative mineral composition is useful for evaluating deformation
mechanismsandkinematicinformation(Žáketal.,2008;Holnessetal.,
2017); (4) anisotropy of magnetic susceptibility is widely used to ana-
lyze pluton fabrics (e.g., Tarling and Hrouda, 1993; Archanjo et al.,
1995; Bouchez, 1997; Borradaile and Henry, 1997; Kruckenberg et al.,
2010; Dietl et al., 2010; Tomek et al., 2017; Žák et al., 2017); (5) mi-
neral-scale electron microprobe and laser ablation-inductively coupled
plasmamassspectrometryanalyses,leadingtotheapplicationofcrystal
thermometers and chemometers are powerful tools to evaluate the
origin, timing and rheologic evolution of compositionally defined
magmatic structures (McBirney, 1993; Solgadi and Sawyer, 2008;
Davidson et al., 2008; Barnes et al., 2016, 2017; Putirka, 2016; Zhang
et al., 2017); and (6) discrete element, computational fluid dynamics,
numericalsimulationsofahydrogranularmagmamushes(e.g.Bergantz
et al. (2017).
5. Conclusions
(1) A complex array of magmatic structures occur in plutons and pro-
vide useful tools for constraining hypersoidus temporal histories,
evolvingrheology,strainfieldsandinferredflowdirections,growth
and cooling information, local and regional tilting, and syn-em-
placement tectonism.
(2) Theabovestructuralhistoriesprovideapowerfulmeansoftestinga
wide array of incremental emplacement, magma chamber growth
and evolution and syn-emplacement tectonic models.
(3) Manyofthesemagmaticstructuresformin “hydrogranular”magma
slurries during local convection, collapse of mushes and strain re-
sulting in local structural diversity. Thus, the notion that magmas
must have≤55% crystals to convect, mix, fractionate or strain to
form compositional and structural diversity at these crustal levels,
needs to be revised.
(4) Researchers using preferred orientations to indicate magma flow
must remember that these reflect strain rather than flow directly:
thelattermustbeinferredfromstructuralgeometriesandtemporal
and kinematic information.
(5) The presence of multiple fabrics in plutons requires caution in the
interpretation of Anisotropy of Magnetic Suceptibility and other
quantitativetoolsthatdefinepreferredorientationsatthesampling
site through use of a single ellipsoid.
(6) Future structural studies are particularly needed in plutons ex-
amining the distribution, styles and formation of compositional
definedstructures,magmaticfolding,shearzones,brittlefaults,and
multiple fabrics in hydrogranular slurries.
Acknowledgements
WethankMarianHolnessandJean-LucBouchezforhelpfulreviews
and Joao Hibbertt for editorial assistance. Scott Paterson and Katie
Ardill acknowledge National Science Foundation support through
grants EAR-0537892 and EAR-1624847. Paterson acknowledges dis-
cussionswithJornKruhlandGeorgeBergantzaboutmagmaticsystems.
Jiří Žák acknowledges financial support from the Czech Science
Foundation through Grant No. 16-11500S and from the Charles
University through Center for Geosphere Dynamics (UNCE/SCI/006)
and project PROGRES Q45.
References
Albertz,M.,Paterson,S.R.,Okaya,D.,2005.Faststrainratesduringplutonemplacement:
magmatically folded leucocratic dikes in aureoles of the Mount Stuart
Batholith,Washington, and the Tuolumne Intrusive Suite, California. Geol. Soc. Am.
Bull. 117, 450–465.
Alasino, P.H., Larrovere, M.A., Rocher, S., Dahlquist, J.A., Basei, M.A., Memeti, V.,
Paterson, S.R.,Galindo, C., Grande,M.M., Neto,M.D.C.C., 2017.Incremental growth
ofanuppercrustal,A-typepluton,Argentina:evidenceofare-usedmagmapathway.
Lithos 284, 347–366.
Archanjo,C.J.,Launeau,P.,Bouchez,J.L.,1995.Magneticfabricvs.magnetiteandbiotite
shape fabrics of the magnetite-bearing granite pluton of Gameleiras (Northeast
Brazil). Phys. Earth Planet. Inter. 89, 63–75.
Ardill, K.E., Paterson, S.R., Barnes, C., Stanback, J., Alasino, P., Werts, K., Memeti, V.,
Teruya,L.,King,J.,Crosbie,S.,2017.MagmaticstructuresintheTuolumneIntrusive
Complex record physical differentiation of magmas at the emplacement level of an
upper crustal batholith. Geol. Soc. Am. 49http://dx.doi.org/10.1130/abs/2017AM-
302303. Abstracts with Programs.
Arzi, A.A., 1978. Critical phenomena in the rheology of partially melted rocks.
Tectonophysics 44, 173–184.
Balk, R., 1937. Structural Behavior of Igneous Rocks, vol. 5. Geological Society of
America Memoir, pp. 177.
Barbey, P., 2009. Layering and schlieren in granitoids: a record of interactions between
magmaemplacement,crystallizationanddeformationingrowingplutons.Geol.Belg.
12, 109–133.
Barbey, P., Gasquet, D., Pin, C., Bourgeix, A.L., 2008. Igneous banding, schlieren and
mafic enclaves in calc-alkaline granites: the Budduso pluton (Sardinia). Lithos 104,
147–163.
Barnes, C.G., Memeti, V., Coint, N., 2016. Deciphering magmatic processes in calc-alka-
line plutons using trace element zoning in hornblende. Am. Mineral. 101, 328–342.
Barnes, C.G., Berry, R., Barnes, M.A., Ernst, W.G., 2017. Trace element zoning in
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
144 223
hornblende: tracking and modeling the crystallization of a calc-alkaline arc pluton.
Am. Mineral. 102, 2390–2405.
Barrière, M., 1981. On curved laminae, graded layers, convection currents and dynamic
crystal sorting in the Ploumanac'h (Brittany) subalkaline granite. Contrib. Mineral.
Petrol. 77, 214–224.
Benn, K., 1994. Overprinting of magnetic fabrics in granites by small strains: numerical
modeling. Tectonophysics 233, 153–162.
Bergantz,G.W.,2000.Onthedynamicsofmagmamixingbyreintrusion:implicationsfor
pluton assembly processes. J. Struct. Geol. 22, 1297–1309.
Bergantz, G.W., Schleicher, J.M., Burgisser, A., 2017. On the kinematics and dynamics of
crystal-rich systems. J. Geophys. Res. Solid Earth 122.
Blumenfeld, P., Bouchez, J.L., 1988. Shear criteria in granite and migmatite deformed in
the magmatic and solid states. J. Struct. Geol. 10, 361–372.
Borradaile, G.J., Henry, B., 1997. Tectonic applications of magnetic susceptibility and its
anisotropy. Earth Sci. Rev. 42, 49–93.
Bouchez, J.L.,1997.Granite isneverisotropic:anintroduction toAMSstudiesofgranitic
rocks. In: Bouchez, J.L., Hutton, D.H.W., Stephens, W.E. (Eds.), Granite: from
Segregation of Melt to Emplacement Fabrics. Kluwer Academic Publishers,
Dordrecht, pp. 95–112.
Bouchez, J.L., Delas, C., Gleizes, G., Nédelec, A., Cuney, M., 1992. Submagmatic micro-
fractures in granites. Geology 20, 35–38.
Boudreau, A.E., 1995. Crystal aging and the formation offine-scale igneous layering.
Contrib. Mineral. Petrol. 54, 55–69.
Brown, M., 2013. Granite: from genesis to emplacement. Bull. Geol. Soc. Am. 125,
1079–1113.
Brown,P.E.,Chambers,A.D.,Becker,S.M.,1986.Alargesoft-sedimentfoldintheLilloise
Intrusion,EastGreenland.In:In:Parsons,I.(Ed.),OriginsofIgneousLayering,NATO
ASI Series C: Mathematical and Physical Sciences, vol. 196. D. Reidel Publishing
Company, Dordrecht, pp. 125–144.
Burgess, S.D., Miller, J.S., 2008. Construction, solidification and internal differentiation
of a large felsic arc pluton: cathedral Peak granodiorite, Sierra Nevada Batholith.
Geol. Soc. Lond. Spec. Publ. 304, 203–233.
Cao,W.,Paterson,S.,Memeti,V.,Mundil,R.,Anderson,J.L.,Schmidt,K.,2015.Tracking
paleodeformationfields in the Mesozoic central Sierra Nevada arc: implications for
intra-arc cyclic deformation and arc tempos. Lithosphere 7, 296–320.
Cashman,K.V.,Marsh,B.D.,1988.Crystalsizedistribution(CSD)inrocksandthekinetics
and dynamics ofcrystallization II.Makaopuhilava lake.Contrib. Mineral.Petrol. 99,
292–305.
Cashman, K.V., Sparks, R.S.J., Blundy, J.D., 2017. Vertically extensive and unstable
magmatic systems: a unified view of igneous processes. Science 355, eaag3055.
Castro, A., Martino, R., Vujovich, G., Otamendi, J., Pinotti, L., D'Eramo, F., Tibaldi, A.,
Viñao, A., 2008. Top-down structures of mafic enclaves within the Valle Fértil
magmatic complex (early Ordovician, san juan, Argentina). Geol. Acta 6, 217–229.
Castro, J., Manga, M., Cashman, K.V., 2002. Dynamics of obsidianflows inferred from
microstructures:insightsfrommicrolitepreferredorientations.EarthPlanet.Sci.Lett.
199, 211–226.
Clarke, D.B., Clarke, G.K.C., 1998. Layered granodiorites at chebutco head, South
mountain batholith, Nova Scotia. J. Struct. Geol. 20, 1305–1324.
Clarke, D.B., Grujic, D., McCuish, K.L., Sykes, J.C.P., Tweedale, F.M., 2013. Ring
schlieren:descriptionandinterpretationoffieldrelationsintheHalifaxpluton,South
mountain batholith, Nova Scotia. J. Struct. Geol. 51, 193–205.
Cloos,E.,1936.DerSierraNevadaPlutoninCalifornien.NeuesJahrbuchfürMineralogie.
Geol. Paleontol. 76, 355–450.
Cloos, E., 1946. Lineation: a critical review and annotated bibliography. Geol. Soc. Am.
Mem. 18. http://dx.doi.org/10.1130/MEM18.
Cloos,H.,1925.EinführungindietektonischeBehandlungmagmatischerErscheinungen:
Das Riesengebirge in Schlesien. Borntraeger, Berlin.
Coleman,D.S.,Gray,W.,Glazner,A.F.,2004.Rethinking theemplacementandevolution
ofzonedplutons:geochronologicevidenceforincrementalassemblyoftheTuolumne
Intrusive Suite, California. Geology 32, 433–436.
Coleman, D.S., Bartley, J.M., Glazner, A.F., Pardue, M.J., 2012. Is chemical zonation in
plutonic rocks driven by changes in source magma composition or shallow-crustal
differentiation? Geosphere 8, 1568–1587.
Compton, R., 1955. Trondhjemite batholith near bidwell bar, California. Geol. Soc. Am.
Bull. 66, 9–44.
Correa-Gomes, L.C., Souza Filho, C.R., Martins, C.J.F.N., Oliveira, E.P., 2001.
Development of symmetrical and asymmetrical fabrics in sheet-like igneous bodies:
theroleofmagmaflow andwall-rock displacements intheoretical andnatural cases.
J. Struct. Geol. 23, 1415–1428.
Costa,A.,Caricchi,L.,Bagdassarov,N.,2009.Amodelfortherheologyofparticlebearing
suspensions andpartially moltenrocks.Geochem. Geophys.Geosys. 10http://dx.doi.
org/10.1029/2008GC002138. Q03010.
Cruden,A.,1990.Flowandfabricdevelopmentduringthediapiricriseofmagma.J.Geol.
98, 681–698.
Davidson, J.P., Font, L., Charlier, B.L.A., Tepley, F.J., 2008. Mineral-scale Sr isotope
variation in plutonic rocks: a tool for unraveling the evolution of magma systems.
Earth Environ. Sci. Trans. R. Soc. Edinb. 97, 35–67.
Didier, J., 1973. Granites and Their Enclaves: the Bearing of Enclaves on the Origin of
Granites. Elsevier, Amsterdam.
Dietl, C., de Wall, H., Finger, F., 2010. Tube-like schlieren structures in the Fürstenstein
Intrusive Complex (Bavarian Forest, Germany): evidence for melt segregation and
magmaflow at intraplutonic contacts. Lithos 16, 321–339.
Dingwell, D.B., 1997. The brittle–ductile transition in high-level granitic magmas: ma-
terial constraints. J. Petrol. 38, 1635–1644.
Dufek, J., Bergantz, G.W., 2007. Suspended load and bed-load transport of particle-laden
gravity currents: the role of particle–bed interaction. Theor. Comput. Fluid Dyn. 21,
119–145.
Emeleus, C.H., Troll, V., 2014. The rum igneous centre, Scotland. Mineral. Magaz. 78.
http://dx.doi.org/10.1180/minmag.2014.078.4.04.
Fernandez, A., 1988. Strain Analysis from Shape Preferred Orientation in Magmatic
Flows, vol. 14. Bulletin of the Geological Institute, University of Uppsala, pp. 61–67.
Fernandez, A.N., Gasquet, D.R., 1994. Relative rheological evolution of chemically con-
trasted coeval magmas: example of the Tichka plutonic complex (Morocco). Contrib.
Mineral. Petrol. 116, 316–326.
Fletcher, R.C., 1982. Analysis of theflow in layeredfluids at small, butfinite, amplitude
with application to mullion structures. Tectonophysics 81, 51–66.
Gaschnig, R.M., Vervoort, J.D., Tikoff, B., Lewis, R.S., 2017. Construction and preserva-
tion of batholiths in the northern US Cordillera. Lithosphere 9, 315–324.
Geoffroy,L.,Callot,J.P.,Aubourg,C.,Moreira,M.,2002.Magneticandplagioclaselinear
fabric discrepancy in dykes: a new way to define theflow vector using magnetic
foliation. Terra Nova. 14, 183–190.
Geshi, N., 2001. Melt segregation by localized shear deformation and fracturing during
crystallisationofmagmainshallowintrusionsoftheOtogevolcaniccomplex,central
Japan. J. Volcanol. Geotherm. Res. 106, 285–300.
Gilbert, G.K., 1906. Gravitational assemblage in granite. Bull. Geol. Soc. Am. 17,
321–328.
Griffiths, R.W., 1986. Thermals in extremely viscousfluids, including the effects of
temperature-dependent viscosity. J. Fluid Mech. 166, 115–138.
Hardee, H.C., 1982. Incipient magma chamber formation as a result of repetitive intru-
sions. Bull. Volcanol. 45, 41–49.
Harm, S., 1991. Linear dilatation structures and syn-magmatic folding in granitoids. J.
Struct. Geol. 13, 625–634.
Hibbard, M.J., Watters, R.J., 1985. Fracturing and diking in incompletely crystallized
granitic plutons. Lithos 18, 1–12.
Hines, R., Paterson, S.R., Memeti, V., Chambers, J.A., 2018;al., in press. Nested incre-
mental growth of zoned upper crustal plutons in the Southern Uplands Terrane, UK:
fractionating, mixing,and contaminated magmafingers. J.Petrol. http://dx.doi.org/
10.1093/petrology/egy034. (in press).
Hodge, K.F., Carazzo, G., Montague, X., Jellinek, A.M., 2012. Magmatic structures in the
TuolumneIntrusiveSuite,California:anewmodelfortheformationanddeformation
of ladder dikes. Contrib. Mineral. Petrol. 164, 587–600.
Holness,M.B.,Vukmanovic,Z.,Mariani,E.,2017.Assessingtheroleofcompactioninthe
formation of adcumulates: a microstructural perspective. J. Petrol. 58, 643–673.
Humphreys, M.C.S., Holness, M.B., 2010. Melt-rich segregations in the Skaergaard
Marginal Border Series: tearing of a vertical silicate mush. Lithos 119, 181–192.
Hutton, D.H.W., 1992. Granite sheeted complexes: evidence for the dyking ascent me-
chanism. Trans. R. Soc. Edinb. Earth Sci. 83, 377–382.
Iezzi, G., Ventura, G., 2000. Kinematics of lavaflows based on fold analysis. Geophys.
Res. Lett. 27, 1227–1230.
Irvine, T.N., 1980. Magmatic density currents and cumulus processes. Am. J. Sci. 280,
1–58.
Irvine, T.N., 1987. Appendix I. Glossary of terms for layered intrusions. In: In: Parsons, I.
(Ed.), Origins of Igneous Layering, NATO ASI Series C: Mathematical and Physical
Sciences, vol. 196. D. Reidel Publishing Company, Dordrecht, pp. 641–647.
Irvine, T.N., Andersen, J.C.O., Brooks, C.K., 1998. Included blocks (and blocks within
blocks) in the Skaergaard intrusion: geologic relations and the origins of rhythmic
modally graded layers. Geol. Soc. Am. Bull. 110, 1398–1447.
Jackson, E.D., 1971. The origin of ultramfic rocks by cumulus processes. Fortschritte für
Mineral. 48, 8–74.
Jerram, D.A., Cheadle, M.J., Philpotts, A.R., 2003. Quantifying the building blocks of
igneous rocks: are clustered crystal frameworks the foundation? J. Petrol. 44,
2033–2052.
John, Barbara E., Blundy, Jonathan D., 1993. Emplacement-related deformation of
granitoid magmas, southern Adamello Massif, Italy. Geol. Soc. Am. Bull. 105 (12),
1517–1541.
John, B.E., Stünitz, H., 1997. Magmatic fracturing and small-scale melt segregation
during pluton emplacement: evidence from the Adamello massif (Italy). In: Bouchez,
J.L., Hutton, D., Stephens, W.E. (Eds.), Granite: from Segregation of Melt to
Emplacement Fabrics. Kluwer Academic Publishers, Dordrecht, pp. 55–74.
Johnson, S.E., Fletcher, J.M., Fanning, C.M., Vernon, R.H., Paterson, S.R., Tate, M.C.,
2003. Structure, emplacement and lateral expansion of the San Jose tonalite pluton,
PeninsularRangesbatholith,BajaCalifornia,Mexico.J.Struct.Geol.25,1933–1958.
Katz, R., Spiegelman, M., Holtzman, B.K., 2006. The dynamics of melt and shear locali-
zation in partially molten aggregates. Nature 442, 676–679.
King, B.C., 1964. The nature of basic igneous rocks and their relations with associated
acid rocks. Part IV. Sci. Prog. 52, 282–292.
Komar,P.D.,1972.Flowdifferentiation inigneousdikesandsills:profilesofvelocityand
phenocryst concentration. Geol. Soc. Am. Bull. 83, 3443–3448.
Kratinová, Z., Ježek, J., Schulmann, K., Hrouda, F., Shail, R.K., Lexa, O., 2010.
Noncoaxial K-feldspar and AMS subfabrics in the Land's End granite, Cornwall: evi-
dence of magmatic fabric decoupling during late deformation and matrix crystal-
lization. J. Geophys. Res. 115 B09104.
Kruckenberg, S.C., Ferré, E.C., Teyssier, C., Vanderhaeghe, O., Whitney, D.L., Seaton,
N.C.A., Skord, J.A., 2010. Viscoplasticflow in migmatites deduced from fabric ani-
sotropy: an example from the Naxos dome, Greece. J. Geophys. Res. 115, B09401.
Larsen, R., Brooks, C.K., 1994. Origin and evolution of gabbroic pegmatites in the
skaergaardintrusion,eastGreenland.J.Petrol.35,1651–1679.http://dx.doi.org/10.
1093/petrology/35.6.1651.
Launeau, P., Cruden, A., 1998. Magmatic fabric acquisition mechanisms in a syenite:
resultsofacombinedanisotropyofmagneticsusceptibilityandimageanalysisstudy.
J. Geophys. Res. 103, 5067–5089.
Lucus, S.B., St-Onge, M.R., 1995. Syn-tectonic magmatism and the development of
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
145 224
compositional layering, Ungawa Orogen (northern Quebec, Canada). J. Struct. Geol.
17, 475–491.
Lusk, A.D.J., Paterson, S.R., Ratschbacher, B.C., Larrovere, M., Alasino, P., Memeti, V.,
Cawood, T.K., Hernandez, R., 2017. Deformation of the uppermost Famatinian
orogen: mapping and structural analysis from the Sierra de Navaráez and Sierra de
Las Planchadas, NW Argentina. Geol. Soc. Am. 49http://dx.doi.org/10.1130/abs/
2017AM-304220. Abstract with Programs.
Macchioli Grande, M., Alasino, P.H., Rocher, S., Larrovere, M.A., Dahlquist, J.A., 2015.
Asymmetric textural and structural patterns of a granitic body emplaced at shallow
levels: the La Chinchilla pluton, northwestern Argentina. J. S. Am. Earth Sci. 64,
58–68.
Marsh, B.D., 2015. Magmatism, Magma, and Magma Chambers. Elsevier.
Martin, D., Griffiths, R.W., Campbell, I.H., 1987. Compositional and thermal-convection
in magma chambers. Contrib. Mineral. Petrol. 96, 465–475.
McBirney, A.R., Noyes, R.M., 1979. Crystallization and layering of the Skaergaard in-
trusion. J. Petrol. 20, 487–554.
McBirney, A.R., 1993. Igneous Petrology, second ed. Jones and Bartlett Publishers,
Boston.
Memeti, V., Paterson, S., Putirka, K., 2014. Formation of the Sierra Nevada Batholith:
Magmatic and Tectonic Processes and Their Tempos, vol. 34. Geological Society of
America Field Guide, pp. 116.
Miller, R.B., Paterson, S.R., 1994. Transition from magmatic to high-temperature solid-
state deformation, MountStuart batholith,Washington. J.Struct. Geol.16,853–865.
Miller, R.B., Paterson, S.R., 2001. Construction of mid-crustal sheeted plutons: examples
from the North Cascades, Washington. Geol. Soc. Am. Bull. 113, 1423–1442.
Moore, J.G., Lockwood, J.P., 1973. Origin of comb layering and orbicular structure,
Sierra Nevada batholith, California. Geol. Soc. Am. Bull. 84, 1–20.
Naslund, H.R., McBirney, A.R., 1996. Mechanisms of formation of igneous layering. In:
Cawthorn, R.G. (Ed.), Layered Intrusions. Elsevier, Amsterdam, pp. 1–43.
Neufeld, J.A., Wettlaufer, J.S., 2008. An experimental study of shear-enhanced convec-
tion in a mushy layer. J. Fluid Mech. 612, 363–385.
Nicolas, A., Ildefonse, B., 1996. Flow mechanism and viscosity in basaltic magma
chambers. Geophys. Res. Lett. 23, 2013–2016.
Olivier, P., de Saint Blanquat, M., Gleizes, G., Leblanc, D., 1997. Homogeneity of granite
fabricsatthemetreanddecametrescale.In:Bouchez,J.L.,Hutton,D.H.W.,Stephens,
W.E. (Eds.), Granite: from Segregation of Melt to Emplacement Fabrics. Kluwer
Academic Publishers, Amsterdam, pp. 113–127.
Park, Y.,Means, D., 1996. Directobservation ofdeformation processes in crystal mushes.
J. Struct. Geol. 18, 847–858.
Parsons, I., 1986. Origins of Igneous Layering. NATO ASI Series, Series C: Mathematical
and Physical Sciences, V. 196. D. Reidel Publishing Company, Dordrecht, Holland.
Passchier, C.W., Trouw, R.A.J., 1996. Microtectonics. Springer, Berlin.
Paterson, S.R., 2009. Magmatic tubes, troughs, pipes, and diapirs: late-stage convective
instabilities resulting in compositional diversity and permeable networks in crystal-
rich magmas of the Tuolumne Batholith, Sierra Nevada, California. Geosphere 5,
496–527.
Paterson,S.R.,Vernon,R.H.,Tobisch,O.T.,1989.Areviewofcriteriaforidentificationof
magmatic and tectonic foliations in granitoids. J. Struct. Geol. 11, 349–363.
Paterson, S.R., T. Fowler, K., Schmidt, K., Yoshinobu, A., Yuan, S., 1998. Interpreting
magmatic fabric patterns in plutons. Lithos 44, 53–82.
Paterson, S.R., Pignotta, G., Vernon, R.H., 2004. The significance of microgranitoid en-
clave shapes and orientations. J. Struct. Geol. 26, 1465–1481.
Paterson, S.R., Vernon, R.H., Žák, J., 2005. Mechanical instabilities and physical accu-
mulation of K-feldspar megacrysts in granitic magma, Tuolumne Batholith,
California, USA. J. Virtual Explor. 18 Paper 01.
Paterson, S.R., Žák, J., Janoušek, V., 2008. Growth of complex magmatic zones during
recycling of older magmatic phases: the Sawmill Canyon area in the Tuolumne
Batholith, Sierra Nevada, California. J. Volcanol. Geotherm. Res. 177, 457–484.
Paterson, S., Okaya, D., Memeti, V., Economos, R., Miller, R., 2011. Magma addition and
flux calculations of incrementally constructed magma chambers in continental
margin arcs: combinedfield, geochronologic, and thermal modeling studies.
Geosphere 7, 1439–1468.
Paterson, S., Memeti, V., Mundil, R., Žák, J., 2016. Repeated, multiscale, magmatic
erosion and recycling in an upper-crustal pluton: implications for magma chamber
dynamics and magma volume estimates. Am. Mineral. 101, 2176–2198.
Perugini, D., Poli, D., 2005. Viscousfingering during replenishment of felsic magma
chambers by continuous inputs of mafic magmas:field evidence andfluid-mechanics
experiments. Geology 33, 5–8.
Peternell, M., Bitencourt, F.M., Kruhl, J.H., 2011. Combined quantification of anisotropy
and inhomogeneityofmagmatic rock fabrics inanoutcrop scaleanalysis recorded in
high resolution. J. Struct. Geol. 33, 609–623.
Petford, N., 2009. Which effective viscosity? Mineral. Mag. 73, 167–191.
Philpotts, A., Asher, P., 1994. Magmaticflow-direction indicators in a giant diabase
feeder dike, Connecticut. Geology 22, 363–366.
Philpotts, A.R., Dickson, L.D., 2000. The formation of plagioclase chains during con-
vective transfer in basaltic magma. Nature 406, 59–61.
Pinotti, L.P., D'Eramo, F.J., Weinberg, R.F., Demartis, M., Tubía, J.M., Coniglio, J.E.,
Radice,S.,Maffini,M.N.,Aragón,E.,2016.Contrastingmagmaticstructuresbetween
small plutons and batholiths emplaced at shallow crustal level (Sierras de Córdoba,
Argentina). J. Struct. Geol. 92, 46–58.
Pitcher, W.S., Berger, A.R., 1972. The Geology of Donegal: a Study of Granite
Emplacement and Unroofing. John Wiley, London.
Pitcher,W.S.,Atherton,M.P.,Cobbing,E.J.,Beckinsale,R.D.,1985.MagmatismataPlate
Edge: the Peruvian Andes. John Wiley, London.
Price,N.J.,Cosgrove,J.W.,1990.AnalysisofGeologicalStructures.CambridgeUniversity
Press.
Putirka, K., 2016. Amphibole thermometers and barometers for igneous systems and
some implications for eruption mechanisms of felsic magmas at arc volcanoes. Am.
Mineral. 101, 841–858.
Puziewicz, J., Wojewoda, J., 1984. Origin ofload structures and graded beddingsof dark
minerals in the Strzegom granites (SW Poland): a sedimentological and petrological
approach. Neues Jahrb. für Mineral. Monatsh. 8, 353–364.
Ramsay, J.G., 1989. Emplacement kinematics of a granite diapir: the Chindamora
Batholith, Zimbabwe. J. Struct. Geol. 11, 191–209.
Reid,J.,John,B.,Murray,D.P.,Hermes,O.D.,Steig,E.J.,1993.Fractionalcrystallization
in granites of the Sierra Nevada: how important is it? Geology 21, 587–590.
Rocher,S.,Alasino,P.H.,MacchioliGrande,M.,Larrovere,M.A.,Paterson,S.R.,2018;al.,
in press. K-feldspar megacryst accumulations formed by mechanical instabilities in
magmachambermargins,Ashapluton,NWArgentina.J.Struct.Geol.112,154–173.
http://dx.doi.org/10.1016/j.jsg.2018.04.017. ISSN 0191-8141.
Rosenberg, C., Berger, A., Schmid, S., 1995. Observations from thefloor of a grantoid
pluton: inferences on the driving force offinal emplacement. Geology 23, 443–446.
Rosenberg, C.L., Handy, M.R., 2005. Experimental deformation of partially melted
granite revisited: implications for the continental crust. J. Metamorph. Geol. 23,
19–28.
Scaillet, B.,Whittington, A., Martel,C., Pichavant, M., Holtz, F.,2000. Phaseequilibrium
constraints on the viscosity of silicic magmas 2: implications for mafic-silicic mixing
processes. Trans. R. Soc. Edinb. Earth Sci. 91, 61–72.
Schleicher, J.M., Bergantz, G.W., Breidenthal, R.E., Burgisser, A., 2016. Time scales of
crystal mixing in magma mushes. Geophys. Res. Lett. 43.
Schmalholz,S.M.,Schmid,D.W.,2012.Foldinginpower-lawviscousmulti-layers.Philos.
Trans. R. Soc. A 370, 1798–1826.
Schofield,N.,Stevenson,C.,Reston,T.,2010.Magmafingersandhostrockfluidizationin
the emplacement of sills. Geology 38, 63–66.
Schulmann, K., Ježek, J., Venera, Z., 1997. Perpendicular linear fabrics in granite: mar-
kers of combined simple shear and pure shearflows? In: Bouchez, J.L., Hutton, D.,
Stephens, W.E. (Eds.), Granite: from Segregation of Melt to Emplacement Fabrics.
Kluwer Academic Publishers, Dordrecht, pp. 159–176.
Seaman,S.J., 2000. Crystal clusters, feldspar glomerocrysts, andmagma envelopes in the
AtascosaLookoutlavaflow,southern Arizona,USA:recordersofmagmaticevents. J.
Petrol. 41, 693–716.
Shannon, J.R., Walker, B.M., Carten, R.B., Geraghty, E.P., 1982. Unidirectional solidifi-
cationtexturesandtheirsignificanceindeterminingrelativeagesofintrusionsatthe
Henderson Mine, Colorado. Geology 10, 293–297.
Shelley, D.M., 1985. Determining paleo-flow directions from groundmass fabrics in the
Lyttleton radial dykes, New Zealand. J. Volcanol. Geotherm. Res. 25, 69–79.
Smith, J.V., Houston, E., 1994. Folds produced by gravity spreading of a banded rhyolite
lavaflow. J. Volcanol. Geotherm. Res. 63, 89–94.
Solgadi, F.,Sawyer,E.W.,2008. Formationofigneous layeringingranodiorite bygravity
flow: afield, microstructure and geochemical study of the Tuolumne Intrusive Suite
at Sawmill Canyon, California. J. Petrol. 49, 2009–2042.
Stickel, J.J., Powell, R.L., 2005. Fluid mechanics and rheology of dense suspensions.
Annu. Rev. Fluid Mech. 37, 129–149.
Streck, M.J., 2008. Mineral textures and zoning as evidence for open system processes.
Rev. Mineral. Geochem. 69, 595–619.
Talbot, J.Y., Martelet, G., Courrioux, G., Chen, Y., Faure, M., 2004. Emplacement in an
extensional setting of the Mont Lozére–Borne granitic complex (SE France) inferred
from comprehensive AMS, structural and gravity studies. J. Struct. Geol. 26, 11–28.
Tarling, D., Hrouda, F., 1993. Magnetic Anisotropy of Rocks. Chapman and Hall.
Tobisch, O.T., McNulty, B.A., Vernon, R.H., 1997. Microgranitoid enclave swarms in
granitic plutons, central Sierra Nevada, California. Lithos 40, 321–339.
Tomek, F., Žák, J., Verner, K., Holub, F.V., Sláma, J., Paterson, S.R., Memeti, V., 2017.
Mineral fabrics in high-level intrusions recording crustal strain and volcano–tectonic
interactions: the Shellenbarger pluton, Sierra Nevada, California. J. Geol. Soc. Lond.
174, 193–208.
Urai, J.L., Spaeth, G., van der Zee, W., Hilger, C., 2001. Evolution of mullion (boudin)
structures in the Variscan of the Ardennes and Eifel. J. Virtual Explor. 3 Paper 1.
van der Molen, I., Paterson, M.S., 1979. Experimental deformation of partially melted
granite. Contrib. Mineral. Petrol. 70, 299–318.
Vance, J.A., 1969. On synneusis. Contrib. Mineral. Petrol. 24, 7–29.
Vaughan, A.P.M., Thistlewood, L., Millar, I.L., 1995. Small-scale convection at the in-
terface between stratified layers of mafic and silicic magma, Campbell Ridges, NW
Palmer Land, Antarctic Peninsula: syn-magmatic way-up criteria. J. Struct. Geol. 17,
1071–1075.
Ventura, G.,2004. The strainpathand kinematics oflava domes: anexample from Lipari
(Aeolian Islands, Southern Tyrrhenian Sea, Italy). J. Geophys. Res. 109 B01203.
Vernon, R.H., 1987. A microstructural indicator of shear sense in volcanic rocks and its
relationship to porphyroblast rotation in metamorphic rocks. J. Geol. 95, 127–133.
Vernon,R.H.,2000.Reviewofmicrostructuralevidenceofmagmaticandsolid-stateflow.
Electron. Geosci. 5, 1–23.
Vernon, R.H., 2004. A Practical Guide to Rock Microstructure. Cambridge University
Press.
Vernon, R.H., Paterson, S.R., 2008a. Mesoscopic structures resulting from crystal accu-
mulation and melt movement in granites. Trans. R. Soc. Edinb. Earth Sci. 97,
369–381.
Vernon,R.H., Paterson, S.R.,2008b.Howextensiveare subsolidus grainshape changes in
cooling granites? Lithos 105, 42–50.
Vigneresse, J.L., Bouchez, J.L., 1997. Successive granitic magma batches during pluton
emplacement: the case study of Cabeza de Araya, Spain. J. Petrol. 38, 1767–1776.
Vigneresse, J.L., Barbey, P., Cuney, M., 1996. Rheological transitions during partial
meltingandcrystallizationwithapplicationtofelsicmagmasegregationandtransfer.
J. Petrol. 37, 1579–1600.
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
146 225
Vigneresse, J.L., 2008. Granitic batholiths: from pervasive and continuous melting in the
lower crust to discontinuous and spaced plutonism in the upper crust. Trans. R. Soc.
Edinb. Earth Sci. 97, 311–324.
Wager, L.R., Brown, G.M., 1968. Layered Igneous Rocks. Oliver and Boyd, Edinburgh.
Wahrhaftig, C., 1979. Significance of asymmetric schlieren for crystallization of granites
in the Sierra Nevada batholith, California. Geol. Soc. Am. 11, 133 Abstracts with
Programs.
Weinberg, R.F., Sial, R.N., Pessoa, R.R., 2001. Magmaflow within the Tavares pluton,
northwesternBrazil:compositionalandthermalconvection.Geol.Soc.Am.Bull.113,
508–520.
Wiebe, R.A., 1974. Coexisting intermediate and basic magmas, Ingonish, cape Breton
Island. J. Geol. 82, 74–87.
Wiebe,R.A.,Collins,W.,1998.Depositionalfeaturesandstratigraphicsectionsingranitic
plutons: implications for the emplacement and crystallization of granitic magma. J.
Struct. Geol. 20, 1273–1289.
Wiebe,R.A.,Jellinek,M.,Markley,M.J.,Hawkins,D.P.,Snyder,D.,2007.Steepschlieren
and associated enclaves in the Vinalhaven granite, Maine: possible indicators for
granite rheology. Contrib. Mineral. Petrol. 153, 121–138.
Wiebe, R.A., Jellinek, A.M., Hodge, K.F., 2017. New insights into the origin of ladder
dikes: implications for punctuated growth and crystal accumulation in the Cathedral
Peak granodiorite. Lithos 277, 241–258.
Yoshinobu, A.S., Wolak, J.M., Paterson, S.R., Pignotta, G.S., Anderson, H.S., 2009.
Determiningrelativemagmaandhostrockxenolithrheologyduringmagmaticfabric
formation in plutons: examples from the middle and upper crust. Geosphere 5,
270–285.
Žák, J., Klomínský, J., 2007. Magmatic structures in the Krkonoše–Jizera Plutonic
Complex, Bohemian Massif: evidence for localized multiphaseflow and small-scale
thermal-mechanical instabilities in a granitic magma chamber. J. Volcanol.
Geotherm. Res. 164, 254–267.
Žák, J., Paterson, S.R., 2009. Magmatic erosion of the solidification front during re-
intrusion: the eastern margin of the Tuolumne batholith, Sierra Nevada, California.
Int. J. Earth Sci. (Geol Rundsch) 99 (4), 801–812. http://dx.doi.org/10.1007/
s00531-009-0423-7.
Žák, J., Paterson, S.R., Memeti, V., 2007. Four magmatic fabrics in the Tuolumne bath-
olith, central Sierra Nevada, California (USA): implications for interpreting fabric
patterns in plutons and evolution of magma chambers in the upper crust. Geol. Soc.
Am. Bull. 119, 184–201.
Žák, J., Verner, K., Týcová, P., 2008. Grain-scale processes in actively deforming magma
mushes: new insights from electron backscatter diffraction (EBSD) analysis of biotite
schlieren in the Jizera granite, Bohemian Massif. Lithos 106, 309–322.
Žák, J., Paterson, S.R., Janoušek, V., Kabele, P., 2009. The Mammoth Peak sheeted
complex,Tuolumnebatholith,SierraNevada,California:arecordofinitialgrowthor
late thermal contraction in a magma chamber? Contrib. Mineral. Petrol. 158,
447–470.
Žák, J., Verner, K., Tomek, F., Johnson, K., Schwartz, J.J., 2017. Magnetic fabrics of arc
plutons reveal a significant Late Jurassic to Early Cretaceous change in the relative
plate motions of the Pacific Ocean basin and North America. Geosphere 13, 11–21.
Zhang, J., Humphreys, M.C.S., Cooper, G.F., Davidson, J.P., Macpherson, C.G., 2017.
Magma mush chemistry at subduction zones, revealed by new melt major element
inversion from calcic amphiboles. Am. Mineral. 102, 1353–1367.
Zibra, I., 2012. Syndeformational granite crystallisation along the Mount Magnet
Greenstone Belt, Yilgarn Craton: evidence of large-scale magma-driven strain loca-
lisation during Neoarchean time. Aust. J. Earth Sci. 59, 793–806.
Zieg, M.J., Marsh, B.D., 2002. Crystal size distributions and scaling laws in the quanti-
fication of igneous textures. J. Petrol. 43, 85–101.
S.R. Paterson et al. Journal of Structural Geology 125 (2019) 134–147
147 226
1
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Magmatically folded and faulted schlieren zones formed
by magma avalanching in the Sonora Pass Intrusive Suite,
Sierra Nevada, California
P .H. Alasino
1,2,
*, K. Ardill
3,
*, J. Stanback
3,
*, S.R. Paterson
3,
*, C. Galindo
4,
*, and M. Leopold
5,
*
1
Centro Regional de Investigaciones Científicas y Transferencia Tecnológica de La Rioja (Prov. de La Rioja–UNLaR–SEGEMAR–UNCa–CONICET), Entre Ríos y Mendoza s/n, Anillaco 5301, Argentina
2
Instituto de Geología y Recursos Naturales, Centro de Investigación e Innovación Tecnológica, Universidad Nacional de La Rioja (INGeReN-CENIIT -UNLaR), Av. Gob. Vernet y Apóstol Felipe,
5300 La Rioja, Argentina
3
Department of Earth Sciences, University of Southern California, 3651 Trousdale Parkway, Los Angeles, California 90089-0740, USA
4
Departamento de Mineralogía y Petrología, Universidad Complutense de Madrid–Instituto de Geociencias (UCM-CSIC), 28040 Madrid, Spain
5
Department of Geology, School of Science, Math, and Engineering, San Juan College, 4601 College Boulevard, Farmington, New Mexico 87402, USA
■ ABSTRACT
The southwestern margin of the Late Cretaceous Sonora Pass Intrusive
Suite, northern Sierra Nevada, California (USA), preserves a densely populated
zone of magmatic structures that record dynamic magmatic layer forma-
tion and deformation (faulting and folding) within a solidifying upper-crustal
magma mush. This zone consists largely of coupled melanocratic (or schlieren)
and leucocratic bands hosted within the 95.6 ± 1.5 Ma Kinney Lakes granodio-
rite (Leopold, 2016), with orientations approximately parallel to the intrusive
margin and with inward younging directions. Schlieren consist of a high modal
abundance of medium-grained ferromagnesian minerals (hornblende + biotite),
zircon, sphene, apatite, opaque minerals, and minor plagioclase and interstitial
quartz. Leucocratic bands are dominated by coarse-grained feldspar + quartz
with minor ferromagnesian and accessory minerals. Whole-rock geochemical
and Sr and Nd isotopic data indicate that the schlieren are derived from the
Kinney Lakes granodiorite by effective mechanical separation of mafic min -
erals and accessory phases.
We interpret that the schlieren zone at the margin of the Kinney Lakes
granodiorite formed by large-scale collapse of crystal mush by “magma ava-
lanching,” facilitated by gravity, local convection, and possibly by host-rock
stoping at the margin. This process eroded a significant portion of the solid -
ifying margin of the chamber and resulted in the formation of magmatically
deformed layered structures, which experienced further mingling, re-intru-
sion, magmatic erosion, and recycling processes. We envisage that magma
avalanching of magma mushes in plutons can be achieved by any unstable
process (e.g., tectonic, fluid-assisted, stoping, or gravity-driven) in large, long-
lived magma-mush chambers.
■ INTRODUCTION
Magmas can experience multiple physical and chemical processes during
their residence time within the crust before eruption or solidification into
plutons. Particularly, in a thermally mature crust, magmas can be maintained
above their solidus for thousands to millions of years and form a complex
interconnected network of crystal-melt mixtures which are commonly not
in isotopic equilibrium (e.g., Broxton et al., 1989; Christensen et al., 1995;
Davidson et al., 2001, 2008; Cooper and Reid, 2003; Costa et al., 2003; Lowery
Claiborne et al., 2006; Barbey et al., 2008; Ramos and Reid, 2005; Wallace
and Bergantz, 2005; Kaiser et al., 2016; Alasino et al., 2017). This crustal-scale
maturity, achieved by episodic intrusions from the lower crust through time
(e.g., de Silva et al., 2006; Paterson et al., 2011; Karakas et al., 2017), allows
magma chambers to evolve to highly dynamic hydrogranular environments
(e.g., Bergantz et al., 2017), in which extensive mixing, mingling, fractionation,
recycling, and contamination can modify the original magma source signa-
tures. In this context, physical and chemical instabilities driven by internal
thermal, compositional, and rheological gradients can form structural and
compositional diversity in chambers by late, local movement of the crys-
tal-rich system (e.g., Weinberg et al., 2001; Paterson et al., 2005, 2016, 2019;
Žák and Klomínský, 2007; Vernon and Paterson, 2008; Ruprecht et al., 2008;
Bachmann and Bergantz, 2008; Paterson, 2009; Pinotti et al., 2016; Bergantz
et al., 2017; Rocher et al., 2018). Deciphering the timing and mechanisms of
magmatic structure formation, and how these structures interact with local and
regional strain fields, is essential to understand long-lived magmatic systems.
This relationship influences our interpretation of chamber construction and
evolution and reconstruction of tectonic records.
Mechanical instabilities can occur along internal contacts either between
separate magma batches (Bergantz, 2000; Žák and Paterson, 2005; Memeti
et al., 2010) or along solidification fronts within magma chambers (Marsh,
1996, 2006; Žák and Paterson, 2010; Rocher et al., 2018). These boundary-layer
GEOSPHERE
GEOSPHERE, v. 15, no. X
https://doi.org/10.1130/GES02070.1
13 figures; 4 tables
CORRESPONDENCE: palasino@conicet.gov.ar
CIT A TION: Alasino, P .H., Ardill, K., Stanback, J., Pater son, S.R., Galindo, C., and Leopold, M., 2019, Mag matically folded and faulted schlieren zones formed
by magma avalanching in the Sonora Pass Intrusive
Suite, Sierra Nevada, California: Geosphere, v. 15,
no. X, p. 1–26, https://doi.org/10.1130/GES02070.1.
Science Editor: Shanaka de Silva
Associate Editor: Alan Whittington
Received 28 September 2018
Revision received 30 April 2019
Accepted 3 July 2019
© 2019 The Authors
This paper is published under the terms of the
CC BY NC license.
*E-mail: palasino@conicet.gov.ar; kardill@usc.edu; jstanbac@usc.edu; paterson@usc.edu;
cgalindo@geo.ucm.es; monika.b.leopold@gmail.com
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
227
2
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
instabilities can take different forms depending on the location of the layer and
crystal growth in the chamber (e.g., Jaupart and Tait, 1995; Bergantz, 2000).
For example, along the horizontal roof of a chamber, magma convection can
produce chemical and physical instabilities that cause crystal-rich batches to
descend, or “drip” , from the roof (e.g., Bergantz and Ni, 1999; Rocher et al.,
2018), while at a vertical or sloping margin, convection would generate a
sidewall current (e.g., Jaupart and Tait, 1995; Žák and Paterson, 2010; Pater-
son et al., 2016, 2019). Marsh (1996, 2006, 2013, 2015) examined the growth of
solidification fronts in chambers where the cooling is faster at the roof than
at deeper levels along walls. He suggested that these fronts may detach from
the roof, generating a crystal-rich magma avalanche toward hotter and deeper
levels of the chamber (e.g., Žák and Paterson, 2010). Davis et al. (2007) used
the concept of transient particle pressure in magma slurries to argue that
magmas affected by earthquake activity could partially fluidize crystal mushes
in seconds, resulting in an instantaneous decrease in pressure in the melt
phase, in situ bubble formation, and thus large-scale destabilization of the
mush, similar to the deformation of wet sediments during seismic loading (e.g.,
Sumita and Manga, 2008). Petford (2009) suggested that excess pressure due
to new bubble formation provides a mechanism for promoting chamber-wide
instability through overpressurization of the magma.
A field-based example of solidification front instability is found in the
Late Cretaceous Sonora Pass Intrusive Suite, Sierra Nevada, California, USA
(e.g., Wahrhaftig, 1979; Macias, 1996, Leopold, 2016) (Fig. 1). The oldest
unit, the Kinney Lakes granodiorite (KLG), intruded the Bummers Flat gra-
nodiorite (BFG) at 95.6 ± 1.5 Ma (Leopold, 2016) and exhibits a ~2 km
2
area
of complex schlieren and a range of other magmatic structures along the
intrusive contact with the BFG (Fig. 2). This zone is characterized largely by
coupled melanocratic (schlieren) and leucocratic mineral layering but also
includes blocks of host-rock material, enclave swarms, and magmatic folds
and faults (e.g., Leopold, 2016). In this paper, we explore the structure and
composition of schlieren in the KLG in the Sonora Pass Intrusive Suite and
their significance at the margins of the intrusion. We present the structural
features of the study area, focusing on a comprehensive data set of field
observations, structural data, whole-rock element geochemistry, and Sr and
Nd isotope compositions of the schlieren troughs in the KLG. Our results
provide insight into solidification-front instabilities in active magma cham -
bers, where large-scale magma avalanching (10
2
–10
3
m wide), caused by
large-scale collapse of a crystal-rich magma mush, may be focused along
the margin of a magma body.
■ GEOLOGICAL SETTING
The Sierra Nevada continental arc formed across a transitional oceanic to
continental margin and was largely constructed during three Mesozoic arc
flareup periods in the Late Triassic, Jurassic, and Cretaceous (Armstrong and
Ward, 1993; DeCelles et al., 2009; Paterson and Ducea, 2015; Kirsch et al., 2016)
(Fig. 1A). The Late Cretaceous arc flareup was 10
2
–10
3
times more voluminous
than previous arc flareups (Paterson and Ducea, 2015) and emplaced calc-al -
kaline tonalite, granodiorite, and granite plutons, which form a large part of
the presently exposed Sierra Nevada batholith.
The Sonora Pass region of the Sierra Nevada batholith, between 38°N and
39°N latitude, was initially studied by Slemmons (1953) and was subsequently
mapped at a regional scale (Giusso, 1981; John, 1983; Huber, 1983; Wahrhaftig,
2000). The Sonora Pass Intrusive Suite (SPIS) is the northernmost concen-
trically zoned intrusive complex of the Cretaceous Sierra Nevada batholith
(Fig. 1A) (e.g., Bateman, 1992; Macias, 1996; Leopold, 2016). It is also the small-
est (~650 km
2
) of four Late Cretaceous nested intrusive complexes ~1000 km
2
that intrude along the axis of the Sierra Nevada batholith (Stern et al., 1981;
Kistler et al., 1986; Bateman, 1992) (Fig. 1A). The SPIS was emplaced during
the peak of the Cretaceous flareup (Paterson and Ducea, 2015; Cao et al., 2015)
at pressures between 2.0 and 3.5 ± 0.6 kbar (Macias, 1996).
The SPIS consists of two main intrusive units, the Kinney Lakes horn-
blende-biotite granodiorite (95.6 ± 1.5 Ma; Leopold, 2016) and the Topaz Lake
biotite granodiorite (90.1 ± 1.1 Ma; Leopold, 2016) (Fig. 1B). The SPIS intruded
into the Bummers Flat granodiorite (109.8 ± 1.2 Ma; Leopold, 2016) in the
west and Paleozoic metasediments in the east. The metasediments have
been interpreted to be remnants of the Snow Lake passive-margin block (e.g.,
Lahren and Schweickert, 1989; Memeti et al., 2010). To the south, the SPIS
intrudes Jurassic mafic to ultramafic plutonic rocks and Cretaceous metavol -
canic units (Lahren et al., 1990; John et al., 1994; Cao et al. 2015). The SPIS is
compositionally normally zoned, with a felsic, porphyritic core (Topaz Lake
granodiorite) intruding the less differentiated, equigranular Kinney Lakes
granodiorite (Macias, 1996; Leopold, 2016). Contacts between the two units
are generally subvertical (Macias, 1996; Leopold, 2016). Initial
87
Sr/
86
Sr
i
ratios
between 0.7055 and 0.7058 within the SPIS indicates that the two units are
isotopically similar (Macias, 1996).
The study area is located in the southwestern part of the SPIS (Fig. 1B)
at the intrusive margin between the KLG and BFG units. A region of concen-
trated magmatic structures and complex schlieren is exposed here, which
has been interpreted to have formed from late-stage magmatic flow (Leo -
pold, 2016).
■ METHODS
Forty-two (42) samples were collected for petrography. One granodiorite
sample of KLG, one leucogranite sample of KLG, three schlieren samples in the
KLG unit, and three samples of felsic layers found between sampled schlieren
(herein termed leucocratic bands; see Table 1) were selected for whole-rock
major- and trace-element analyses using inductively coupled plasma–mass
spectrometry (ICP-MS) at Activation Laboratories, Ancaster, Ontario, Canada,
under the “4LithoResearch” package, following the procedure described at
http://www.actlabs.com (Table 2). Leucocratic band and schlieren samples
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
228
3
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
4220000
0 10 km
STUDY AREA
KLG
95 ± 2 Ma
TLG
90 ± 1 Ma
SPIS
Kinney Lakes
granodiorite (KLG)
Topaz Lake
granodiorite (TLG)
109 ± 1 Ma
BFG
Bridgeport, CA
(32 km)
Sonora, CA
(77 km)
108
Sonora
Pass
(2922 m
elevation)
240000
280000
260000
4240000
Gabbro and diorite bodies
within the BFG
Metamorphic rocks,
Host rocks to the SPIS
Bummers Flat granodiorite
(BFG)
Cretaceous plutons,
4260000
LA
California
SF
Nevada
o
Pacific
200 km
Ocean
SP
T
JM
MW
Tertiary volcanic cover
undivided
undivided
Sierra Nevada
B
A
Figure 1. Simplified regional geological map
of the Sonora Pass Intrusive Suite. (A) Lo-
cation of the Sierra Nevada arc section in
dark gray, and the study area, labeled SP .
T—Tuolumne Intrusive Suite; JM—John
Muir Intrusive Suite; MW—Mount Whitney
Intrusive Suite; SF—San Francisco; LA—
Los Angeles. (B) Map of the main plutonic
units of the Sonora Pass Intrusive suite
(SPIS) and plutonic host rocks relevant to
this study. Box shows the location of this
study (see Fig. 2). Geologic map of the So-
nora Pass Intrusive Suite is simplified from
John (1983) and unpublished mapping (D.A.
John, 2015, personal commun.). U-Pb zir-
con ages shown are from Leopold (2016).
Coordinate system is Universal Transverse
Mercator zone 11S (in meters), North Amer-
ican Datum 1983. CA—California.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
229
4
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Figure 2. Geologic map of the southwest-
ern intrusive contact of the Kinney Lakes
granodiorite (KLG) with the Bummers Flat
granodiorite (BFG). Map shows represen-
tative structural data from both igneous
units. Multiple magmatic foliations are
recognized. Structural data in black corre-
spond to the regional magmatic foliations;
structural data in red correspond to the
orientation of the schlieren. Blue line de-
lineates the extent of the schlieren zone
found at the margin. Local younging direc-
tion determined from graded layers and
trough cut-offs is also shown; the arrow-
head indicates its sense. The strike/dip (in
degrees) used for the orientation of intru-
sive contact is according to the right-hand
rule. Stereonets (equal-area, lower-hemi-
sphere projections) show poles to planes
of magmatic structures, such as regional
magmatic foliations, schlieren orientation,
and magmatic folds and faults in the KLG.
Coordinate system is Universal T ransverse
Mercator zone 11S (in meters), North Amer-
ican Datum 1983.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
230
5
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
were paired, i.e., they were collected from the same locality and are in contact.
Two additional samples of KLG were sampled by J. Miller (personal commun.,
2018) and analyzed using X-ray fluorescence at Pomona College, Claremont,
California (see Lackey et al. [2012] for further details regarding X-ray fluor -
escence analysis), for whole-rock major elements, and for trace elements
in the same samples at the Washington State University GeoAnalytical Lab,
Pullman, Washington, USA, using ICP-MS following the procedure described
at ht tps://en vironm ent.wsu.edu/facilities/geoanalytical-lab/technical-notes /icp
-ms -method/. Additionally, six representative igneous samples of nearby units
from this work and that of J. Miller (personal commun., 2018) (two mafic sam-
ples and two granodiorite samples of the BFG, and two granodiorite samples
of the Topaz Lake granodiorite unit) were selected for major- and trace-ele-
ment analysis. Nineteen (19) samples of the SPIS from Macias (1996) are also
included for comparison.
Strontium (Sr) and neodymium (Nd) isotopic analyses of six representative
KLG samples (one granodiorite, one leucogranite, two felsic layers, and two
schlieren) were carried out at the Geochronology and Isotope Geochemistry
Centre of Complutense University, Madrid, Spain (Table 3). Whole-rock pow-
ders were decomposed in 4 ml HF and 2 ml HNO
3
in Teflon digestion bombs
heated for 48 h at 120 °C and finally in 6 M HCl. Samarium (Sm) and Nd were
determined by isotope dilution using spikes enriched in
149
Sm and
150
Nd. Ion
exchange techniques were used to separate the elements for isotopic analysis.
Sr and rare earth elements (REEs) were separated using Bio-Rad AG50 × 12
cation exchange resin. Sm and Nd were further separated from the REE group
using Bio-Beads coated with 10% HDEHP [Di-(2-ethylhexyl) phosphoric acid].
Isotopic analyses were made on an automated multicollector VG SECTOR 54
mass spectrometer. Analytical uncertainties are estimated to be 0.006% for
143
Nd/
144
Nd and 0.1% for
147
Sm/
144
Nd. Replicate analyses of the JNdi Nd-isotope
standard yielded an average
143
Nd/
144
Nd ratio of 0.512106 ± 0.000005 (2σ) with n
= 7 (accepted value 0.5121 15 ± 0.000002; T anaka et al., 2000), and the NBS-987
Sr-isotope standard yielded an average
87
Sr/
86
Sr ratio of 0.710248 ± 0.000019
(2σ) with n = 12. Errors are quoted throughout as two standard deviations
from measured or calculated values.
■ FIELD AND STRUCTURAL RELATIONSHIPS AT THE KINNEY
LAKES GRANODIORITE MARGIN
Intrusive Contact between the Kinney Lakes Granodiorite and
Bummers Flat Host Rock
The southwestern intrusive contact between the BFG and the KLG is
relatively straight in the northern part of the study area, with a general NNW-
SSE–trending strike, but is highly irregular (at the 100 m scale) in the south,
where the contact strikes east-west (Fig. 2). The contact between the KLG and
the BFG is highly discordant. It varies from moderately dipping (~60° NE) to
vertical along strike, truncating all structures in the BFG. Several pegmatites
and aplite dikes (up to 1 m wide) from the KLG intrude the BFG. The KLG also
formed a small cupola (0.1 km
2
) in the BFG (Fig. 2). Local (centimeter to deci-
meter scale) solid-state shears are found in the BFG close to the margins of
the KLG. No solid-state deformation was found in the KLG within the study
area, although it is reported further to the east, near the contact with metased-
imentary host rocks and older mafic plutons (Leopold, 2016).
A rich array of magmatic structures is recognized at the margin of the
KLG. These include: regional and local magmatic fabrics, schlieren and trough
structures, mafic magmatic enclaves, leucogranite and mafic dikes, magma
mingling zones, host-rock xenoliths (of BFG), cognate inclusions (of earlier,
recycled KLG), and magmatic faults and folds. The zone of highly concentrated
magmatic structures extends ~800 m from the BFG-KLG contact to the blue
dashed line in Figure 2 and is discussed in detail below.
Regional Magmatic Fabrics
The dominant regional magmatic fabric in the KLG is NW-SE–striking and
steeply dipping to subvertical (60°–90°). The average orientation in the study
area is 154/90 (strike/dip according to the right-hand rule) (Fig. 2; n = 130
measurements). It is defined by aligned hornblende, biotite, and plagioclase.
TABLE 1. SUMMARY OF PETROGRAPHIC OBSERVATIONS OF SOME REPRESENTATIVE
SAMPLES WITH GEOCHEMICAL DATA, SONORA PASS INTRUSIVE SUITE
Sample Rock type Mineralogy Observations
SWP312 Granodiorite Pl > Qtz > HblBt > Kfs > Spn > Opq > Ap > Zrn
SWP2782 Schlieren Hbl > Bt > Opq > Spn > Pl > Qtz > Ap > Zrn Kfeldspar is rare
SWP2901 Schlieren Hbl > Bt > Opq > Pl > Spn > Qtz > Ap > Zrn Finegrained quartz; Kfeldspar is rare
SWP3071 Schlieren Bt > Hbl > Spn > Opq > Pl > Qtz > Ap > Zrn Finegrained quartz; Kfeldspar is rare
SWP2783 Leucocratic band PlQtz > Kfs > HblBt > OpqSpn > Ap > Zrn
SWP2902 Leucocratic band Pl > Qtz > Kfs > Bt > Hbl > Spn > Opq > Ap > Zrn
SWP3072 Leucocratic band Pl > Qtz > Kfs > Bt > Hbl > Opq > Spn > Ap > Zrn Secondary muscovite
SWP311 Leucogranodiorite Pl > QtzKfs > Bt Hornblende and accessories are rare; secondary muscovite
Abbreviations: Pl—plagioclase; Qtz—quartz; Hbl—hornblende; Bt—biotite; Kfs—Kfeldspar; Spn—sphene; Opq—opaques; Ap—apatite; Zrn—zircon.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
231
6
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
TABLE 2. MAJOR AND TRACEELEMENT COMPOSITIONS OF THE IGNEOUS SAMPLES
Host rock Sonora Pass Intrusive Suite
Unit BFG BFG BFG BFG KLG KLG KLG KLG KLG KLG KLG KLG KLG KLG KLG
Sample SWP272* SWP2722* SWP2792* SICN3
†
SWP312* SICN17
†
SICN25
†
S146
§
S102
§
S723
§
S724
§
S522
§
S228
§
R529
§
S288
§
Major elements (wt%)
SiO
2
48.34 43.71 67.06 67.59 67.16 66.14 64.98 61.2 65.4 65.2 66.6 64 67.7 64.9 68
TiO
2
1.68 1.08 0.51 0.36 0.52 0.56 0.55 0.80 0.57 0.56 0.63 0.63 0.48 0.55 0.44
Al
2
O
3
17.96 14.03 15.51 14.66 15.81 15.49 15.66 17.2 15.6 15.7 15.7 16.3 15.1 15.8 15.2
FeO
t
10.82 13.71 4.81 3.01 3.26 4.31 4.26 6.25 4.46 4.99 4.22 5.37 3.99 4.51 3.90
MgO 7.43 12.99 1.54 0.96 1.14 1.64 1.53 2.59 1.40 1.94 1.54 2.19 1.46 1.78 1.38
MnO 0.17 0.18 0.09 0.06 0.06 0.07 0.06 0.09 0.07 0.08 0.06 0.09 0.06 0.07 0.07
CaO 9.14 9.03 3.72 2.75 3.60 4.06 3.91 5.55 3.99 4.14 3.95 4.56 3.54 4.11 3.29
Na
2
O 3.18 1.49 3.49 3.32 3.63 3.32 3.36 3.73 3.80 3.38 3.67 3.51 3.33 3.67 3.47
K
2
O 1.18 2.35 2.22 3.31 3.21 3.01 3.41 1.98 2.67 2.89 3.13 2.73 3.46 2.85 3.35
P
2
O
5
0.1 0.06 0.11 0.1 0.13 0.16 0.15 0.21 0.16 0.16 0.16 0.16 0.12 0.16 0.12
LOI (%) 1.48 1.95 0.85 — 0.46 — — 0.55 1.80 0.75 0.55 0.45 0.25 0.45 0.65
Trace elements (ppm)
Cs 5.3 20.5 2.1 3.9 1.6 5.6 4.2 — — — — — — — —
Rb 53 108 87 115 95 85.1 94.7 69.3 94.2 105 104 107 131 92.1 118
Sr 754 478 374 406 496 519 572 641 562 451 543 517 431 504 426
Ba 244 275 859 1462 674 938 1141 790 823 791 649 753 784 856 679
La 10.8 7.47 32.2 35.6 21.7 15.2 21.1 27.2 23.2 27.2 22.3 23.1 22.1 25.8 19.2
Ce 24.7 18 57.9 69.4 34.7 34.5 41.2 55.6 41.7 49.5 41.0 47.3 46.0 44.7 32.2
Pr 3.39 2.49 6.36 4.9 3.63 7 5.1 6.46 4.66 5.51 4.67 5.58 5.25 4.97 3.48
Nd 15.1 10.9 22.3 20.1 13.4 13.3 19 24.8 17.7 20.7 17.9 20.5 19.4 18.4 12.9
Sm 3.73 2.8 4.13 3.3 2.44 1.6 3.4 4.43 3.13 3.63 3.17 3.80 3.34 3.25 2.29
Eu 1.18 1.02 0.84 — 0.66 — — 1.15 0.82 0.87 0.85 0.95 0.83 0.87 0.62
Gd 3.52 2.77 3.28 — 1.79 — — 3.76 2.50 3.16 2.58 3.25 2.78 2.75 1.90
Tb 0.55 0.4 0.5 — 0.24 — — 0.51 0.32 0.44 0.33 0.47 0.38 0.36 0.25
Dy 3.17 2.36 2.88 — 1.31 — — 2.71 1.58 2.49 1.59 2.64 2.05 1.89 1.29
Ho 0.59 0.43 0.56 — 0.23 — — 0.50 0.28 0.49 0.29 0.52 0.39 0.35 0.24
Er 1.63 1.15 1.73 — 0.63 — — 1.31 0.76 1.36 0.74 1.43 1.08 0.92 0.64
Tm 0.21 0.15 0.26 — 0.09 — — 0.19 0.10 0.20 0.10 0.22 0.16 0.13 0.09
Yb 1.29 0.92 1.7 — 0.55 — — 1.15 0.64 1.30 0.65 1.36 1.04 0.82 0.58
Lu 0.18 0.13 0.26 — 0.08 — — 0.17 0.10 0.20 0.10 0.20 0.16 0.13 0.09
U 0.9 0.74 3.45 0.5 3.55 1.4 0.2 2.90 4.38 4.10 3.57 3.96 6.41 3.72 4.64
Th 2 0.94 15.8 11 .2 10.8 5.2 3.7 11 .3 11 .8 16.4 12.7 11 .2 15.2 17.8 23.0
Y 17.1 12.2 17.8 11 .5 6.7 11 .2 9.9 — — — — — — — —
Nb 3.6 1.2 7.9 9.5 2.2 8 7.4 10.8 7.93 10.4 8.72 10.2 10.7 8.96 8.68
Zr 72 45 153 134 99 122 101 150 124 146 105 133 112 117 108
Hf 1.9 1.3 3.9 4.6 2.6 3.2 2.8 — — — — — — — —
Ta 0.34 0.14 1.08 1.3 0.47 4.9 2.4 — — — — — — — —
Ga 17 15 17 16.3 18 18.1 17.4 — — — — — — — —
Sc 43 21 8 5.9 4 9.1 8.8 — — — — — — — —
Notes: Total iron expressed as FeO
total
. Dashes indicate no value given (not analyzed). BFG—Bummers Flat granodiorite; KLG—Kinney Lakes granodiorite; SCH—schlieren in KLG; LB—leucocratic band in KLG; Lg—
leucogranite in KLG; TLG—Topaz Lake granodiorite; LOI—loss on ignition.
*Samples from this work.
†
Samples from J. Miller, personal commun., 2018.
§
Samples from Macias (1996).
(continued)
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
232
7
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
TABLE 2. MAJOR AND TRACEELEMENT COMPOSITIONS OF THE IGNEOUS SAMPLES ( continued)
Sonora Pass Intrusive Suite
Unit KLG SCH SCH SCH LB LB LB Lg TLG TLG TLG TLG TLG TLG TLG
Sample R489
§
SWP2782* SWP2901* SWP3071* SWP2783* SWP2902* SWP3072* SWP311* SICN10
†
SICN12
†
S504
§
S684
§
S604
§
S246
§
S1034
§
Major elements (wt%)
SiO
2
65.7 54.25 53.95 61.51 65.91 69.46 67.3 73.49 68.67 68.32 67.6 67 67.7 68.1 68.5
TiO
2
0.65 1.62 1.79 1.21 0.37 0.55 0.49 0.13 0.46 0.45 0.44 0.46 0.41 0.41 0.32
Al
2
O
3
15.9 14.78 12.52 13.1 16.48 14.45 16.09 14.95 15.66 15.73 16 15.6 16 15.2 15.5
FeO
t
4.44 10.59 12.73 8.11 3.2 3.81 3.52 1.25 2.95 3.04 3.38 3.35 3.28 3.17 2.37
MgO 1.58 4.73 4.87 3.32 1.07 1.32 1.3 0.31 0.05 0.05 1.01 1.16 1.06 0.88 0.61
MnO 0.06 0.21 0.25 0.16 0.05 0.07 0.07 0.03 2.66 2.74 0.05 0.06 0.05 0.05 0.05
CaO 4.01 6.22 4.95 3.91 4.57 2.96 3.34 1.75 0.89 0.94 3.39 3.3 3.48 2.97 2.46
Na
2
O 3.78 2.96 2.15 2.71 3.9 2.97 3.74 3.10 3.15 2.91 4.24 4.08 4.27 4.05 4.07
K
2
O 2.91 2.78 3.7 3.95 2.47 4.14 3.32 4.59 3.87 3.79 2.91 3.08 2.9 3.28 4.05
P
2
O
5
0.18 0.42 0.43 0.26 0.13 0.12 0.11 0.03 3.68 4.07 0.16 0.14 0.15 0.15 0.13
LOI (%) 0.3 — — 0.73 0.48 0.48 0.79 0.39 0.15 0.15 1.05 0.5 0.5 0.8 0.5
Trace elements (ppm)
Cs — 2.3 4.2 3.3 1.3 2 2 2.3 6 7 — — — — —
Rb 101 102 154 136 53 106 89 127 122 134 133 115 117 132 149
Sr 563 368 212 222 579 357 357 249 563 578 615 557 614 507 566
Ba 827 621 488 438 916 671 423 415 744 1227 704 736 742 675 1133
La 24.6 55.6 129 60.4 16.5 37.6 33.4 17.3 24.2 16.1 25.5 18.8 20.9 28.2 32.2
Ce 44.8 142 227 115 34.3 62.4 52.6 21.4 49.9 35.1 47.2 33.6 38.9 48.4 51.8
Pr 5.05 17.4 22.4 12.4 3.92 6.04 5.13 1.66 5.9 7.1 5.21 3.78 4.36 5.30 5.57
Nd 19.4 67.4 79.9 43.6 14.6 20.6 16.8 4.73 16.5 13.1 18.6 13.8 15.6 19.2 19.8
Sm 3.44 13 13.9 8.28 2.61 3.52 2.93 0.7 3.5 3.1 3.12 2.42 2.67 3.13 3.10
Eu 0.93 2.68 2.59 1.63 0.73 0.82 0.76 0.26 — — 0.78 0.65 0.72 0.75 0.73
Gd 2.76 9.8 10.3 6.34 1.96 2.62 2.18 0.57 — — 2.39 1.87 2.08 2.45 2.25
Tb 0.34 1.37 1.44 0.89 0.3 0.36 0.31 0.09 — — 0.29 0.23 0.25 0.30 0.27
Dy 1.69 7.52 7.93 4.93 1.52 2.05 1.73 0.49 — — 1.44 1.12 1.26 1.49 1.33
Ho 0.30 1.39 1.46 0.93 0.28 0.39 0.32 0.1 — — 0.27 0.20 0.23 0.28 0.24
Er 0.77 3.71 4.21 2.53 0.82 1.12 0.91 0.29 — — 0.75 0.52 0.63 0.77 0.67
Tm 0.11 0.56 0.62 0.374 0.119 0.16 0.148 0.05 — — 0.11 0.07 0.09 0.11 0.10
Yb 0.65 3.47 4.02 2.51 0.78 1.09 1.01 0.37 — — 0.70 0.46 0.58 0.73 0.66
Lu 0.10 0.54 0.64 0.393 0.112 0.165 0.164 0.06 — — 0.10 0.07 0.09 0.11 0.10
U 3.97 5.38 8.75 8.39 1.6 3.99 7.37 3.44 4.3 6.6 7.27 4.53 3.66 5.40 4.05
Th 12.1 18.3 78.6 50.2 4.54 32.8 33 22.1 12.7 11 .9 13.42 10.72 10.39 7.76 14.89
Y — 43.2 46.4 28.1 8.9 12 10.2 3.1 8.5 8.5 — — — — —
Nb 9.26 20.1 26 16.1 2.7 5 4.6 <0.2 8.9 7.3 8.43 6.93 8.20 9.68 9.36
Zr 123 324 405 268 94 117 118 54 110 98.8 127 106 105 122 120
Hf — 7.8 10.4 7.1 2.3 2.9 3.1 1.6 3.5 2.8 — — — — —
Ta — 3 3.64 2.16 0.64 0.94 0.82 0.21 3.4 5.1 — — — — —
Ga — 22 23 19 17 15 17 13 20.1 20.3 — — — — —
Sc — 20 23 15 5 5 5 1 5.6 7.1 — — — — —
(continued)
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
233
8
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
A second magmatic fabric of equal intensity but unevenly distributed in the
study area, also defined by hornblende and biotite, is ENE striking and is
also steeply dipping to subvertical (average orientation 078/86; n = 82 mea-
surements; Fig. 2). Lineation plunges defined by hornblende are steep to
subvertical (Leopold, 2016). Relative timing relationships between the two
fabrics are unclear and, in many cases, contradictory. These two fabrics are
widespread across the KLG unit (Macias, 1996; Leopold, 2016).
Compositionally Defined Layering and Structures
Planar Schlieren and Trough Structures
The most common type of compositional layering observed within the KLG
consists of alternating schlieren and leucocratic bands (or felsic schlieren)
that form planar to trough-like geometries (Figs. 3A–3D). Layer thickness is
on the order of centimeters to meters. Schlieren consist of medium-grained
ferromagnesian minerals hornblende and biotite, zircon, sphene, apatite, and
opaque minerals, ± quartz and plagioclase (Fig. 3A; T able 1). Leucocratic bands
between schlieren contain abundant coarse-grained feldspar (plagioclase
dominates over microcline) + quartz with minor ferromagnesian and accessory
minerals (Table 1). Commonly, schlieren have a basal surface with a sharp
contact defined by a thin, densely packed layer of mafic minerals (<1 cm),
whereas the upper contact is commonly less well defined and gradational into
the leucocratic bands at the centimeter scale (Figs. 3A, 4, 5).
At the map scale, schlieren broadly strike margin parallel, with an approx-
imately NNW-SSE orientation in the north of the study area. However,
orientations are highly variable at the outcrop scale (Fig. 2). There is no map-
scale schlieren alignment in the southern part of the study area. Schlieren dips
range widely from 2° to 88° but are on average steeply dipping to subvertical
toward the SW and NE (Fig. 2). The steepest dips are mainly found near the
NNW-SSE–striking contact, while shallow dips are dominant near the east-
west–striking contact (Fig. 2).
Locally, alternating schlieren and leucocratic bands contain two mineral
orientations: the dominant fabric is defined by mafic minerals located in
the schlieren aligned parallel to the base of the layer, and the second, less
intense fabric of similar strike but with a steeper dip preferentially forms in the
leucocratic band. In some places, the NW-striking regional magmatic fabric
overprints the schlieren (Fig. 4). Magmatic mineral lineations in the schlieren,
defined by hornblende and sphene, are approximately downdip with a ten -
dency to plunge toward the northeast.
Local younging directions are determined from trough cross-cutting rela-
tionships, comparable with those found in sedimentary layered rocks (e.g.,
Weinberg et al., 2001; Paterson, 2009; Figs. 3B, 3D, 4), as well as mineral
size and density grading in planar schlieren (Figs. 3A, 4, 5). Together these
structures indicate younging to the northeast (mean vector 048), toward the
interior of the KLG (Figs. 2, 4, 5).
TABLE 2. MAJOR AND TRACEELEMENT COMPOSITIONS
OF THE IGNEOUS SAMPLES (continued)
Sonora Pass Intrusive Suite
Unit TLG TLG TLG TLG TLG
Sample S696
§
S1024
§
S1184
§
S1244
§
S489
§
Major elements (wt%)
SiO
2
68.2 71.5 71 70.9 68.5
TiO
2
0.49 0.25 0.26 0.31 0.36
Al
2
O
3
15.4 14.3 15.3 15 15.2
FeO
t
3.64 2.11 2.03 2.46 2.88
MgO 1.19 0.51 0.48 0.61 0.65
MnO 0.06 0.04 0.04 0.05 0.06
CaO 3.55 1.9 2.11 2.46 2.64
Na
2
O 3.96 4.01 4.01 4.13 4.33
K
2
O 2.58 3.56 3.99 3.26 3.22
P
2
O
5
0.17 0.09 0.09 0.13 0.14
LOI (%) 0.7 0.65 0.55 0.45 1
Trace elements (ppm)
Cs — — — — —
Rb 101 162 166 160 141
Sr 578 408 500 512 570
Ba 497 670 1149 125 675
La 26.6 25.8 24.0 — 29.2
Ce 45.0 44.3 42.3 — 49.6
Pr 4.95 4.84 4.60 — 5.28
Nd 18.7 16.2 16.4 — 18.3
Sm 3.05 2.61 2.67 — 2.95
Eu 0.76 0.58 0.63 — 0.71
Gd 2.45 1.96 1.90 — 2.26
Tb 0.30 0.25 0.23 — 0.27
Dy 1.47 1.25 1.13 — 1.34
Ho 0.27 0.24 0.21 — 0.25
Er 0.71 0.70 0.60 — 0.68
Tm 0.10 0.11 0.09 — 0.10
Yb 0.65 0.75 0.59 — 0.68
Lu 0.10 0.12 0.09 — 0.10
U 4.82 2.70 3.47 — 6.20
Th 14.9 17.8 18.9 — 18.9
Y — — — — —
Nb 8.53 9.69 8.67 — 9.41
Zr 122 113 103 — 133
Hf — — — — —
Ta — — — — —
Ga — — — — —
Sc — — — — —
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
234
9
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
A
B
C
D
TABLE 3. Rb-Sr AND Sm-Nd DATA FOR IGNEOUS SAMPLES FROM THE SONORA PASS INTRUSIVE SUITE
Sample Rb
(ppm)
Sr
(ppm)
87
Rb/
86
Sr
87
Sr/
86
Sr
87
Sr/
86
Sr
t
* Sm
(ppm)
Nd
(ppm)
147
Sm/
144
Nd
143
Nd/
144
Nd
143
Nd/
144
Nd
t
ε
Nd(t)
†
Kinney Lakes granodiorite
SWP312 95 496 0.550 0.70648 0.70573 2.44 13.4 0.1101 0.51245 0.51238 –2.60
Schlieren in Kinney Lakes granodiorite
SWP290-1 154 212 2.101 0.70851 0.70567 13.9 79.9 0.1052 0.51242 0.51235 –3.09
SWP307-1 136 222 1.772 0.70817 0.70578 8.28 43.6 0.1148 0.51242 0.51235 –3.14
Leucocratic band in Kinney Lakes granodiorite
SWP290-2 106 357 0.859 0.70696 0.70580 3.52 20.6 0.1033 0.51242 0.51236 –3.04
SWP307-2 89 357 0.721 0.70679 0.70582 2.93 16.8 0.1054 0.51242 0.51235 –3.11
Leucogranite in Kinney Lakes granodiorite
SWP311 127 249 1.475 0.70802 0.70603 0.70 4.73 0.0895 0.51243 0.51237 –2.84
Note: The decay constants used in the calculations are the values λ
87
Rb = 1.42 × 10
−11
and λ
147
Sm = 6.54 × 10
−12
yr
−1
recommended by the International Union of
Geological Sciences (IUGS) Subcommission on Geochronology (Steiger and Jäger, 1977).
*t—time used for the calculation of the isotopic initial ratios; t = 95 Ma.
†
Epsilon-Nd values were calculated relative to a present-day chondrite (
143
Nd/
144
Nd)
today
CHUR
= 0.512638; (
143
Sm/
144
Nd)
today
CHUR
= 0.1967, where CHUR is chondritic
uniform reservoir.
Figure 3. Field photos within the schlieren zone showing typical
texture, magma mingling, and typical truncation features with en-
clave swarms. (A) Close-up view of melanocratic and leucocratic
layers with a sharp contact at the base of the schlieren and east-
ward grading (right-hand side of the photo). (B) Schlieren layering,
in places graded, re-intruded by leucogranite and dikes of Kinney
Lakes granodiorite (KLG). The package is gently folded. Load-cast
features are shown by red arrows. The width of outcrop in the
photo is ~4 m. (C) Steep truncation surface of shallowly dipping
schlieren by an enclave-rich unit. Height of outcrop in photo is
~2 m. (D) Enclave-rich unit truncating, along a magmatic angular
unconformity, shallowly dipping troughs and planar schlieren
(some truncation surfaces are evident at the lower right of the
image) formed at the Bummers Flat granodiorite–KLG contact.
An enclave-rich zone below the schlieren package represents a
separate magma batch from the upper unit. Height of outcrop
in photo is ~3 m.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
235
10
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Schlieren region
Kinney Lakes granodiorite
Schlieren
1m
Truncation
Y
Y
Y
Gradation
61
81
54
N
Schlieren orientation 18
Younging direction
Y
Gradational schlieren
Leucogranite dike
Regional magmatic fabric 18
Schlieren package 2
(dashed where faint
or graded)
Fault
Fold axis
Comb layering
Leucogranite dike
Schlieren orientation
18
Mineral lineation
Regional magmatic fabric
Younging direction
Y
18
Kinney Lakes granodiorite
Schlieren package 1
1 m
Dike
Truncation
54
18
Graded schlieren
57
Y
Y
regional fabric
schlieren
fabric in
normal fault
64
N
Figure 4. Sloped outcrop facing north (dip toward viewer) of a package of truncated schlieren troughs indicating an orientation of younging toward the northeast in the Kinney Lakes granodiorite
margin. Schlieren dip increases toward the northeast. Coarse schlieren troughs show mineral size grading, indicating consistent younging toward the northeast. Some schlieren in the southwestern
corner of the photo are magmatically folded. The regional fabric is well-defined SE-NW striking (164/81, right-hand rule). At left of the photo, a leucogranite dike truncates the schlieren packages.
Figure 5. Sloped outcrop facing south (dip toward viewer) of sharply truncated schlieren packages in the Kinney Lakes granodiorite (KLG) margin. The dominant regional magmatic fabric is approx-
imately east-west (e.g., 270/64, right-hand rule). In some places, the regional fabric overprints the fabric in the schlieren packages. On the right-hand (western) side of the outcrop is a schlieren-rich
area. Note that the right package truncates the left package and is thus younger. The late dike also supports this because it originates in the right package, drained from KLG magma, and intrudes the
left. Layering strikes roughly north-south and is moderately to steeply dipping (351/54). T extural features include comb layering and gradational layering. Magmatic mineral lineation in the schlieren
plane defined by hornblende and sphene are approximately down dip, with a tendency toward the NE (54/027; plunge trend). Graded schlieren layers young eastwards in the right-hand-side (western)
package, toward the truncation boundary, and most layers have been magmatically folded (axial plane: 085/57), with some layers being folded by multiple events. The left-hand-side (eastern) package
has east-west–striking, shallowly dipping schlieren layers (060/18), which are gradually more diffuse eastwards. T op-to-the-west normal faulting is observed close to the moderately dipping trunca-
tion boundary. At the faults, magmatic fabric defined by hornblende and biotite is deflected into the fault plane. Schlieren grading suggests that schlieren layers young upwards (and to the south).
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
236
11
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Mafic Enclaves
Magmatic mafic enclaves are common within the schlieren zone. Enclaves
are up to 20 cm length in the long axis, are rarely contained in the basal sec-
tions of schlieren bands, and are more commonly found within meter-scale
enclave-rich zones, or enclave swarms (Figs. 3C, 3D, 6A, 6B). In many cases,
enclave swarms also contain crystal clots of hornblende and biotite (Fig. 6B).
Two types of enclaves are distinguished at the KLG margin: type 1 has an
aphanitic texture, while type 2 contains millimeter-sized feldspar phenocrysts,
forming a porphyritic texture. Enclaves locally have felsic rims, or schlieren
rims, but more commonly are in direct contact with the host magma.
Within enclave swarms, enclaves may have almost all geometries and sizes:
spherical to ellipsoidal, angular to rounded, and quadruple pronged (all seen at
a single outcrop: Figs. 3D, 6B). In most enclave swarms, the enclaves appear to
have no preferred orientation (Fig. 6B), with some exceptions (Fig. 6A). Where
found in isolation, enclaves commonly appear lenticular to elongate parallel
with the regional fabric (e.g., Fig. 7A; n = 17 enclave long-axis measurements).
Dikes and Small Granitic Bodies
Fine-grained leucogranite bodies and dikes intrude the KLG (Figs. 3B, 3C,
4, 5, 6C). They are dominated by quartz and feldspar, with approximate widths
between 10 cm and 5 m and lengths >5 m. Dikes are found in all orientations,
but a limited number of measurements show a weak preference for dikes
striking NW-SE (average 334/82; n = 10). Dikes are usually straight, but are
also found folded or faulted, particularly where associated with schlieren
layers and trough structures (Figs. 5, 7A). Mafic dikes are rare, and typically
disaggregated or associated with enclave swarms (Fig. 8). Pegmatite dikes are
observed in addition (see Vugs and Pegmatites section below).
Magma Mingling Zones
Mingling between schlieren and the non-layered granodiorite (KLG) occurs
throughout the margin zone. Schlieren are commonly locally re-intruded by
B
C D
A
Figure 6. Field photos within the schlieren zone
illustrating interactions of schlieren with enclave
swarms and stoped blocks. (A) Enclave swarm sur-
rounds block of Kinney Lakes granodiorite (KLG).
A meter-scale vug in the upper left of the image
records local boiling. Schlieren are observed in the
top right of the image. (B) Close-up view of enclave
swarm and cognate inclusion shown in A. Enclaves
are a wide range of shapes with no clear alignment,
include schlieren in the interstitial regions, and are
sharply truncated by the cognate inclusion. (C) Lay-
ered, stoped block in KLG (likely a cognate inclusion)
is cut by leuco granite veins, and normal faults indi-
cated by a red arrow are found in the schlieren layers
below the block. Scale bar is 15 cm. (D) Stoped
blocks of mafic Bummers Flat granodiorite in the
KLG with angular and subrounded edges. Largest
block in the photo is ~2 m wide.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
237
12
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
C
F
E
A
B
D
Figure 7. Field photos within the schlie-
ren zone illustrating the different types of
deformation-related structures observed.
(A) Folded schlieren layer and leucogran-
ite dike, where the regional NW-striking
fabric represents the axial planar cleav-
age (direction of pencil). Enclaves are also
aligned with the fabric. (B) Layered schlie-
ren package broken up by faulting. 19 cm
notebook for scale. (C) Schlieren package
displaced by right-lateral fault running
from left to right in the image (above pen-
cil). (D) Outcrop-scale magmatic folds of
schlieren layering with a local axial-planar
magmatic foliation. Local truncations are
visible within the folds, documenting mul-
tiple magmatic events. (E) Another view of
the outcrop shown in D showing folding
and magmatic faulting of schlieren layers.
(F) Outcrop-scale view of the schlieren zone
with folded and faulted layers.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
238
13
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
the KLG unit (Fig. 8). The mingling process is commonly associated with
some magmatic deformation at the outcrop scale, such as faulting and folding
(Figs. 3B, 8). At the scale of a single layer, contacts between schlieren and leuco-
cratic layers are in some places lobate. Locally, features at least superficially
similar to load casts are found at the base of schlieren (Fig. 3B). Enclaves also
in some cases have lobate boundaries with the host magma.
Xenolith Blocks and Cognate Inclusions
Isolated igneous blocks are ubiquitous near the KLG margin in the mapped
area (Figs. 2, 6A–6D). The most common types of blocks are xenoliths of the
BFG (sheeted granodiorite, tonalite, and diorite) and cognate inclusions of
the KLG. Blocks vary in size from several centimeters to as much as 300 m in
length, and the largest blocks found were of BFG xenoliths (Figs. 2, 6C, 6D).
At the map scale, xenoliths are irregular to rectangular in shape with either
sharp or subangular corners (Figs. 2, 6C). In outcrop-sized and smaller blocks,
margins of BFG xenoliths are commonly sharp and angular, but locally are
lobate (Fig. 6D) or, in rare cases, diffuse with the host KLG.
Cognate inclusions of mafic KLG are similar in shape and proportion to BFG
xenoliths, but generally with lobate and subrounded edges (Figs. 6A, 6B). They
have a similar texture to the KLG in the study area, but differences include a
finer grain size, higher content of mafic minerals (a mafic granodiorite), and
internal layered structures.
Both types of blocks are commonly associated with leucogranite dikes and
veins, where blocks are broken into smaller pieces and/or fractured (Figs. 6C,
6D). Where blocks are associated with schlieren, layer orientations are in some
cases arranged parallel to the edges of the largest blocks, with highly variable
dips (Fig. 2). The blocks appear to have little to no effect on the orientation of
the regional magmatic foliations.
Vugs and Pegmatites
A small number of miarolitic cavities (centimeter scale) and vugs (meter
scale) composed of pegmatitic quartz, feldspar, and minor biotite and tour-
maline are found in the study area (Fig. 6A). Vugs have spherical to ellipsoidal
geometries. Pegmatite dikes are more common, and sometimes are folded
(Fig. 7A).
Deformation-Related Structures
Magmatic Faults
Magmatic faults are observed at the KLG margin where they offset pla-
nar schlieren and troughs. Faults are mostly normal in the sense of slip, are
1m
Re-intruded mafic dike
Dikes
Kinney Lakes granodiorite
Schlieren
Leucogranite dike
Mafic enclave
Schlieren region
K-feldspar phenocrysts
Schlieren region outlined
Re-intruded magma
Mineral lineation
Regional magmatic fabric
Sharply truncated layers
Fault
Schlieren orientation 18
Younging direction
Y
Y
72
N
18
67
Figure 8. Horizontal outcrop viewed in the southeast direction of deformed schlieren
layers in the Kinney Lakes granodiorite (KLG) margin. The regional fabric is well-
defined striking north-south (018/67 , right-hand rule). Some portions of the KLG in
this region have alkali-feldspar clusters, which is atypical for the margins of this unit
(the core is more porphyritic). Schlieren dip steeply (357/72) roughly parallel to the
regional magmatic mineral foliation. Magmatic mineral lineation is roughly down
dip (62/085). Cross-cutting schlieren layers indicate a younging direction toward the
north. Schlieren are locally re-intruded by KLG, and nearby, schlieren are folded and
faulted magmatically (see red dashed line). The disaggregation of the mafic dike
shows magma mingling with the KLG. All structures, including the mafic dike, are
cut by late leucogranite veins.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
239
14
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
oriented at a high angle to the schlieren layering and have minimal displace-
ments <1 m (Figs. 6C, 7B, 7C, 7E, 8). A single magmatic fault may displace
multiple schlieren, and within a single outcrop there is evidence for several
distinct faulting events (Fig. 5, 7B).
Measurements on fault slip planes indicate an average sense of slip par-
allel to the contact margin (NNW-SSE trending) with a dip of ~75° and sense
of movement down to the east (see stereonet in Fig. 2; n = 4). In some cases,
magmatic normal faults offset leucogranite dikes that intrude the schlieren
zone (Fig. 5). Leucogranite veins are commonly found along the fault plane
(Figs. 7B, 7C, 7E). At the fault, the local magmatic fabric in schlieren, defined by
hornblende and biotite, was deflected into the fault plane in the magmatic state
(Fig. 5). No solid-state deformation is found along faults at the outcrop scale.
Magmatic Folds
Like magmatic faults, magmatic folds at the KLG margin are observed
deforming multiple schlieren and trough structures. Magmatic folds vary in
wavelength and amplitude over short distances (Figs. 3B, 5, 7A, 7D–7F). Locally,
they deform a small number of schlieren layers with thicknesses from top to
bottom <2 m. Some schlieren are folded by multiple events (Fig. 5). Folds are
distributed throughout the margin zone (Fig. 7F), commonly associated with
either magmatic faults (Figs. 5, 7E) or host-rock blocks. In folded schlieren,
the orientation of minerals is deflected to a new orientation parallel to the
axial plane (Fig. 5). Folds have multiple orientations of axial planes that may
be associated with the emplacement of host-rock blocks. A limited number
of measured fold axial planes excluding those associated with the host-rock
blocks show a weak preference for contact-parallel strikes and a restricted
span of dip values (~70°–90°) (see stereonet in Fig. 2; n = 7).
■ SPATIOTEMPORAL RELATIONSHIPS BETWEEN STRUCTURAL
ELEMENTS OF THE SCHLIEREN ZONE
The wide variation in magmatic structures preserved at the margin of the
KLG allows us to investigate the relative chronology of magmatic events. Below
we present evidence pertaining to relative timing relationships between dif-
ferent structures.
The formation of the schlieren, troughs, and enclave swarms were closely
intertwined in the study area. Schlieren and enclave swarms are commonly
interbedded. Enclave swarms in some cases truncate schlieren packages
(Fig. 3C) and vice versa (Fig. 3D). Trough truncations in addition indicate
repeated magmatic erosion and redeposition to form new schlieren (Figs. 3B, 4).
Schlieren structures (both planar schlieren and troughs) were subsequently
re-intruded by KLG (Fig. 8) and deformed by magmatic faults and folds (Figs. 5,
7C–7E). The mingling and re-intrusion of KLG into schlieren (Fig. 8) suggests
that the schlieren were in a mush state and could be partly reworked. However,
in some cases (e.g., Fig. 8) schlieren behaved as rigid blocks that were dis-
placed and recycled as one cohesive package into the KLG.
The NW-SE and ENE-WSW regional magmatic fabrics defined by horn -
blende and biotite are locally defined by stretched, aligned enclaves (Fig. 7A).
The NW-SE fabric is roughly parallel to the orientation of schlieren, as well
as the KLG-BFG contact in the north of the study area. At the map scale, the
NW-SE fabric is discordant with the southern contact of the KLG-BFG. Within
individual schlieren structures, the local schlieren fabric may be highly discor-
dant with both regional fabrics, and in some cases, the local schlieren fabric
is overprinted by the regional fabric(s). In folded schlieren, the regional fabric
may be deflected to a new orientation parallel to the axial plane (Figs. 5, 7D).
Stoped blocks of BFG and cognate inclusions of KLG in the study area
are typically angular, with a few subrounded shapes, indicating magmatic
interaction of the blocks with the KLG. Blocks in some cases truncate enclave
swarms (Fig. 6B) and have deformed schlieren at the map and outcrop scale
(Figs. 2, 6C), suggesting that they are closely linked in time with both the
formation and deformation of the compositionally defined structures. The
cognate inclusions of KLG, generally of a more mafic granodiorite composition,
also contain a magmatic fabric rotated (as much as 90°) from the strike of the
regional magmatic fabrics within the KLG, and were broken and fractured by
leucocratic veins and dikes. Indeed, blocks including dike fragments that are
in turn surrounded by schlieren indicate that the blocks may have experienced
multiple episodes of fracture and movement in the chamber (Fig. 3D).
Field evidence suggests that the leucogranite bodies and dikes were melts
draining from the KLG mush (Fig. 5). Dikes are considered in most cases as
late features due to their generally high silica compositions and their cross-cut-
ting relationships with stoped blocks (Fig. 6C), schlieren structures (Figs. 3B,
4, 5, 8), and magmatic faults (Fig. 5). Pegmatite dikes and vugs also fit into
this category. However, dikes are locally folded and in some cases included
as a component of blocks (Figs. 3D, 5), suggesting that they also experienced
magmatic deformation and formed at different times during the formation of
the KLG margin zone.
Deformation-related structures such as magmatic faults and folds deform
multiple, cohesive schlieren packages (Figs. 5, 7B–7E). Thus, they must slightly
postdate the formation of the layers. However, Figure 5 shows the complex-
ity in timing relationships, as a package of magmatically folded schlieren is
truncated by another package of folded and faulted schlieren. This indicates
that deformation was synchronous with the generation of planar schlieren and
troughs. In addition, folding and faulting were occurring at the same time (Fig. 5).
■ PETROGRAPHY AT THE KINNEY LAKES GRANODIORITE MARGIN
Kinney Lakes Granodiorite and Leucogranite
The KLG contains feldspars and quartz together with hornblende, bio-
tite, sphene, magnetite, and accessory apatite and zircon (Table 1). The
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
240
15
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
grain size of quartz and feldspars (~2–4 mm) in the host unit are consid-
erably larger than in both the schlieren and leucocratic bands. Evidence
for intracrystalline deformation is found only in quartz, which displays
undulose extinction.
The leucogranite within the KLG unit is composed of plagioclase, quartz,
alkali feldspar, and biotite (Table 1). Biotite typically has ragged grain bound-
aries at some ends (Fig. 9A). In rare cases, it shows undulose extinction where
it is pinned against larger feldspar grains. Plagioclase contains magmatic
microstructures such as euhedral shapes, concentric compositional zoning,
and growth twins (Fig. 9A). Secondary muscovite and sericite are found
in some plagioclase crystals. Quartz shows sweeping undulose extinction
(Fig. 9A).
Schlieren and Leucocratic Bands
Schlieren bands are dominated by euhedral hornblende, biotite, magne-
tite, and sphene, with accessory apatite and zircon (Table 1). Plagioclase and
quartz occupy the interstitial regions (~0.5–1 mm size), while alkali feldspar is
rare. Hornblende, biotite, and sphene are aligned, forming a magmatic fabric
in the schlieren, and do not show evidence for lattice distortion. In contrast
quartz, occupying interstitial pockets between phenocrysts, shows undulose
extinction of both sweeping and chessboard structures (Fig. 9B). Feldspars
record intracrystalline deformation in the form of rare deformation twinning,
but microstructures are largely magmatic, as growth twinning and zoning are
ubiquitous (Fig. 9C). Plagioclase crystal clots are observed, where crystals
bt
plag
kfsp
Pl
Qtz
Hbl kfs
Pl
Qtz
Pl
Pl
Qtz
Pl
Hbl
Spn
Pl
Bt
Kfs
Kfs
Qtz
Bt
kfs
A B
C D
Figure 9. Photomicrographs of leucogranite, schlie-
ren, and felsic layers at the Kinney Lakes granodiorite
margin, all under cross-polarized light. Scale bars are
0.5 mm. (A) Sample SWP311 of leucogranite showing
the ragged edges of biotite intergrown with plagioclase,
alkali feldspar, and quartz, interpreted to be a magmatic
growth feature. (B) Sample SWP 290-1 of a schlieren
layer showing chessboard undulose extinction in quartz,
with unmodified, magmatic margins and dihedral angles.
(C) Sample SWP 290-1 of a schlieren layer showing a clot
of plagioclase crystals. Note that the plagioclase twins
are largely growth twins, with a few exceptions that
are deformation twins. Plagioclase-plagioclase bound-
aries are low angle or parallel to growth twins, with few
boundaries at high angle to growth twinning. (D) Sample
SWP 290-2 felsic band in schlieren showing graphic mi-
crostructure between plagioclase and alkali feldspar, and
between alkali feldspar and quartz. This microstructure
appears to form preferentially within pools and along
grain boundaries. Pl—plagioclase; Qtz—quartz; Hbl—
hornblende; Bt—biotite; Kfs—K-feldspar; Spn—sphene.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
241
16
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
truncate growth zoning, indicative of synneusis or mechanical clumping. Qual-
itative observations of grain boundaries and dihedral angles are comparable
to those described by Holness et al. (2018) for plutonic units dominated by
magmatic crystallization (Fig. 9C). Rare quartz subgrains, fractured amphibole,
and glide twins in plagioclase are suggestive of some late crystal deformation
during settling and compaction in the presence of melt.
Leucocratic bands between schlieren are dominated by plagioclase and
alkali feldspar, with lesser amounts of euhedral hornblende, biotite, and mag-
netite (T able 1). Leucocratic bands display all of the microstructures contained
within schlieren, but at a lower intensity. For example, undulose extinction in
quartz is not as pervasive as in the schlieren. Graphic texture is common in
the leucocratic bands (Fig. 9D). In these samples, it forms from the intergrowth
between plagioclase and alkali feldspar and between quartz and alkali feld-
spar, occurring at and along grain boundaries and as pools between grains.
■ WHOLE-ROCK GEOCHEMISTRY
Southwestern Host Rock of the Sonora Pass Intrusive Suite
Two samples of the Bummers Flat granodiorite (BFG) host rock show con-
tents of SiO
2
~67%, TiO
2
~0.4%, FeO
total
of 3% and 4.8%, CaO ~3%, and alkalis
~6% with a Na
2
O/K
2
O ratio ≥1 (Table 2). In the alkalis versus silica classification
diagram, the compositions plot in the granodiorite field (Fig. 10A). The samples
have ASI (aluminum saturation index) values of ~1 .06 (i.e., slightly peraluminous),
and they plot in the medium- to high-K fields on a K
2
O versus SiO
2
diagram
(Fig. 10B). They show contents of Rb ≤ 115 ppm, Ba ≤1462 ppm, Sr ≤406 ppm,
Nb ≤9.5 ppm, Y ≤17 .8 ppm, Yb = 1.7 ppm (one value), and Zr ≤153 ppm (T able 2).
The chondrite-normalized (N) REE pattern of one sample shows La
N
/Yb
N
ratio of
~1 1 and a moderate Eu anomaly (Eu
N
/Eu
N
* = 0.7) (not shown).
Compared to the BFG, mafic intrusions show lower SiO
2
contents (≤48%)
with relatively high TiO
2
(≤1.7%), FeO
t
(≤13.7%), CaO (~9%), and alkalis (~4%)
with variable values of the Na
2
O/K
2
O ratio (Table 2). These two mafic samples
in the alkalis versus silica classification diagram plot in the gabbro and foid
gabbro fields (Fig. 10A). In a K
2
O versus SiO
2
diagram, the gabbro falls in the
medium-K field whereas the foid gabbro plots in the high-K field (Fig. 10B).
ASI values are ~0.7 , i.e., metaluminous. They show relatively low contents of
Rb ≤108 ppm, Ba ≤275 ppm, Nb ≤3.6 ppm, Y ≤17 .1 ppm, Yb ≤1.3 ppm, and Zr
≤72 ppm, and high Sr contents of ≤754 ppm (Table 2). REE patterns of the two
gabbroic samples are flat (La
N
/Yb
N
ratios ~5) with either no Eu anomaly or a
positive one (Eu
N
/Eu
N
* = 1 and 1.12) (not shown).
Sonora Pass Intrusive Suite
The older plutonic unit, the Kinney Lakes granodiorite (KLG), does not
show a large variation in major elements. It has restricted contents of SiO
2
(61%–68%), TiO
2
(0.4%–0.8%), FeO (3.3%–6.2%), CaO (3.3%–5.5%), and alkalis
(5.7%–6.8%), with a Na
2
O/K
2
O ratio usually >1 (T able 2). Most of the samples fall
in the granodiorite field in the alkalis versus silica classification diagram, and
in the high-K field on a K
2
O versus SiO
2
diagram (Figs. 10A, 10B). ASI values
range from 0.97 to 1.01 (average 0.98, n = 10), i.e., metaluminous to slightly
peraluminous. In line with major elements, contents of Rb (85–131 ppm), Ba
(649–1141 ppm), and Sr (426–641 ppm) have restricted ranges. In contrast,
a wider range in values of Nb (2.2–10.8 ppm), Y (6.7–1 1.2 ppm), Yb (0.5–1.3 ppm),
and Zr (99–150 ppm) is observed (T able 2). The variation in the heavy REE con-
tents is reflected in the variable REE patterns with a wide range of values of the
La
N
/Yb
N
ratio from 10 to 24 (Fig. 10C). There are small negative Eu anomalies
in the studied samples (Eu
N
/Eu
N
* = 0.79–0.97).
The leucogranite sample has contents of SiO
2
= 73%, TiO
2
= 0.1%, FeO
t
= 1 .2%,
CaO = 1.7%, and alkalis = 7 .7% (Na
2
O/K
2
O ratio ≤1) (Table 2). It is classified as
granite in the alkalis versus silica classification diagram, and in the high-K field
on a K
2
O versus SiO
2
diagram with tendency toward the leuco granite field
(Figs. 10A, 10B). ASI value is equal to 1.13, i.e., peraluminous. The leucogranite
has contents of Rb = 127 ppm, Ba = 415 ppm, Sr = 249 ppm, Nb < 0.2 ppm, Y
= 3.1 ppm, Yb = 0.4 ppm, and Zr = 54 ppm (T able 2). It exhibits a steeply dipping
REE pattern (La
N
/Yb
N
= 28) with a positive Eu anomaly (Eu
N
/Eu
N
* = 1 .26) (Fig. 10C).
The Topaz Lake granodiorite, the youngest unit in the suite, also shows a
restricted range in major elements, although in comparison to the KLG it tends
toward more felsic end members. It shows ranges in SiO
2
contents between
67%–71%, TiO
2
= 0.2%–0.5%, FeO = 2%–3.6%, CaO = 0.9%–3.6%, and alkalis
= 6.5%–8.3% (Table 2). The Na
2
O/K
2
O ratio is usually >1. The samples plot in
the granodiorite and granite fields in the alkalis versus silica classification
diagram, and in the medium- to high-K fields on a K
2
O versus SiO
2
diagram
(Figs. 10A, 10B). ASI values range from 0.99 to 1.05 (average 1.01, n = 13), i.e.,
metaluminous to slightly peraluminous. The granitic samples have contents of
Rb = 101–166 ppm, Ba = 125–1227 ppm, Sr = 408–615 ppm, Nb = 6.9–9.7 ppm, Y
= 8.5 ppm (two values), Yb = 0.5–0.7 ppm, and Zr = 99–133 ppm (Table 2). The
REE patterns show La
N
/Yb
N
ratios between 22 and 29 and weak to moderate
Eu anomalies (Eu
N
/Eu
N
* = 0.8–0.9) (not shown).
Schlieren Zone in the Kinney Lakes Granodiorite
Schlieren in the KLG have relatively low SiO
2
contents (54%–61%) but high
TiO
2
(1.2%–1.8%), FeO (8.1%–12.7%), CaO (3.9%–6.2%), and alkalis (5.7%–6.6%)
(T able 2). The Na
2
O/K
2
O ratio is usually <1. Samples plot between the monzodi-
orite and tonalite fields in the alkalis versus silica classification diagram, and
in the high-K field on a K
2
O versus SiO
2
diagram (Figs. 10A, 10B). ASI values
range from 0.80 to 0.86 (average 0.82, n = 3), i.e., metaluminous. They have
contents of: Rb = 102–154 ppm, Ba = 438–621 ppm, Sr = 212–368 ppm, Nb
= 16–26 ppm, Y = 28–46 ppm, Yb = 2.5–4 ppm, and Zr = 268–405 ppm (Table 2).
REE patterns show La
N
/Yb
N
ratios from 10 to 19 and moderate Eu anomalies
(Eu
N
/Eu
N
* = 0.66–0.73) (Fig. 10C).
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
242
17
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Leucocratic bands have high SiO
2
contents (66%–69%) but low TiO
2
(0.37%–
0.55%), FeO (3.2%–3.8%), CaO (2.9%–4.5%), and alkalis (6.4%–7 .1%) (Table 2)
with variable values of the Na
2
O/K
2
O ratio. The samples plot in the granodiorite
field in the alkalis versus silica classification diagram, and they plot in the medium-
to high-K fields on a K
2
O versus SiO
2
diagram (Figs. 10A, 10B). ASI values range
from 0.96 to 1 .03 (average 0.99, n = 3), i.e., metaluminous to slightly peraluminous.
They have relatively high contents of Ba (423–916 ppm) and Sr (357–579 ppm) but
lower Rb (53–106 ppm), Nb (2.7–5 ppm), Y (8.9–12 ppm), Yb (0.8–1 .1 ppm), and Zr
(94–1 18 ppm) (T able 2). REE patterns show La
N
/Yb
N
ratios from 12 to 21 and weak
or no Eu anomalies (Eu
N
/Eu
N
* = 0.83–0.99) (Fig. 10C).
■ Rb-Sr AND Sm-Nd WHOLE-ROCK ISOTOPE COMPOSITIONS
A reference age of 95 Ma is used for calculating isotope compositions at
the time (t) of magma crystallization in the KLG, from the weighted U-Pb sen-
sitive high-resolution ion microprobe (SHRIMP) zircon age of one granodiorite
sample (Leopold, 2016). Compared with the published data of eight samples
of the KLG (0.7056 ≤
87
Sr/
86
Sr
(t)
≤ 0.7058) (Macias, 1996), our samples (SWP -31 1
and SWP-312) of the KLG yield similar
87
Sr/
86
Sr
(t)
values of 0.7060 and 0.7057
(Table 3). The leucogranite and granodiorite samples have ε
Nd(t)
values of −2.8
and −2.6 respectively (Table 3).
Diorite
Monzonite
Syenite
Gabbro
Gabbro
diorite
Monzo
diorite
Quartz
monzonite
Foid syenite
Foid monzo-
syenite
Foid
gabbro
Granite
Foidolite
2.5
10
25
Granodiorite
Alkaline
Midalkaline
40 50 60 70 80
SiO (%)
2
2
4
6
8
10
12
14
Monzo
gabbro
Tonalite
Subalkaline
Na O + K O (%)
2 2
Foid monzo-
gabbro
1
La Ce Pr NdPmSmEuGd Tb Dy Ho ErTmYb Lu
1000
Rock/chondrites
10
100
1
10
100
1000
Cs Rb Ba Th K Nb La Ce Sr Nd ZrSm Ti Y Yb
Rock/bulk earth
KLG samples
40 50 60 70 80
0
1
2
3
4
5
K O (%)
2
Leucogranites
Low-K
Medium-K
SiO (%)
2
High-K
SPIS
Kinney Lakes granodiorite
Schlieren samples
Leucocratic band samples
Bummers Flat granodiorite
Mafic intrusions in Bummers
Flat granodiorite
Leucogranite samples
Topaz Lake granodiorite
A
D C
B
Figure 10. (A, B) Total alkali versus silica diagram
after Middlemost (1994) (A) and K
2
O versus SiO
2
diagram with classification boundaries after Le Mai -
tre (1989) (B) for the studied samples. The alkaline,
mid-alkaline, and subalkaline magmatic lineages
in (A) are defined by sigma isopleths (after Ritt -
mann, 1957). (C) Chondrite-normalized (after Sun
and McDonough, 1989) rare earth element plot for
the samples of igneous layering and the host Kinney
Lakes granodiorite (KLG). (D) Bulk earth-normalized
(after Hickey et al., 1986) values plot for the samples
of igneous layering and the host KLG. SPIS—Sonora
Pass Intrusive Suite.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
243
18
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Two samples (SWP-290-1 and SWP-307-1) of schlieren have
87
Sr/
86
Sr
(t)
val-
ues of 0.7056 and 0.7057 , and ε
Nd(t)
−3.1 (Table 3). The leucocratic counterparts
of these layers (samples SWP-290-2 and SWP-307-2) yield
87
Sr/
86
Sr
(t)
and ε
Nd(t)
values of 0.7058 and ~−3, respectively (Table 3).
The analyzed samples are depleted in Rb, with whole-rock contents
between 89 and 154 ppm (Table 2). One granodiorite sample (SWP-312) and
two leucocratic bands (SWP-290-2 and SWP-307-2) have high Sr contents
(357–496 ppm), but the two schlieren and leucogranite have <250 ppm, result-
ing in a wide range of Rb/Sr ratios. Including the KLG leucogranite (sample
SWP -31 1), the six samples fall on an errorchron corresponding to an age of 94
± 19 Ma (mean square of weighted deviates or MSWD = 13). However, if the
leucogranite sample is not considered, an isochron (MSWD = 2.6) is defined
with an age of 91.9 ± 9.6 Ma and an initial
87
Sr/
86
Sr ratio of 0.70581 ± 0.00018
(Figs. 11A, 11B).
■ DISCUSSION
Origin of the Schlieren in the Kinney Lakes Granodiorite
Petrographic observations indicate that both schlieren and leucocratic
bands contain the same mineral assemblage as the host granodiorite but have
different modal abundances (Table 1; see also Reid et al., 1993; Solgadi and
Sawyer, 2008; Paterson, 2009). This is consistent with Sr and Nd whole-rock
isotope compositions, where the studied layered samples and the host granodi-
orite show restricted ranges of
87
Sr/
86
Sr
(t)
= 0.7056–0.7060 and ε
Nd(t)
= −2.6 to −3.1,
regardless of the wide range in SiO
2
content (54%–74%) (Fig. 1 1C; Tables 2, 3).
They plot in the
87
Sr/
86
Sr
(t)
versus ε
Nd(t)
diagram without any dispersion, indicat-
ing that all samples crystallized from the same magma (Fig. 1 1D). In the same
way, the acceptable isochron of five samples with an age of 91.9 ± 9.6 Ma and
2.2 0.6 1.0 1.4 1.8
0.7065
0.7085
87 86
Sr/ Sr
87 86
Rb/ Sr
Age = 91.9 9.6 Ma
87 86
Initial Sr/ Sr = 0.70581 0.00018
MSWD = 2.6
±
±
2.2 0.6 1.0 1.4 1.8
0.7065
0.7085
87 86
Sr/ Sr
87 86
Rb/ Sr
Age = 94 19 Ma
87 86
Initial Sr/ Sr = 0.70582 0.00037
MSWD = 13
±
±
Sch
LB
KLG
Lg
50 60 70 80
0.705
0.707
87 86
Sr/ Sr
(t)
Kinney Lakes granodiorite
Leucocratic band samples
Schlieren samples
0.709 0.705
87 86
Sr/ Sr
(t)
0
2
+2
4
6
0.707
ε
Nd(t)
SiO (%)
2
SPIS field
Leucogranite sample
B A
D C
Figure 11. (A) Rb-Sr errorchron plot which includes
two schlieren samples (Sch), two leucocratic band
samples (LB), a leucogranite (Lg), and one sample
of the Kinney Lakes granodiorite (KLG). (B) Rb-Sr
isochron plot excluding sample SWP311, a leuco-
granite (Lg in panel A). The data are given in T able 3.
MSWD—mean square of weighted deviates. (C, D)
SiO
2
versus
87
Sr/
86
Sr initial (C) and
87
Sr/
86
Sr initial
versus ε
Nd(t)
plots (D) for the studied granitic rocks.
In C, the Sonora Pass Intrusive Suite (SPIS) field is
formed by one sample of this work and 20 samples
(nine samples of KLG and 11 samples of T opaz Lake
granodiorite) from Macias (1996).
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
244
19
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
an initial
87
Sr/
86
Sr = 0.70581 is consistent with a co-magmatic origin (Fig. 1 1B),
and further suggests that the Rb-Sr age records whole-rock system closure
immediately after crystallization of the pluton at 95.6 ± 1.5 Ma (Leopold, 2016).
The KLG leucogranite sample (SWP -31 1) excluded in the isochron, with a high
abundance of secondary muscovite, probably reflects some perturbation of
the Rb-Sr systematics after crystallization by a local fluid event.
Chemically, the schlieren, leucocratic band samples, leucogranite, and host
granodiorite of Kinney Lakes define a clear trend on major-element diagrams
(trend 1 in Fig. 12). The SiO
2
versus FeO
t
+ MgO + TiO
2
and the TiO
2
+ FeO
t
+
K
2
O versus P
2
O
5
plots show trends with correlation coefficients ( r) close to 0.9
for all samples (Figs. 12A, 12B). Trend 1 is consistent with crystal fractionation,
in which dense mafic crystals represented in the schlieren are separated from
KLG magma resulting in a feldspar- and silica-rich melt (leucocratic bands
and leucogranite). A similar trend (trend 2 in Fig. 12) is shown in the same
diagrams for the granitic samples of the SPIS (excluding the layered struc-
tures and the leucogranite), where the T opaz Lake granodiorite samples have
a more differentiated composition than KLG samples. However, on the FeO
t
versus CaO plot, only the SPIS granitic samples form a clear trend (trend 2)
t
TiO + FeO + MgO (%)
2
40 50 60 70 80
0
10
20
SiO (%)
2
r = 0.96
Hbl + Bt
Qtz
100 1 10
Eu (ppm)
0.1
100
1000
Sr (ppm)
r = 0.41
Pl
r = 0.99
0 10 20 30
0
10
20
30
40
Sc (ppm)
t
FeO + MgO (%)
r = 0.77
TREND 2
TREND 1
TREND 2
TREND 1
TREND 2
Kinney Lakes granodiorite
Leucocratic band samples
Schlieren samples
Leucogranite sample
Topaz Lake granodiorite
r = 0.95
r = 0.97
Hbl + Pl
3
5
7
CaO (%)
TREND 2
0 5 10 15
1
t
FeO
0 10 20 30
0
0.1
0.2
0.3
0.4
P O (%)
2 5
Bt + Ap
r = 0.95
r = 0.61
TREND 1
TREND 2
t
TiO + FeO + K O (%)
2 2
Hbl
A
D E
C B
Figure 12. Plots for the Sonora Pass Intrusive Suite and igneous layering samples. (A) SiO
2
versus TiO
2
+ FeO
total
+ MgO. (B) TiO
2
+ FeO
t
+ K
2
O versus P
2
O
5
. (C) FeO
t
versus CaO.
(D) FeO
t
+ MgO versus Sc. (E) Eu versus Sr. Note log scale in E. Arrows reflect crystallization of different minerals based on the whole-rock geochemistry; the sense of the arrow
indicates a greater modal mineral proportion. T rend 1 is defined by schlieren, leucocratic band samples, leucogranite, and samples of Kinney Lakes granodiorite. T rend 2 is formed by
granitic samples of the Sonora Pass Intrusive Suite, excluding the layered structures and the leucogranite. Hbl—hornblende; Bt—biotite; Qtz—quartz; Ap—apatite; Pl—plagioclase.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
245
20
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
(Fig. 12C). Schlieren samples move away from the differentiation trend 2,
which is dominated by fractionation of hornblende + plagioclase, toward
values higher in FeO
t
by almost exclusive fractionation of mafic minerals and
subordinate plagioclase.
Trace-element abundances in the layered structures are also in concor-
dance with a selective separation of mafic minerals from the KLG magma.
The high modal proportion of biotite, hornblende, and accessory minerals
(including zircon, sphene, apatite, and opaque minerals) in the schlieren is
reflected in the enrichment of REEs, Y , and Zr relative to the leucocratic band
samples and the leucogranite (Fig. 10D). Moreover, the acceptable correlation
between FeO
t
+ MgO against Sc, which is highly fractionated by hornblende,
describes a good correlation (r = 0.99) for trend 1 (Fig. 12D). In contrast,
samples of trend 1 show no correlation in the logarithmic plot of Eu versus
Sr concentrations (Fig. 12E). The Sr decrease with decreasing Eu showing
the role of feldspar fractionation is only evident in SPIS samples (trend 2).
Therefore, during schlieren formation, feldspar fractionation was decoupled
in relation to that of the mafic minerals. We envisage that this fractionation
at the solidification front could have been aided by external factors, such as
strain rate and imposed stress, to expel lighter minerals and/or felsic liquid
from the schlieren.
REE patterns of both layered and non-layered rocks are mostly parallel
(Fig. 10C) but with differences in both their abundances and in the magnitude
of the Eu anomaly. This behavior, also reported in other layered structures in
the Sierra Nevada (e.g., Reid et al., 1993; Solgadi and Sawyer, 2008), implies
that the schlieren-forming process did not discriminate between the different
types of REE-bearing minerals hosted in the KLG magma. These different
minerals, mostly accessory phases, were distributed almost homogeneously
in the study area but concentrated in variable proportions between schlieren
and KLG, which supports the idea that schlieren formed from the KLG. This is
also supported by the parallel REE patterns of the schlieren samples to those
of the host KLG. Schlieren show a higher abundance of most of the elements
and a moderate negative anomaly of Eu (Eu
N
/Eu
N
* = 0.66–0.73) due to a lower
plagioclase content. The total REE content in the schlieren (260–505 ppm) far
exceeds the range of values of the KLG (76–130 ppm). In contrast, the lower
modal content of the mafic minerals and high modal content of feldspar in
the leucogranite sample show a depletion in heavy REEs, with a total con-
tent of 48 ppm and a positive Eu anomaly (Eu
N
/Eu
N
* = 1.28). The leucocratic
bands show both identical REE patterns and a similar content of total REEs
(78–138 ppm) to those of KLG, except for a weak or absent Eu anomaly (Eu
N
/Eu
N
*
= 0.84–0.99), indicating a higher fraction of plagioclase.
From a mass-balance perspective, a simple two-component linear mixing
calculation (e.g., Fourcade and Allegre, 1981) on major elements suggests that
the mixing between the leucocratic band and schlieren (model A in T able 4) can
reproduce the bulk composition of the KLG with a very acceptable regression
line (r
2
= 0.99). However, this mixture of 14% schlieren and 86% leucocratic
band is unrealistic in a closed-system model (where there is no melt migra-
tion outside of the schlieren zone) because proportions of the schlieren and
leucocratic bands in the field are close to 1:1 (Figs. 3A, 7F). Such a model
requires that a significant proportion of magma, compositionally similar to
the leucocratic bands, is expelled toward the chamber interior, or to shallower
levels of the plumbing system. Using the leucogranite sample as the differ-
entiated end member (model B in Table 4) provides a similar value of the
regression line (r
2
= 0.95) to model A but with a bulk mixture of ~42% schlie-
ren and 58% leucogranite. This implies that leucogranite dikes may reflect
the fractionated melts removed from schlieren, rather than the felsic layers
found in between the schlieren.
Therefore, we speculate that the magmatic system was an open system,
where leucogranitic melts may have been extracted from the schlieren zone.
The melt could either have escaped the system to the chamber interior or to
shallower levels, readily diffused, and mixed back in with melts of the KLG,
or locally pooled and formed dikes and leucogranite bodies. This can explain
why the chemical signature of the differentiated end member from schlieren
is difficult to trace.
In summary, whole-rock chemical data support the hypothesis that the
main mechanism of schlieren formation was fractional crystallization in an
isotopically closed system, mainly by effective segregation of mafic min -
erals plus accessory phases with drainage and migration of leucogranitic
melts back into the KLG or to shallower levels. The decoupling of plagioclase
± quartz in the schlieren chemistry illustrates that liquid-crystal fractionation
was selective and did not depend solely on gravitational forces, but that other
factors, such as deformation-assisted reorganization of the crystal mush, were
required to promote leucogranitic melt migration away from the schlieren
zone (see next).
TABLE 4. TWOCOMPONENT MIXTURE CALCULATION
C
i
M
= xC
i
A
+ (1 – x)C
i
B
C
i
M
= xC
i
A
+ C
i
B
– xC
i
B
C
i
M
– C
i
B
= x(C
i
A
– C
i
B
)
Model A
C
i
A
—average of leucocratic band samples (SWP2783, SWP2902, and SWP3072)
C
i
B
—average of schlieren samples (SWP2782, SWP2901, and SWP3071)
C
i
M
—average of Kinney Lakes granodiorite (KLG) samples (SICN17, SICN25,
SWP312, S102, S723, S724, S522, S228, R529, and S288)
R
2
= 0.99 C
i
A
in the mixing = 86%
Model B
C
i
A
—leucogranite sample (SWP311)
C
i
B
—average of schlieren samples (SWP2782, SWP2901, and SWP3071)
C
i
M
—average of KLG samples (SICN17, SICN25, SWP312, S102, S723, S724,
S522, S228, R529, and S288)
R
2
= 0.95 C
i
A
in the mixing = 58%
Notes: Twocomponent mixture calculation based on Fourcade and Allegre (1981)
C
i
M
—content of i element in the mixture formed by two components; C
i
A
—content of
i element in felsic member; C
i
B
—content of i element in mafic member; x—weight
proportion of felsic to mafic member in the mixture for each element (i).
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
246
21
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Spatiotemporal and Structural Relationships in the Schlieren Zone
Field and Structural Data
The orientation of the schlieren zone parallel to the orientation of the KLG
margin indicates that it was controlled by the geometry of the magma cham-
ber. The growth and younging direction of the schlieren zone, determined by
mineral gradation of the layers and cross-cutting relationships, was toward
the interior of the chamber, where regional magmatic fabrics developed
coincidentally with the formation and deformation of schlieren (Figs. 2, 4, 6;
Spatiotemporal Relationships between Structural Elements of the Schlieren
Zone section).
The intrusive contact between the BFG and the KLG is highly discordant
and steep, which, together with the existence of numerous BFG host-rock
blocks and KLG cognate inclusions in the schlieren zone, supports the inter-
pretation that the host rock was removed and earlier parts of the KLG margin
were recycled during the formation of the schlieren (e.g., Irvine et al., 1998).
This is most noticeable in the southern part of the intrusive contact. In this
area, the greater variability in the dip of schlieren (Fig. 2) may be explained
by the sinking of blocks during schlieren formation. At the same time, enclave
swarms commonly truncate schlieren packages, and are themselves truncated
by overlying schlieren packages, implying episodic, or cyclic, erosional and
magma mingling processes during schlieren formation, all of which likely
occurred in situ (e.g., Paterson et al., 2016).
The associated magmatic faults and folds within schlieren packages suggest
that the crystal mush was deformed magmatically in a highly dynamic environ-
ment. Faults affecting both planar and multiply folded schlieren packages attest
to several distinct deformation events during both formation and collapse of
the crystal mush (Paterson, 2009; Humphreys and Holness, 2010). Folds have
variable axial-planar orientations, reflecting the heterogeneity of folding and
deformation at the outcrop scale during magma avalanching. Magmatically
folded schlieren are commonly truncated by younger packages of folded and
faulted schlieren recording multiple phases of deformation at a single outcrop.
Petrographic and Microstructural Data
Schlieren, leucocratic bands, and the KLG unit show predominantly
magmatic microstructures (Fig. 9). Some quartz grains show evidence of intra-
crystalline deformation in all samples, which could be attributed to stresses
during cooling of late-stage melts. A small number of feldspars show defor-
mation twinning. However, the overall amount of deformation is minimal and
localized, and there is only slight variation between the type and intensity of
the microstructures contained within schlieren and the host granodiorite. This
suggests that the overall crystal framework is magmatic with sub-magmatic to
solid-state microstructures developed locally due to crystal-crystal stresses and
not a throughgoing tectonic strain. Both truncation of compositional zoning
in feldspar and interstitial quartz-feldspar in the schlieren are consistent with
mineral compaction and melt removal processes in the crystal mush (Fig. 9C).
A Model for the Generation of the Schlieren Zone in the Kinney Lakes
Granodiorite
Our model for the formation of the schlieren zone involves the construction
of an incipient solidification front on an upper margin of the chamber (Fig. 13).
This front, primarily controlled by the geometry of the magma-chamber margin,
experiences a larger heat loss than the interior, resulting in an increase in the
crystallinity, density, and viscosity of the magma (Marsh, 1996). These denser,
crystal-rich magmas become unstable against the boundary with the more-in-
terior, less-viscous magmas and occasionally collapse, leading to downward
transport and avalanching of dense crystal slurries by gravity-assisted flow
«
Gravity collapse
Magmatic stoping
Normal faulting
Magma mush avalanche
Magmatic folds
Folding by
stoped blocks
«
Gravity collapse
MAGMA CHAMBER
New magma
intrusion
Avalanche leaves behind
schlieren, tracing the
flow of the magma mush
Formation of schlieren
with migration of felsic
melt to the interior of
chamber
Magmatic deformation of the
schlieren zone in a wall of
the chamber
Figure 13c
MAGMA CHAMBER
Some minerals and blocks
causing instability at the
pluton boundary
SOLIDIFICATION FRONT
HOST ROCK
A
B
C
Figure 13. Interpretive block diagrams illustrating the magmatic evolution at the margin of the
Kinney Lakes granodiorite (KLG). (A) Formation of an early solidification front in the KLG. (B) Ini -
tiation of gravitational collapse of the solidification front and formation of layering by laminar
flow. Instability causes crystals to avalanche along the pluton boundary, and incorporates pieces
of host rocks and cognate inclusions. (C) Advanced stage of magma avalanching with syn-mag-
matic deformation structures (fault and fold) related to the transport and reorganization of the
crystal-rich pile on a wall of the chamber. The propagation of normal magmatic faults shows
directions of slip parallel to the contact margin.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
247
22
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
(Irvine et al., 1998; Solgadi and Sawyer, 2008; Žák and Paterson, 2010; Paterson
et al., 2016, 2019). Based on field relationships, schlieren formation was discon -
tinuous and repeated in time. Magmatic erosion and redeposition of schlieren
was pervasive, and commonly the schlieren packages record evidence of early
and later repeated deformation events during magma avalanching.
Field evidence suggests that other mechanisms such as stoping and the
presence of a fluid phase could also have triggered a large-scale destabilization
of the solidification front. Blocks of host rocks falling into the crystal mush may
have created instability by increasing the weight of the mush and exceeding
the yield stress of the magma. The presence of a late, exsolved fluid phase,
recorded by miarolitic cavities, vugs, and pegmatites, could have generated
progressive fluidization by excess pressure in the hydrogranular medium,
which is generally governed by multi-phase interactions of silicate liquid, crys-
tals, and bubbles (e.g., Sumita and Manga, 2008; Petford, 2009; Bergantz et
al., 2017). In this context, the downward collapse of the crystal-rich margin
transported crystals and surrounding magma along the pluton margin, gen-
erating the most commonly observed schlieren with a normal density grading
and inverse size gradation toward the top of the layer. Their formation was
accompanied by strain, compaction, and filter pressing. As a result, both the
dense crystals (biotite, hornblende, sphene, and opaque minerals) and in some
cases small mafic enclaves are sorted against steeply dipping walls. However,
the chemistry suggests that the simple model of physical, mechanical separa-
tion of suspended minerals from the magma, combined with crystal settling
and compaction, does not fully explain the composition or structural features
of the schlieren, and that additional processes are required. During magma
mush avalanching, local convection and late hypersolidus strain may have
promoted density-driven mineral sorting against steeply dipping walls, forming
a transient crystal framework (e.g., Irvine, 1987; Hodson, 1998, Blanchette et
al., 2004). This framework was connected to felsic intergranular melt, which
may have been displaced toward the interior of the magma body, outside in
the BFG host rock or to higher (unexposed) levels. The felsic intergranular melt
could have partially contributed to the formation of the leucocratic bands, or
remixed with the host magma. We envision that both gravity and local con-
vection driven by internal thermal, compositional, and rheological gradients
(Martin et al. 1987) acted on the hydrogranular slurry and mechanically sepa-
rated the silica-rich melt from the dense crystals.
Kinematic indicators in the schlieren zone are consistent with syn-magmatic
deformation and displacement of magma downwards. The propagation of
normal magmatic faults with directions of slip parallel to the contact margin
resembles slope instability, where material moves downslope in response to
gravity (Figs. 2, 5). We find that magmatic faulting and folding record local
deformation of the crystal mush after magma avalanching eroded and depos-
ited schlieren. Faults and folds are evidence supporting slumping and settling
of the crystal-rich pile. Additionally, magmatic folding and faulting could have
been triggered by tectonic or internal magma-chamber processes at a larger
scale. For example, the folds with NNW-SSE–trending axial planes with steep
dip (~80°) and contact-parallel strike (see above) are also consistent with the
shortening direction of regional tectonic forces. Magma-chamber processes
could have included the intrusion of new magma batches, or overpressure-
or convection-related forces in the interior parts of the magma body directed
toward chamber margins. However, to test this hypothesis further, it will be
necessary to obtain a greater number of structural measurements of folds.
Possible Causes of Magma Avalanches in Chambers
The formation of a schlieren zone by the growth and collapse of a solidifica -
tion front requires not only enough melt present in the mush to accommodate
strain during magmatic faulting, folding, and stoping of the xenolith blocks,
but also a strong crystal framework to preserve the layering relatively intact,
especially during deformation. In other words, the magma must be sufficiently
rich in crystals and therefore denser than its surroundings to descend into the
chamber, but equally be plastic enough to flow under strain, that is, behave
as a visco-plastic material. (e.g., Bergantz et al., 2015; Schleicher et al., 2016).
In this context, a number of possible causes could trigger an extensive
visco-plastic collapse of a magma mush: (1) differential cooling in the magma
chamber, in which chamber margins cool faster than its interior, would pro-
mote the growth of crystal-rich solidification fronts that are prone to collapse
into the hotter interior (Marsh, 1996, 2006, 2013, 2015; Žák and Paterson, 2010);
(2) density inversion of crystal-rich peripheral mush over hotter crystal-poor
interior magma could trigger magma avalanching preferentially in the upper
part of the chamber (roof) or in irregular steep sections of the wall; (3) the
arrival of a new magma batch could allow the crystal mush to unlock and
fluidize (e.g., Solgadi and Sawyer, 2008; Burgisser and Bergantz, 2011; Pat -
erson et al., 2016; Bergantz et al., 2017), which could contribute to magma
avalanching either by magma injection to the solidification front that pene -
trates and spreads magma by porous media flow through the crystal mush,
or by the fluid phase generating an excess pressure (by bubble formation) in
the hydrogranular mixture (e.g., Sumita and Manga, 2008; Petford, 2009); (4)
the presence of host-rock blocks and cognate inclusions in the margin could
increase the overall density of the solidifying crystal-melt pile, whereby the
erosive component of the magma avalanching process can act to remove ear-
lier-formed parts of the margin (cognate inclusions) and increase the density
of the magma pile, helping it to overcome the flow resistance of the medium
(e.g., Paterson et al., 2016); and (5) natural seismic sources such as tectonic and
volcanic earthquakes could induce destabilization by fluidization and decrease
in pressure in the crystal mush (e.g., Sumita and Manga, 2008; Paterson et al.,
2019, and references therein).
We envisage that most of the causes described for the development of
large-scale (10
2
–10
3
-m-wide) magma avalanches in peripheral mush zones
occur in large, long-lived magma-mush chambers that mechanically mingle
and mix, rather than small, isolated magma bodies (e.g., Solgadi and Sawyer,
2008; Pinotti et al., 2016). A long-lived, thermally mature magmatic plumb-
ing system allows potentially multiple magma chambers to form at shallow
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
248
23
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
crustal levels (e.g., Alasino et al., 2017). Thus, the generation of these multi-
ple chambers by the ascent of different magma batches can be promoted by
gravitational collapse and local convection of a ductile downward flow of the
earlier magma mushes.
Implications for the Behavior of Long-Lived Magmatic Systems, and
the Significance of Structural and Compositional Diversity in Upper-
Crustal Magma Chambers
Uranium-lead (U-Pb) zircon geochronology illustrates that the SPIS is an
example of a long-lived magmatic system, emplaced over at least 2.9 m.y.,
and possibly >8.1 m.y. (Leopold, 2016), comparable to other zoned intrusive
suites across the Sierran crest (e.g., Coleman et al., 2004; Frazer et al., 2014;
Paterson et al., 2016). Like other Sierran intrusive suites, the SPIS displays
map-scale textural and compositional homogeneity within identified units (e.g.,
restricted Sr isotope values in the SPIS; Macias, 1996; Leopold, 2016). However,
it displays remarkable mesoscale and microscale structural heterogeneity, the
schlieren zone being a key example (see also Burgisser and Bergantz, 2011).
Schlieren were created and deformed in the magmatic state. The scale of the
schlieren zone, ~800 m wide, suggests that an extensive area of the crystal
mush was mobilized at the outermost (oldest and coolest) margin of the SPIS.
The complexity in the structure within a single outcrop suggests that multiple,
repeated avalanche events took place with the avalanche area, or front, migrat-
ing inwards. Additionally, it suggests that this system was dynamic even at
shallow crustal levels (2.0–3.5 ± 0.6 kbar; Macias, 1996).
The structurally complex schlieren zone suggests that a high level of insta-
bility and mobility existed at the margin zone of the granitic magma mush.
Our field observations of magmatic faulting and magmatic fold structures
are remarkably similar to the model observations from the magma “mixing
bowl” of Bergantz et al. (2015). In addition, this model predicts that due to
the hydrogranular behavior of the mush, variable gradients in porosity exist
within the mush (Bergantz et al., 2015; Schleicher and Bergantz, 2017), which
may explain the formation of the leucogranite dikes as well as the re-intrusion
and reorganization of schlieren layers post-avalanching.
The schlieren zone at the KLG margin is one spectacular example of local
compositional and structural diversity generated at the emplacement level, in
an upper-crustal magma body. While the mineral components of schlieren and
leucogranite layers are locally sourced from the host magma mush, dynamic
processes operated to mechanically sort crystals by density and mode. The
same dynamic processes (e.g., avalanching, mixing) may have also occurred
in parts of the KLG unit where schlieren are sparse. KLG magma may have
later erased these layers and/or modified them by protracted magma mixing,
mobilization, and/or magma flow in hotter, more melt-rich areas of the chamber.
This would reproduce the map-scale homogeneous appearance of the inner
KLG, which experienced a slower cooling rate than the outermost margin.
In the case of the schlieren zone at the KLG margin, however, a rich array of
structures is preserved, perhaps capturing different magmatic conditions than
found in other KLG domains.
■ CONCLUSIONS
Whole-rock data reveal that the granitic rocks of the Sonora Pass Intrusive
Suite in the northern Sierra Nevada have a relatively narrow range of silica
content (61%–71%) and show geochemical signatures typical of high-K calc-al-
kaline granites, with metaluminous to slightly peraluminous compositions.
Evidence presented here suggests that the southwestern intrusive margin
of the KLG is a remarkable example of the dynamic processes that can occur
along a pluton–wall rock contact. Large-scale downward magma avalanching
events at an early-formed solidification front generated an ~800-m-wide region
of igneous layering (schlieren and leucocratic bands) recording the complex
interaction between magmatic structure formation and deformation at map to
outcrop scales. Schlieren (SiO
2
= 54%–61%) were formed by the accumulation
of medium-grained ferromagnesian minerals (hornblende + biotite), zircon,
sphene, apatite, and opaque minerals, and minor plagioclase and interstitial
quartz. Associated leucocratic bands (SiO
2
= 66%–69%) are dominated by
coarse-grained feldspar (plagioclase over microcline) + quartz with minor ferro-
magnesian and accessory minerals and secondary muscovite, and potentially
formed by a mix of melt removed from schlieren with melt remaining in the
surrounding host. A third group of rocks defined as leucogranite (SiO
2
= 73%)
are found in dikes that intrude schlieren packages. They have a similar mineral-
ogy to the leucocratic bands but are generally more felsic in composition. The
leucogranite dikes may reflect the fractionated melts removed from schlieren.
Whole-rock chemical data support that the main mechanism of formation of
schlieren and leucocratic bands was by fractional crystallization in an isotopi-
cally closed system, mainly by an effective segregation of mafic minerals plus
accessory phases with drainage and diffusion of leucogranitic melts back into
the KLG. The accumulation of mafic and accessory minerals along steep walls
displaced the intergranular felsic melt toward the interior of the pluton or to
shallower levels of the plumbing system by deformation-assisted mechanical
separation in the hydrogranular medium, promoted by both gravity and local
convection. This process was repeated and episodic, which allowed mixing,
magmatic erosion and recycling, re-intrusion, and magmatic deformation
(faulting and folding) to occur at the KLG margin.
ACKNOWLEDGMENTS
Financial support was provided by the National Science Foundation grant EAR-1624847 to
S.R. Paterson and a CONICET (Argentine National Council of Scientific and Technical Research)
external fellowship awarded to P .H. Alasino for his research stay at the University of Southern
California, Los Angeles, California, USA. L. Ardill and K. O’Rourke are thanked for the assis-
tance in mapping work. We thank M. Holness (University of Cambridge) and an anonymous
reviewer for their very constructive reviews that significantly improved the manuscript. We
also would like to thank A. Whittington (editor) for his helpful comments and suggestions on
our manuscript.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
249
24
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
REFERENCES CITED
Alasino, P .H., Larrovere, M.A., Rocher, S., Dahlquist, J.A., Basei, M.A.S., Memeti, V., Paterson,
S., Galindo, C., Grande, M.M., and da Costa Campos Neto, M., 2017 , Incremental growth of
an upper crustal, A-type pluton, Argentina: Evidence of a re-used magma pathway: Lithos,
v. 284–285, p. 347–366, ht tps:// doi .org /1 0 .1 0 1 6 /j .lithos .20 1 7 .04 .0 1 3.
Armstrong, R.L., and Ward, P .L., 1993, Late Triassic to earliest Eocene magmatism in the North
American Cordillera: Implications for the Western Interior Basin, in Caldwell, W.G.E., and
Kauffman, E.G., eds., Evolution of the Western Interior Basin: Geological Association of Can-
ada Special Paper 39, p. 49–72.
Bachmann, O., and Bergantz, G.W., 2008, Deciphering magma chamber dynamics from styles of
compositional zoning in large silicic ash flow sheets: Reviews in Mineralogy and Geochem -
istry, v. 69, p. 651–674, ht tps:// doi .org /1 0 .21 38 /rmg .20 08 .69 .1 7 .
Barbey, P ., Gasquet, D., Pin, C., and Bourgeix, A.L., 2008, Igneous banding, schlieren and mafic
enclaves in calc-alkaline granites: The Budduso pluton (Sardinia): Lithos, v. 104, p. 147–163,
ht tps:// doi .org /1 0 .1 0 1 6 /j .lithos .20 07 .1 2 .0 04.
Bateman, P .C., 1992, Plutonism in the central part of Sierra Nevada batholith, California: U.S. Geo-
logical Survey Professional Paper 1483, 186 p., ht tps:// doi .org /1 0 .31 33 /pp1 483.
Bergantz, G.W., 2000, On the dynamics of magma mixing by reintrusion: Implications for pluton
assembly processes: Journal of Structural Geology, v. 22, p. 1297–1309, ht tps:// doi .org /1 0
.1 0 1 6 /S0 1 91 -81 41 (0 0)0 0 053 -5.
Bergantz, G.W., and Ni, J., 1999, A numerical study of sedimentation by dripping instabilities in
viscous fluids: International Journal of Multiphase Flow 25, v. 2, p. 307–320, ht tps:// doi .org
/1 0 .1 0 1 6 /S030 1 -9322 (98)0 0 050 -0.
Bergantz, G.W., Schleicher, J.M., and Burgisser, A., 2015, Open-system dynamics and mixing
in magma mushes: Nature Geoscience, v. 8, p. 793–796, ht tps:// doi .org /1 0 .1 038 /ngeo2534.
Bergantz, G.W., Schleicher, J.M., and Burgisser, A., 2017 , On the kinematics and dynamics of crys-
tal-rich systems: Journal of Geophysical Research: Solid Earth, v. 122, p. 6131–6159, https://
doi .org /1 0 .1 0 02 /20 1 7JB0 1 421 8.
Blanchette, F ., Peacock, T ., and Bush, J.W.M., 2004, The Boycott effect in magma chambers: Geo-
physical Research Letters, v. 31, L05611, ht tps:// doi .org /1 0 .1 029 /20 03GL0 1 9235.
Broxton, D.E., Warren, R.G., Byers, F .M., and Scott, R.B., 1989, Chemical and mineralogic trends
within the Timber Mountain–Oasis Valley Caldera Complex, Nevada: Evidence for multiple
cycles of chemical evolution in a long-lived silicic magma system: Journal of Geophysical
Research, v. 94, p. 5961–5986, ht tps:// doi .org /1 0 .1 029 /JB094iB05p05961.
Burgisser, A., and Bergantz, G.W., 2011, A rapid mechanism to remobilize and homogenize crys-
talline magma bodies: Nature, v. 471, p. 212–215, ht tps:// doi .org /1 0 .1 038 /nature09799.
Cao, W., Paterson, S., Memeti, V., Mundil, R., Anderson, J.L., and Schmidt, K., 2015, Tracking
paleodeformation fields in the Mesozoic central Sierra Nevada arc: Implications for intra-arc
cyclic deformation and arc tempos: Lithosphere, v. 7 , p. 296–320, ht tps:// doi .org /1 0 .1 1 30 /L389 .1.
Christensen, J.N., Halliday, A.N., Lee, D.-C., and Hall, C.M., 1995, In situ Sr isotopic analysis by
laser ablation: Earth and Planetary Science Letters, v. 136, p. 79–85, ht tps:// doi .org /1 0 .1 0 1 6
/0 0 1 2 -821X (95)0 0 1 81 -6.
Coleman, D.S., Gray, W., and Glazner, A.F ., 2004, Rethinking the emplacement and evolution of
zoned plutons: Geochronologic evidence for incremental assembly of the Tuolumne Intrusive
Suite, California: Geology, v. 32, p. 433–436, ht tps:// doi .org /1 0 .1 1 30 /G20220 .1.
Cooper, K.M., and Reid, M.R., 2003, Re-examination of crystal ages in recent Mount St. Helens
lavas: Implications for magma reservoir processes: Earth and Planetary Science Letters, v. 213,
p. 149–167 , ht tps:// doi .org /1 0 .1 0 1 6 /S0 0 1 2 -821X (03)0 0262 -0.
Costa, F ., Chakraborty, S., and Dohmen, R., 2003, Diffusion coupling between trace and major ele-
ments and a model for calculation of magma residence times using plagioclase: Geochimica
et Cosmochimica Acta, v. 67 , p. 2189–2200, ht tps:// doi .org /1 0 .1 0 1 6 /S0 0 1 6 -7037 (02)0 1 345 -5.
Davidson, J., Tepley, F ., III, Palacz, Z., and Meffan-Main, S., 2001, Magma recharge, contamina-
tion and residence times revealed by in situ laser ablation isotopic analysis of feldspar in
volcanic rocks: Earth and Planetary Science Letters, v. 184, p. 427–442, ht tps:// doi .org /1 0 .1 0 1 6
/S0 0 1 2 -821X (0 0)0 0333 -2.
Davidson, J.P ., Font, L., Charlier, B.L.A., and T epley, F .J., 2008, Mineral-scale Sr isotope variation in
plutonic rocks: A tool for unraveling the evolution of magma systems: Transactions of the Royal
Society of Edinburgh: Earth Sciences, v. 97 , p. 35–67 , ht tps:// doi .org /1 0 .1 0 1 7 /S026359330 0 0 0 1 504.
Davis, M., Koenders, M.A., and Petford, N., 2007 , Vibro-agitation of chambered magma: Journal
of Volcanology and Geothermal Research, v. 167 , p. 24–36, ht tps:// doi .org /1 0 .1 0 1 6 /j .jv olgeores
.20 07 .07 .0 1 2.
DeCelles, P .G., Ducea, M.N., Kapp, P ., and Zandt, G., 2009, Cyclicity in Cordilleran orogenic systems:
Nature Geoscience, v. 2, p. 251–257 , ht tps:// doi .org /1 0 .1 038 /ngeo469.
de Silva, S., Zandt, G., Trumbull, R., Viramonte, J.G., Salas, G., and Jiménez, N., 2006, Large
ignimbrite eruptions and volcano-tectonic depressions in the Central Andes: A thermome-
chanical perspective, in Troise, C., De Natale, G., and Kilburn, C.R.J., eds., Mechanisms of
Activity and Unrest at Large Calderas: Geological Society of London Special Publication 269,
p. 47–63, ht tps:// doi .org /1 0 .1 1 44 /GSL .SP .20 06 .269 .0 1 .04.
Fourcade, S., and Allegre, C.J., 1981, Trace elements behavior in granite genesis—A case study:
The calc-alkaline plutonic association from the Querigut complex (Pyrénées, France): Contri-
butions to Mineralogy and Petrology, v. 76, p. 177–195, ht tps:// doi .org /1 0 .1 0 07 /BF0 0371 958.
Frazer, R.E., Coleman, D.S., and Mills, R.D., 2014, Zircon U-Pb geochronology of the Mount Givens
Granodiorite: Implications for the genesis of large volumes of eruptible magma: Journal of
Geophysical Research: Solid Earth, v. 1 19, p. 2907–2924, h t t p s : / / d o i . o r g / 1 0 . 1 0 0 2 / 2 0 1 3 J B 0 1 0 7 1 6.
Giusso, J.R., 1981, Preliminary geologic map of the Sonora Pass 15-minute quadrangle, California:
U.S. Geological Survey Open-File Report 81-1170, scale 1:62,500.
Hickey, R.L., Frey, F .A., Gerlach, D.C., and Lopez‐Escobar, L., 1986, Multiple sources for basal-
tic arc rocks from the southern volcanic zone of the Andes (34°–41°S): Trace element and
isotopic evidence for contributions from subducted oceanic crust, mantle, and continental
crust: Journal of Geophysical Research: Solid Earth, v. 91, p. 5963–5983, ht tps:// doi .org /1 0
.1 029 /JB091iB06p05963.
Hodson, M.E., 1998, The origin of igneous layering in the Nunarssuit syenite, South Greenland:
Mineralogical Magazine, v. 62, p. 9–27 , ht tps:// doi .org /1 0 .1 1 80 /0 026461 98547 431.
Holness, M.B., Clemens, J.D., and Vernon, R.H., 2018, How deceptive are microstructures in granitic
rocks? Answers from integrated physical theory, phase equilibrium, and direct observations:
Contributions to Mineralogy and Petrology, v. 173, ht tps:// doi .org /1 0 .1 0 07 /s0 041 0 -0 1 8 -1 488 -8.
Huber, N.K., 1983, Preliminary geologic map of the Pinecrest quadrangle, central Sierra Nevada,
California: U.S. Geological Survey Miscellaneous Field Studies Map MF-1437 , scale 1:62,500.
Humphreys, M.C.S., and Holness, M.B., 2010, Melt-rich segregations in the Skaergaard Marginal
Border Series: Tearing of a vertical silicate mush: Lithos, v. 119, p. 181–192, ht tps:// doi .org
/1 0 .1 0 1 6 /j .lithos .20 1 0 .06 .0 06.
Irvine, T .N., 1987 , Layering and related structures in the Duke Island and Skaergaard intrusions:
Similarities, differences, and origins, in Parsons, I., ed., Origins of Igneous Layering: Dor-
drecht, D. Reidel Publishing Company, NATO ASI Series C196, p. 185–245, ht tps:// doi .org /1 0
.1 0 07 /978 -94 -0 1 7 -2509 -5_6.
Irvine, T .N., Andersen, J.C.O., and Brooks, C.K., 1998, Included blocks (and blocks within blocks)
in the Skaergaard intrusion: Geologic relations and the origins of rhythmic modally graded
layers: Geological Society of America Bulletin, v. 110, p. 1398–1447 , ht tps:// doi .org /1 0 .1 1 30
/0 0 1 6 -7 606 (1 998)1 1 0 <1 398: IBABWB>2 .3 .CO;2.
Jaupart, C., and Tait, S., 1995, Dynamics of differentiation in magma reservoirs: Journal of Geo-
physical Research, v. 100, p. 17 ,615–17 ,636, ht tps:// doi .org /1 0 .1 029 /95JB0 1 239.
John, D.A., 1983, Distribution, ages, and petrographic character of Mesozoic plutonic rocks, Walker
Lake 1° by 2° quadrangle, California and Nevada: U.S. Geological Survey Miscellaneous Field
Studies Map 1382-B, scale 1:250,000, ht tps:// doi .org /1 0 .31 33 /mf1 382B.
John, D.A., Schweickert, R.A., and Robinson, A.C., 1994, Granitic rocks in the Triassic-Jurassic
magmatic arc of western Nevada and eastern California: U.S. Geological Survey Open-File
Report 94-148, 61 p., ht tps:// doi .org /1 0 .31 33 /ofr941 48.
Kaiser, J.F ., de Silva, S., Schmitt, A.K., Economos, R., and Sunagua, M., 2016, Million-year melt–
presence in monotonous intermediate magma for a volcanic–plutonic assemblage in the
Central Andes: Contrasting histories of crystal-rich and crystal-poor super-sized silicic magmas:
Earth and Planetary Science Letters, v. 457 , p. 73–86, ht tps:// doi .org /1 0 .1 0 1 6 /j .epsl .20 1 6 .09 .048.
Karakas, O., Degruyter, W., Bachmann, O., and Dufek, J., 2017 , Lifetime and size of shallow magma
bodies controlled by crustal-scale magmatism: Nature Geoscience, v. 10, p. 446–450, https://
doi .org /1 0 .1 038 /ngeo2959.
Kirsch, M., Paterson, S.R., Wobbe, F ., Martínez Ardila, A.M., Clausen, B.L., and Alasino, P .H., 2016,
Temporal histories of Cordilleran continental arcs: Testing models for magmatic episodicity:
American Mineralogist, v. 101, p. 2133–2154, ht tps:// doi .org /1 0 .21 38 /am -20 1 6 -571 8.
Kistler, R.W., Chappell, B.W., Peck, D.L., and Bateman, P .C., 1986, Isotopic variation in the Tuolumne
Intrusive Suite, central Sierra Nevada, California: Contributions to Mineralogy and Petrology,
v. 94, p. 205–220, ht tps:// doi .org /1 0 .1 0 07 /BF0 0592937 .
Lackey, J.S., Cecil, M.R., Windham, C.J., Frazer, R.E., Bindeman, I.N., and Gehrels, G.E., 2012, The
Fine Gold Intrusive Suite: The roles of basement terranes and magma source development
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
250
25
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
in the Early Cretaceous Sierra Nevada batholith: Geosphere, v. 8, p. 292–313, ht tps:// doi .org
/1 0 .1 1 30 /GES0 07 45 .1.
Lahren, M.M., and Schweickert, R.A., 1989, Proterozoic and Lower Cambrian miogeoclinal rocks
of Snow Lake pendant, Yosemite-Emigrant Wilderness, Sierra Nevada, California: Evidence
for major Early Cretaceous dextral translation: Geology, v. 17 , p. 156–160, ht tps:// doi .org /1 0
.1 1 30 /0 091 -7 61 3 (1 989)0 1 7 <0 1 56: P ALCMR>2 .3 .CO;2.
Lahren, M.M., Schweickert, R.A., Mattinson, J.M., and Walker, J.D., 1990, Evidence of uppermost
Proterozoic to Lower Cambrian miogeoclinal rocks and the Mojave–Snow Lake fault: Snow
Lake pendant, central Sierra Nevada, California: Tectonics, v. 9, p. 1585–1608, ht tps:// doi .org
/1 0 .1 029 /TC0 09i0 06p0 1 585.
Le Maitre, R.W., ed., 1989, A Classification of Igneous Rocks and Glossary of Terms: Oxford, UK,
Blackwell Scientific, 193 p.
Leopold, M.B., 2016, Structure, construction, and emplacement of the Late Cretaceous Sonora Pass
intrusive suite: Central Sierra Nevada batholith, California [M.S. thesis]: San José, California,
San José State University, 100 p.
Lowery Claiborne, L., Miller, C.F ., Walker, B.A., Wooden, J.L., Mazdab, F .K., and Bea, F ., 2006, Track-
ing magmatic processes through Zr/Hf ratios in rocks and Hf and Ti zoning in zircons: An
example from the Spirit Mountain batholith, Nevada: Mineralogical Magazine, v. 70, p. 517–543,
ht tps:// doi .org /1 0 .1 1 80 /0 026461 067050348.
Macias, S.E., 1996, The Sonora Intrusive Suite: Constraints on the assembly of a Late Cretaceous,
concentrically-zoned granitic pluton of the Sierra Nevada batholith [M.S. thesis]: Seattle,
University of Washington, 66 p.
Marsh, B.D., 1996, Solidification fronts and magmatic evolution: Mineralogical Magazine, v. 60,
p. 5–40, ht tps:// doi .org /1 0 .1 1 80 /minmag .1 996 .060 .398 .03.
Marsh, B.D., 2006, Dynamics of magma chambers: Elements, v. 2, p. 287–292, ht tps:// doi .org /1 0
.21 1 3 /gselements .2 .5 .287 .
Marsh, B.D., 2013, On some fundamentals of igneous petrology: Contributions to Mineralogy and
Petrology, v. 166, p. 665–690, ht tps:// doi .org /1 0 .1 0 07 /s0 041 0 -0 1 3 -0892 -3.
Marsh, B.D., 2015, Magmatism, magma, and magma chambers, in Watts, A., ed., Treatise on
Geophysics (second edition), Volume 6: Crustal and Lithosphere Dynamics: Oxford, Elsevier,
p. 273–323, ht tps:// doi .org /1 0 .1 0 1 6 /B978 -0 -444 -53802 -4 .0 0 1 1 6 -0.
Martin, D., Griffiths, R.W., and Campbell, I.H., 1987 , Compositional and thermal convection in
magma chambers: Contributions to Mineralogy and Petrology, v. 96, p. 465–475, ht tps:// doi
.org /1 0 .1 0 07 /BF0 1 1 66691.
Memeti, V., Paterson, S., Matzel, J., Mundil, R., and Okaya, D., 2010, Magmatic lobes as “snap-
shots” of magma chamber growth and evolution in large, composite batholiths: An example
from the Tuolumne intrusion, Sierra Nevada, California: Bulletin of the Geological Society of
America, v. 122, p. 1912–1931, ht tps:// doi .org /1 0 .1 1 30 /B30 0 04 .1.
Middlemost, E.A., 1994, Naming materials in the magma/igneous rock system: Earth-Science
Reviews, v. 37 , p. 215–224, ht tps:// doi .org /1 0 .1 0 1 6 /0 0 1 2 -8252 (94)90 029 -9.
Paterson, S.R., 2009, Magmatic tubes, pipes, troughs, diapirs, and plumes: Late-stage convec-
tive instabilities resulting in compositional diversity and permeable networks in crystal-rich
magmas of the Tuolumne batholith, Sierra Nevada, California: Geosphere, v. 5, p. 496–527 ,
ht tps:// doi .org /1 0 .1 1 30 /GES0 021 4 .1.
Paterson, S.R., and Ducea, M.N., 2015, Arc magmatic tempos: Gathering the evidence: Elements,
v. 11, p. 91–98, ht tps:// doi .org /1 0 .21 1 3 /gselements .1 1 .2 .91.
Paterson, S.R., Vernon, R.H., and Žák, J., 2005, Mechanical instabilities and physical accumulation
of K-feldspar megacrysts in granitic magma, Tuolumne batholith, California, USA: Journal
of the Virtual Explorer, Electronic Edition, v. 18, 1, ht tps:// doi .org /1 0 .3809 /jvir tex .20 05 .0 0 1 1 4.
Paterson, S.R., Okaya, D., Memeti, V., Economos, R., and Miller, R.B., 2011, Magma addition and
flux calculations of incrementally constructed magma chambers in continental margin arcs:
Combined field, geochronologic, and thermal modeling studies: Geosphere, v. 7 , p. 1439–1468,
ht tps:// doi .org /1 0 .1 1 30 /GES0 0696 .1.
Paterson, S., Memeti, V., Mundil, R., and Žák, J., 2016, Repeated, multiscale, magmatic erosion and
recycling in an upper-crustal pluton: Implications for magma chamber dynamics and magma vol-
ume estimates: American Mineralogist, v. 101, p. 2176–2198, ht tps:// doi .org /1 0 .21 38 /am -20 1 6 -557 6.
Paterson, S.R., Ardill, K., Vernon, R., and Žák, J., 2019, A review of mesoscopic magmatic struc-
tures and their potential for evaluating the hypersolidus evolution of intrusive complexes:
Journal of Structural Geology, v. 125, p. 134–147 , ht tps:// doi .org /1 0 .1 0 1 6 /j .jsg .20 1 8 .04 .022.
Petford, N., 2009, Which effective viscosity?: Mineralogical Magazine, v. 73, p. 167–191, https://
doi .org /1 0 .1 1 80 /minmag .20 09 .073 .2 .1 67 .
Pinotti, L.P ., D’Eramo, F .J., Weinberg, R.F ., Demartis, M., Tubía, J.M., Coniglio, J.E., Radice, S.,
Maffini, M.N., and Aragón, E., 2016, Contrasting magmatic structures between small plutons
and batholiths emplaced at shallow crustal level (Sierras de Córdoba, Argentina): Journal of
Structural Geology, v. 92, p. 46–58, ht tps:// doi .org /1 0 .1 0 1 6 /j .jsg .20 1 6 .09 .0 09.
Ramos, F .C., and Reid, M.R., 2005, Distinguishing melting of heterogeneous mantle sources
from crustal contamination: Insights from Sr isotopes at the phenocryst scale, Pisgah Crater,
California: Journal of Petrology, v. 46, p. 999–1012, ht tps:// doi .org /1 0 .1 093 /petrology /egi0 08.
Reid, J.B., Murray, D.P ., Hermes, O.D., and Steig, E.J., 1993, Fractional crystallization in granites
of the Sierra Nevada: How important is it?: Geology, v. 21, p. 587–590, ht tps:// doi .org /1 0 .1 1 30
/0 091 -7 61 3 (1 993)021 <0587: FCIGO T>2 .3 .CO;2.
Rittmann, A., 1957 , On the serial character of igneous rocks: Egyptian Journal of Geology, v. 1,
p. 23–48.
Rocher, S., Alasino, P .H., Macchioli Grande, M., Larrovere, M.A., and Paterson, S.R., 2018, K-feld-
spar megacryst accumulations formed by mechanical instabilities in chamber margins, Asha
pluton, NW Argentina: Journal of Structural Geology, v. 112, p. 154–173, ht tps:// doi .org /1 0
.1 0 1 6 /j .jsg .20 1 8 .04 .0 1 7 .
Ruprecht, P ., Bergantz, G.W., and Dufek, J., 2008, Modeling of gas-driven magmatic overturn:
Tracking of phenocryst dispersal and gathering during magma mixing: Geochemistry Geo-
physics Geosystems, v. 9, Q07017 , ht tps:// doi .org /1 0 .1 029 /20 08GC0 02022.
Schleicher, J.M., and Bergantz, G.W., 2017 , The mechanics and temporal evolution of an open-sys-
tem magmatic intrusion into a crystal-rich magma: Journal of Petrology, v. 58, p. 1059–1072,
ht tps:// doi .org /1 0 .1 093 /petrology /egx045.
Schleicher, J.M., Bergantz, G.W., Breidenthal, R.E., and Burgisser, A., 2016, Time scales of crystal
mixing in magma mushes: Geophysical Research Letters, v. 43, p. 1543–1550, ht tps:// doi .org
/1 0 .1 0 02 /20 1 5GL067372.
Slemmons, D.B., 1953, Geology of the Sonora Pass region [Ph.D. thesis]: Berkeley, University
of California, 201 p.
Solgadi, F ., and Sawyer, E.W., 2008, Formation of igneous layering in granodiorite by gravity
flow: A field, microstructure and geochemical study of the Tuolumne Intrusive Suite at Saw -
mill Canyon, California: Journal of Petrology, v. 49, p. 2009–2042, ht tps: // d oi .o rg / 1 0 . 1 093
/petrology /egn056.
Steiger, R.H., and Jäger, E., 1977 , Subcommission on Geochronology: Convention on the use
of decay constants in geo- and cosmochronology: Earth and Planetary Science Letters, v. 1,
p. 369–371, ht tps:// doi .org /1 0 .1 0 1 6 /0 0 1 2 -821X (77)90 060 -7 .
Stern, T .W., Bateman, P .C., Morgan, B.A., Newell, M.F ., and Peck, D.L., 1981, Isotopic U-Pb ages of
zircon from the granitoids of the central Sierra Nevada, California: U.S. Geological Survey
Professional Paper 1185, 17 p., ht tps:// doi .org /1 0 .31 33 /pp1 1 85.
Sumita, I., and Manga, M., 2008, Suspension rheology under oscillatory shear and its geophys-
ical implications: Earth and Planetary Science Letters, v. 269, p. 468–477 , ht tps:// doi .org /1 0
.1 0 1 6 /j .epsl .20 08 .02 .043.
Sun, S.-s., and McDonough, W.F ., 1989, Chemical and isotopic systematics of oceanic basalts:
Implications for mantle composition and processes, in Saunders, A.D., and Norry, M.J., eds.,
Magmatism in the Ocean Basins: Geological Society of London Special Paper 42, p. 313–345,
ht tps:// doi .org /1 0 .1 1 44 /GSL .SP .1 989 .042 .0 1 .1 9.
T anaka, T ., T ogashi, S., Kamioka, H., Amakawa, H., Kagami, H., Hamamoto, T ., Yuhara, M., Orihashi,
Y., Yoneda, S., Shimizu, H., Kunimaru, T ., Takahashi, K., Yanagi, T ., Nakano, T ., Fujimaki, H.,
Shinjo, R., Asahara, Y ., Tanimizu, M., and Dragusanu, C., 2000, JNdi-1: A neodymium isoto-
pic reference in consistency with La Jolla neodymium: Chemical Geology, v. 168, p. 279–281,
ht tps:// doi .org /1 0 .1 0 1 6 /S0 0 09 -2541 (0 0)0 0 1 98 -4.
Vernon, R.H., and Paterson, S.R., 2008, Mesoscopic structures resulting from crystal accumula-
tion and melt movement in granites: Transactions of the Royal Society of Edinburgh: Earth
Sciences, v. 97 , p. 369–381, ht tps:// doi .org /1 0 .1 0 1 7 /S026359330 0 0 0 1 51 6.
Wahrhaftig, C., 1979, Significance of asymmetric schlieren for crystallization of granites in the
Sierra Nevada batholith, California: Geological Society of America Abstracts with Programs,
v. 11, p. 133.
Wahrhaftig, C., 2000, Geological map of the Tower Peak quadrangle, central Sierra Nevada, Cali-
fornia: U.S. Geological Survey Miscellaneous Investigation Series Map I-2697 , scale 1:62,500.
Wallace, G.S., and Bergantz, G.W., 2005, Reconciling heterogeneity in crystal zoning data: An
application of shared characteristic diagrams at Chaos Crags, Lassen Volcanic Center, Cali-
fornia: Contributions to Mineralogy and Petrology, v. 149, p. 98–112, ht tps:// doi .org /1 0 .1 0 07
/s0 041 0 -0 04 -0639 -2.
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
251
26
Alasino et al. | Magmatically folded and faulted schlieren zones in the Sonora Pass Intrusive Suite GEOSPHERE | V olume 15 | Number X
Research Paper
Weinberg, R.F ., Sial, R.N., and Pessoa, R.R., 2001, Magma flow within the Tavares pluton, north -
western Brazil: Compositional and thermal convection: Geological Society of America Bulletin,
v. 113, p. 508–520, ht tps:// doi .org /1 0 .1 1 30 /0 0 1 6 -7 606 (20 0 1)1 1 3 <0508: MFWTTP>2 .0 .CO;2.
Žák, J., and Klomínský, J., 2007 , Magmatic structures in the Krkonoše–Jizera Plutonic Complex,
Bohemian Massif: Evidence for localized multiphase flow and small-scale thermal-mechanical
instabilities in a granitic magma chamber: Journal of Volcanology and Geothermal Research,
v. 164, p. 254–267 , ht tps:// doi .org /1 0 .1 0 1 6 /j .jv olgeores .20 07 .05 .0 06.
Žák, J., and Paterson, S.R., 2005, Characteristics of internal contacts in the Tuolumne Batholith,
central Sierra Nevada, California (USA): Implications for episodic emplacement and physical
processes in a continental arc magma chamber: Geological Society of America Bulletin, v. 1 17 ,
p. 1242–1255, ht tps:// doi .org /1 0 .1 1 30 /B25558 .1.
Žák, J., and Paterson, S.R., 2010, Magmatic erosion of the solidification front during reintrusion:
The eastern margin of the Tuolumne batholith, Sierra Nevada, California: International Journal
of Earth Sciences, v. 99, p. 801–812, ht tps:// doi .org /1 0 .1 0 07 /s0 0531 -0 09 -0423 -7 .
Downloaded from https://pubs.geoscienceworld.org/gsa/geosphere/article-pdf/doi/10.1130/GES02070.1/4811976/ges02070.pdf
by University of Southern California user
on 16 August 2019
252
MostGranitoidRocksareCumulates:
DeductionsfromHornblendeCompositions
andZirconSaturation
CalvinG.Barnes
1
*,KevinWerts
1
,ValiMemeti
2
andKatieArdill
3
1
Department of Geosciences, Texas Tech University, Lubbock, TX 79409, USA;
2
Department of Geological
Sciences, California State University, Fullerton, Fullerton, CA 92834, USA;
3
Department of Earth Sciences,
University of Southern California, Los Angeles, CA 90089, US
*Corresponding author. Telephone: 1-806-834-7389. Fax: 1-806-742-0100. E-mail: cal.barnes@ttu.edu
Received April 2, 2019; Accepted February 6, 2020
ABSTRACT
Cumulate processes in granitic magma systems are thought by some to be negligible and by
others to be common and widespread. Because most granitic rocks lack obvious evidence of accu-
mulation, such as modal layering, other means of identifying cumulate rocks and estimating pro-
portions of melt lost must be developed. The approach presented here utilizes major and trace
element compositions of hornblende to estimate melt compositions necessary for zircon satur-
ation. It then compares these estimates with bulk-rock compositions to estimate proportions of
extracted melt. Data from three arc-related magmatic systems were used (English Peak pluton,
Wooley Creek batholith, and Tuolumne Intrusive Complex). In all three systems, magmatic horn-
blende displays core-to-rim decreases in Zr, Hf, and Zr/Hf. This zoning indicates that zircon must
have fractionated during crystallization of hornblende, at temperatures greater than 800
C. This T
estimate is in agreement with Ti-in-zircon thermometry, which yields a maximum T estimate of
855
C. On the basis of this evidence, concentrations of Zr in melts from which hornblende and zir-
con crystallized were calculated by (1) applying saturation equations to bulk-rock compositions, (2)
applying saturation equations to calculated melt compositions, and (3) using hornblende/melt par-
tition coefficients for Zr. The results indicate that melt was lost during crystallization of the granitic
magmas, conservatively at least as much as 40 %. These results are in agreement with published
estimates of melt loss from other plutonic systems and suggest that bulk-rock compositions of
many granitic rocks reflect crystal accumulation and are therefore inappropriate for use in thermo-
dynamic calculations and in direct comparison of potentially consanguineous volcanic and plutonic
suites.
Keywords: zircon saturation; crystal accumulation; hornblende trace elements; melt composition
INTRODUCTION
Development of the zircon saturation thermometer
(Watson & Harrison, 1983) provided a means by which
conditions of crystallization of granitic rocks could be
quantified, without reliance on exchange thermometers
(e.g. two-feldspar and Fe–Ti oxide thermometers;
Andersen & Lindsley, 1988; Fuhrman & Lindsley, 1988),
which tend to be reset during slow cooling (e.g.
Anderson, 1996; Andersonetal., 2008). However, appli-
cation of the Watson & Harrison (1983) algorithm, and
revised equations (Boehnke et al., 2013; Gervasoni
et al., 2016; Borisov & Aranovich, 2019), to arc-related
plutonic rocks tended to yield temperatures at or near
the H
2
O-saturated solidus, which suggests that zircon
fractionation could not substantially affect melt compo-
sitions. In contrast, among the most siliceous parts of
arc plutonic suites, Zr concentrations typically decrease
with increasing SiO
2
(Fig. 1); this variation is generally
interpreted as fractionation of zircon from a more mafic
magma (e.g. Miller et al., 2003; Deering & Bachmann,
V C The Author(s) 2020. Published by Oxford University Press. All rights reserved. For permissions, please e-mail: journals.permissions@oup.com 1
JOURNAL OF
PETROLOGY
Journal of Petrology, 2020, 1–14
doi: 10.1093/petrology/egaa008
Advance Access Publication Date: 11 March 2020
Original Article
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
253
2010; Barnes et al., 2016a) or as residual zircon in the
source. In the case of high-T magmas, Harrison et al.
(2007) explained this discrepancy as the effect of crys-
tallization driving melt composition toward zircon sta-
bility at T higher than expected if the magma’s bulk
composition controlled zircon saturation. This explan-
ation is supported by the fact that in many granitic sys-
tems, Ti-in-zircon thermometry (Ferry & Watson, 2007)
yields higher T estimates of zircon crystallization than
saturation thermometers (e.g. Miller et al., 2003, 2007;
Claiborne et al., 2006; Harrison et al., 2007; review by
Sie ´geletal., 2018).
The observation of decreasing Zr with increasing
SiO
2
suggests that zircon saturation occurred at sub-
stantially higher temperature (T) than the wet solidus
(Harrison et al., 2007). If the low temperatures of zircon
saturation are correct, they lead to assumptions con-
cerning the thermal and mechanical state of the mag-
mas; for example, magmatic T below critical lock-up
conditions and therefore magmas that are not erupt-
able. Likewise, if zircon was not saturated in the melt
until near-solidus temperatures, then presumably ante-
crystic and inherited zircons should dissolve at higher
magmatic temperatures (e.g. Miller et al., 2003, 2007).
Because antecrystic and inherited zircons are present in
many arc magmas, their presence indicates either that
zircon was saturated at higher temperatures than indi-
cated by saturation thermometry, or that these zircons
survived because of large size, or were armored by
other minerals (Miller et al., 2003, 2007; Sie ´gel et al.,
2018).
The differences in temperatures obtained from zircon
saturation T and Ti-in-zircon thermometry (see review
by Sie ´gel et al. 2018) may have a number of causes, as
follows. (1) In high-T magmas undergoing fractionation,
zircon should precipitate atT higher than the calculated
saturationT (Harrisonetal., 2007). (2) Saturation therm-
ometry may be applied to rock compositions outside
the calibration range of the experiments (see Watson &
Harrison, 1983; Hanchar & Watson, 2003). (3) For
crystal-rich volcanic rocks and particularly for coarse-
grained plutonic rocks, the bulk composition used in zir-
con saturation equations may not represent a melt com-
position but instead may represent a cumulate (Deering
& Bachmann, 2010; Lee & Morton, 2015; Barnes et al.,
2016a). It is the last possibility that we explore in this
paper.
Crystal accumulation in mafic plutonic systems is
widely accepted, in part because of field evidence such
as modal layering, and compositional evidence such as
cryptic layering. In contrast, granitic plutons (sensu
lato) generally lack widespread modal layering and may
or may not display widely variable mineral composi-
tions. Moreover, compositional trends in many plutonic
systems are tightly constrained and may be interpreted
as liquid lines of descent. However, linear composition-
al trends may equally well illustrate magma mixing or
accumulation of crystals in cotectic proportions.
Determining whether crystal accumulation occurred in
50
100
150
200
250
300
Zr (ppm)
50
50
100
150
150
200
250
250
300
0
100
200
300
40 50 60 70 80
SiO (wt%)
2
Heiney Bar pluton
English Peak pluton
late-stage English Peak pluton
border unit
Yellowjacket Ridge unit
Chimney Rock unit
central zone
lower zone
upper zone
late-stage intrusion
Wooley Creek batholith
Kuna Crest
Tuolumne complex
equigranular Half Dome
porphyritic Half Dome
Cathedral Peak
Johnson Granite
southern Half Dome lobe
M from 0.9 to 2.0
M from 0.9 to 2.0
M from 0.9 to 2.0
potential zircon
cumulates
potential zircon
cumulates
potential zircon
cumulates
T range (°C)
(B et al.) 617–725
(W & H) 678–788
T range (°C)
(B et al.) 655–760
(only two > 715)
(W & H) 725–785
T range (°C)
(B et al.) 660–726
(W & H) 720–775
Fig. 1. Variation of bulk-rock Zr and SiO
2
for all rock types in
each plutonic suite. Samples used in this study are outlined
and are those that meet the criterion for M. The ranges of zir-
con saturation temperatures for samples with M 2 are listed
in each panel, utilizing the Boehnke et al. (2013; B et al.) and
Watson & Harrison (1983; W & H) equations. The former equa-
tion yields temperature estimates near or below the commonly
accepted H
2
O-saturated solidus for tonalitic–granitic magmas,
whereas the latter algorithm yields somewhat higher tempera-
ture estimates.
2 JournalofPetrology, 2020, Vol. 0, No. 0
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
254
granitic systems has a number of important implica-
tions. For example, if the majority of granitic rocks es-
sentially represent melts (e.g. Coleman et al., 2012),
then compositional data may be used to directly assess
liquid lines of descent and evaluate relationships be-
tween magma batches. In addition, if rock compositions
equal melt compositions, then comparisons between
granitic and felsic volcanic rocks will directly address
the potential connection between plutonic and volcanic
rocks (e.g. Glazner et al., 2015). However, if granitic
rocks are cumulate and do not reflect melt composi-
tions, then they are potentially inappropriate as input
for thermodynamic models such as RhyoliteMELTS
(Gualda et al., 2012) and may not be directly compar-
able with felsic volcanic rocks (Keller et al., 2015). We
begin with a review of previous studies in which crystal
accumulation was cited as an important process in the
evolution of granitic (sensulato) magmas. We then pre-
sent one approach for testing accumulation and esti-
mating proportions of lost melt.
Early examples of potential crystal accumulation
focused on concentric compositional zoning in plutons
of various sizes. This zoning was explained in terms of
side-wall crystal nucleation and accumulation (e.g.
Bateman & Nokleberg, 1978; Bateman & Chappell,
1979; McCarthy & Groves, 1979; Srogi & Lutz, 1996,
1997; Chappell et al., 1998). Recognition that some plu-
tons are tilted (e.g. Flood & Shaw, 1979; Barnes, 1983)
led to recognition of vertical zoning and the likelihood
of upward melt migration (e.g. Flood & Shaw, 1979;
Wiebe, 1996; Wiebe & Collins, 1998; Bachl et al., 2001;
Miller & Miller, 2002; Harper et al., 2004; Walker et al.,
2007; Turnbull et al., 2010), as did recognition of leuco-
cratic granitic caps and cupolas above granitic to grano-
dioritic plutons (e.g. Mahood & Cornejo, 1992; Barnes
et al., 2001). A variety of evidence for crystal accumula-
tion is cited, and includes textures (e.g. Wiebe &
Collins, 1998; Bachl et al., 2001; Harper et al., 2004;
Turnbull et al., 2010), magmatic structures (e.g. Wiebe
& Collins, 1998; Weinberg, 2006; Vernon & Paterson,
2008; Paterson, 2009; Wiebe et al., 2017), apparent dis-
equilibrium between mineral and bulk-rock element
ratios (e.g. Barnesetal., 2016a; Wertsetal., 2018, 2020),
and geochemical variation (e.g. McCarthy & Groves,
1979; Chappellet al., 1998; Miller & Miller, 2002; Harper
et al., 2004; Claiborne et al., 2006; Walker et al., 2007;
Deering & Bachmann, 2010; Lee & Morton, 2015;
Barnesetal., 2016a; Schaenetal., 2017, 2018).
In some cases in which geochemical evidence is
used to support crystal accumulation, the cumulate
rocks were found to be enriched in Zr compared with
likely parental compositions, which indicates that zircon
was a cumulate phase (e.g. Claiborne et al., 2006;
Turnbull, 2010; Schaen et al., 2017, 2018; Ratschbacher
et al., 2018). However, in other cases, including exam-
ples discussed below, plutons that display geochemical
or textural evidence for crystal accumulation are char-
acterized by low Zr contents, despite the fact that bulk-
rock Zr contents decrease as SiO
2
increases. In such
cases, calculated melt Zr contents needed for zircon sat-
uration may be much higher than bulk-rock Zr contents
(see, e.g. Milleretal., 2007).
Although many researchers recognized the presence
and role of crystal cumulates in granitic systems, there
are relatively few examples in which the proportions of
accumulation or melt loss have been estimated. In part,
this lack arises from the fact that many granitic magmas
evolve along a multiply saturated cotectic, making iden-
tification of parental, cumulate, and fractionated com-
positions difficult (e.g. Deering & Bachmann, 2010).
Nevertheless, Srogi & Lutz (1996, 1997) proposed a
method based on major element compositions that
yielded estimates of melt loss from 10 to 60 %. Other
workers relied primarily on variations in trace element
abundances and ratios to estimate melt loss. For ex-
ample, Bachl et al. (2001) inferred significant propor-
tions of melt loss by comparing Rb, Sr, Ba, and rare
earth elements (REE) between cumulates and a calcu-
lated parental composition. Similarly, Deering &
Bachmann (2010), Lee & Morton (2015) and Schaen
etal. (2017) used trace element models to estimate per-
centages of extracted melt of 20–40 %, 28–40 %, and c.
50 %, respectively.
In this paper, we pursue a different approach to iden-
tify cumulate granitic rocks and estimate percentages of
melt lost. In particular, we utilize trace and major elem-
ent zoning trends in calcic amphibole (hereafter horn-
blende: Hbl) and temperature estimates based on Hbl
compositions to estimate the temperatures and melt
compositions at which zircon began to crystallize. We
then compare Zr contents necessary for zircon satur-
ation with bulk-rock compositions to estimate the pro-
portion of melt lost.
SUMMARYOFGEOLOGICALSETTING
We utilize bulk-rock and mineral compositions from
three arc-related plutonic complexes in the North
American Cordillera. The English Peak pluton and
Wooley Creek batholith crop out in the Klamath
Mountain province (northern California) and the third,
the Cretaceous Tuolumne Intrusive Complex, is
exposed in the Sierra Nevada of California. The geo-
logical setting, age, and compositional variation of each
of these systems are well documented: English Peak by
Barnes et al. (2016b, 2017) and Ernst et al. (2016);
Wooley Creek by Barnes (1983), Barnes et al. (1986,
1990, 2016a) and Coint et al. (2013b); Tuolumne
Intrusive Complex by, for example, Bateman &
Chappell (1979), Kistler et al. (1986), Huber et al. (1989)
and Memetietal. (2010b, 2014).
The two Klamath Mountain plutons are Middle to
Late Jurassic in age (Coint et al. 2013b; Ernst et al.,
2016). The English Peak pluton is a composite pluton
that consists of two small satellite bodies and a large
central mass, which is subdivided into a gabbroic to
tonalitic early stage and a tonalitic to granitic late stage
(Barnes et al., 2016b, 2017; Ernst et al., 2016;
JournalofPetrology, 2020, Vol. 0, No. 0 3
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
255
Supplementary Fig. 1; supplementary data are available
for downloading at http://www.petrology.oxfordjour
nals.org). Our focus in this study is on tonalitic to granit-
ic samples from late-stage border unit, Yellowjacket
Ridge unit, and Chimney Rock unit, with additional data
from felsic components of one satellite body, the
Heiney Bar pluton.
The Wooley Creek batholith (Barnes, 1983; Barnes
et al., 1986, 1990, 2016a; Coint et al., 2013b;
Supplementary Fig. 2) is also composite and is tilted
such that older gabbroic to tonalitic rocks are exposed
in the deeper, northeastern part of the batholith (lower
zone) and younger tonalitic to granitic rocks are
exposed in the southwestern part (upper zone). Our
focus is on tonalitic to granitic rocks in the upper zone.
The Tuolumne Intrusive Complex (Supplementary
Fig. 3) is one of the large (1100 km
2
), nested, Late
Cretaceous batholiths exposed along the crest of the
Sierra Nevada (e.g. Bateman & Chappell, 1979; Kistler
et al., 1986; Huber et al., 1989; Bateman, 1992; Memeti
et al., 2010b, 2014). It was emplaced over 10 Myr, be-
tween 95 and 85 Ma (Coleman et al., 2004; Paterson
et al., 2016) at depths of 6–10 km (Ague & Brimhall,
1988). It is broadly inwardly zoned in age and compos-
ition, from the outer Kuna Crest unit inward to equigra-
nular Half Dome, porphyritic Half Dome, Cathedral
Peak, and Johnson Granite Porphyry units (Bateman &
Chappell, 1979; Kistleretal., 1986; Colemanetal., 2004;
Memeti et al., 2010b; Paterson et al., 2016). The Kuna
Crest unit is the most compositionally variable, ranging
from gabbro to granite and minor leucogranite. Half
Dome and Cathedral Peak units are mainly granodiorite
with lesser tonalite and granite.
Samples used in this study contain calcic amphibole
and biotite as ferromagnesian silicates. The calcic
amphibole is mainly magnesio-hornblende (hereafter
Hbl), with lesser amounts of tschermakite. In all three
systems, Hbl coexists with intermediate to sodic plagio-
clase, quartz, biotite, Fe–Ti oxides, K-feldspar, and scant
allanite. K-feldspar is mainly interstitial to poikilitic in
the English Peak pluton and Wooley Creek batholith,
but occurs as phenocrysts–megacrysts in interior units
of the Tuolumne complex (e.g. Bateman & Chappell,
1979). Magmatic titanite is widespread in the Tuolumne
complex but is absent in both Klamath plutons except
for a small, late-stage body in the southern Wooley
Creek batholith. Accessory Fe–Ti oxides are relatively
abundant in the Tuolumne Intrusive Complex, in which
magnetite> ilmenite. In contrast, Fe–Ti oxides are scant
in the English Peak and Wooley Creek plutons, and il-
menite> magnetite.
In addition, in the Discussion section, we compare
bulk compositions of the plutonic rocks with published
data on volcanic rocks, including Cretaceous volcanic
rocks exposed within roof pendants in the central Sierra
Nevada (375–38
N). These Sierran data include sam-
ples from the Saddlebag Lake, Piute, Cinko Lake, Iron
Mountain, and Ritter Range pendants (e.g. Fiske &
Tobisch, 1994; Greene & Schweickert, 1995; Memeti
et al., 2010b; Paterson & Memeti, 2014). These volcanic
packages are intruded by voluminous upper-crustal plu-
tons, which are in some cases broadly coeval with
exposed volcanic rocks (e.g. Andersonetal., 2008).
The dominant rock types are rhyolite and dacite clas-
tic- to phenocryst-bearing tuffs (Paterson & Memeti,
2014). Andesitic volcanic rocks are relatively less com-
mon, and basalts are rare. Fiske & Tobisch (1994)
reported kilometer-thick rhyolite ash-flow tuffs and a
megabreccia unit containing blocks of tuff, lavas and
tuff breccias, that together form the caldera fill se-
quence of the Minarets caldera. The Merced Peak vol-
canic complex is characterized by rhyolitic to dacitic
ash-flow tuffs and lavas and has a wider compositional
range than the Minarets caldera tuff units (Peck, 1980;
Lowe, 1995). Feldspar phenocrysts are abundant, to-
gether with lesser amounts of quartz and biotite in felsic
samples. Variably altered Hbl phenocrysts are present
in Beartrap Lake and a few Minarets caldera samples.
METHODSANDDATASOURCES
New bulk-rock data were obtained by X-ray fluores-
cence (Tuolumne and English Peak) and by inductively
coupled plasma atomic emission spectrometry (Wooley
Creek). Major element compositions of Hbl were deter-
mined by electron microprobe (EMP) analysis in various
laboratories over the past 30 years (Werts et al., 2020).
Typical operating conditions were 15–20 kV accelerating
voltage, 20 nA beam current, andc.1lm spot size using
natural and synthetic standards. All Hbl trace element
abundances were determined by laser ablation induct-
ively coupled plasma mass spectrometry (LA-ICP-MS)
at Texas Tech University using a NewWave 213 nm
solid-state laser and Agilent 7500CS quadrupole ICP-
MS system. Nominal conditions were 40lm spot diam-
eter and laser pulse rate of 5 Hz. Trace element abun-
dances were normalized to that of CaO. Precision,
determined by repeated analysis of basaltic glass
BHVO-2g, is 21–92 %, and <6 % for most trace ele-
ments. Additional details have been provided by Barnes
et al. (2017) and Werts et al. (2020). For English Peak
and Tuolumne samples, analytical spots were the same
for EMP and LA-ICP-MS analyses. However, owing to
the ‘legacy’ nature of many EMP analyses for Wooley
Creek Hbl, many EMP and LA-ICP-MS spots are not
fiducial.
Zircon separates from the English Peak pluton and
Wooley Creek batholith were analyzed for trace ele-
ments by LA-ICP-MS at Texas Tech. Analytical condi-
tions were 8–9 J cm
–2
, 5 Hz repetition rate, and 30lm
spot size using NIST 610 as the calibration standard and
zircons 91500 and Temora 2 as internal standards.
Hornblende and bulk-rock data for the English Peak
pluton were published by Barnes et al. (2016b, 2017)
and for the Wooley Creek batholith by Coint et al.
(2013a, 2013b). Bulk-rock data for the Tuolumne com-
plex come from Bateman & Chappell (1979), Burgess &
Miller (2008), Gray et al. (2008), Economos et al. (2009)
4 JournalofPetrology, 2020, Vol. 0, No. 0
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
256
and Coleman et al. (2012), supplemented by unpub-
lished data of the authors. Tuolumne amphibole com-
positions are from Barnes et al. (2016c) and Werts et al.
(2020).
Data for volcanic rocks compiled here include sam-
ples from the Minarets caldera (Fiske & Tobisch, 1994;
Lowe, 1995) and the Merced Peak volcanic–plutonic
complex (Peck & van Kooten, 1983). Samples from the
Minarets caldera yielded ages between 98 and 101 Ma
(Fiske & Tobisch, 1994; U–Pb in zircon). Two samples
from the Merced Peak complex yielded ages c.95Ma
(Lowe, 1995; U–Pb in zircon). Volcanic rocks in the
Saddlebag Lake pendant at Sawmill Canyon yielded
ages of 95–106 Ma (Paterson & Memeti, 2014; Ardill
etal., 2018), whereas at Virginia Canyon, volcanic rocks
interbedded with sediments yielded an age of 117 Ma
(Caoetal., 2015). Cretaceous volcanic rocks to the north
of the Tuolumne Intrusive Complex, in the Cinko Lake
pendant, range in age from 99 to 107 Ma (Memeti et al.,
2010a).
Thermometry
The zircon saturation thermometer strictly applies to
samples in which the parameter M ¼ molar
(Naþ Kþ 2Ca)/(AlSi) is between 09 and 19(Watson &
Harrison, 1983). To use somewhat more data from the
English Peak and Wooley Creek systems, we expanded
this range slightly, such that samples used in this study
haveM values from 09to20, which encompasses rock
compositions from tonalite to granite and incorporates
26 analyses from English Peak, 19 from Wooley Creek,
and >200 from the Tuolumne complex. Temperatures
of Hbl crystallization were calculated using the
pressure-independent thermometer of Putirka [2016;
equation (5)], which is a function of Si, Ti, total Fe, and
Na in the amphibole. Ti-in-zircon temperatures (Ferry &
Watson, 2007) were calibrated for conditions in which
a(SiO
2
) ¼ a(TiO
2
) ¼ 1. All samples are quartz-bearing,
so that use of a(SiO
2
) ¼ 1 is probably appropriate.
However, the Ti-bearing phases are titanite and mag-
netite6 ilmenite in the Tuolumne complex and ilmenite
þ magnetite in the English Peak and Wooley Creek plu-
tons. The calculations therefore used a(TiO
2
) ¼ 05 for
the English Peak and Wooley Creek plutons (Schiller &
Finger, 2019). By comparison, published Ti-in-zircon
temperatures for titanite-bearing samples from the
Tuolumne complex used a(TiO
2
) ¼ 075 [Matzel et al.,
2007; compare Claiborne et al., 2006, who deduced an
a(TiO
2
)of07 for titanite-bearing samples].
DISCUSSION
Zirconfractionationandmineralzoning
A number of recent studies have shown that composi-
tions of calcic amphibole may be used to estimate crys-
tallization T and compositions of the equilibrium melt
phase (e.g. Ridolfi et al., 2010; Ridolfi & Renzulli, 2012;
Putirka, 2016; Zhang et al., 2017; Humphreys et al.,
2019). Trace element zoning patterns in Hbl have also
been shown to provide a record of changing melt com-
positions during crystallization (Barnes et al., 2016c,
2017; Werts et al., 2018, 2020). Therefore, combined
major and trace element analyses of Hbl may be
used to determine core-to-rim change in T and melt
composition during crystallization (Barnes et al.,
2016b, 2017).
In the samples we studied, we found that Ti in Hbl is
closely correlated with T, which is consistent with em-
pirical (Otten, 1984) and experimental (Ernst & Liu,
1998) studies, and is also consistent with the import-
ance of Ti in the Putirka (2016) Hbl thermometer. We
therefore use decreasing Ti content as a proxy for
decreasing temperature of Hbl crystallization (also see
Barnes et al., 2017). Specifically, we use Ti as deter-
mined from LA-ICP-MS analysis because it correlates
directly with trace element abundances from the same
analytical volume. In general, the highest Ti contents
were measured in crystal interiors and the lowest in
crystal rims. In each plutonic suite, the abundances of
Zr and Hf are positively correlated with Ti and decrease
from crystal interiors to margins (Fig. 2).
At the calculated temperatures of Hbl core crystal-
lization (c. 860–820
C), experimentally determined par-
tition coefficients (Nandedkar et al., 2016) for Zr range
from 07to 13 and for Hf from 12to 17. However,
most other studies yielded lower partition coefficients
(034–068 for Zr; 082–095 for Hf; Sisson, 1994; Klein
et al., 1997; Bachmann et al., 2005; Padilla & Gualda,
2016). Thus, even if the highest partition coefficients
(Nandedkar et al., 2016) are used, ones in which Zr and
Hf maybeslightlycompatiblein Hbl, acompletefrac-
tionating assemblage would involve plagioclase> Hbl
6 biotite6 quartz6 Fe–Ti oxides. The partition coeffi-
cients of Zr and Hf in the other phases are nearly zero,
which will result in bulk partition coefficients <10.
Thus, both Zr and Hf should display mildly incompat-
ible behavior. If this were the case, both elements
should increase in abundance from crystal interiors
torims andbe anticorrelatedwithTi, contrary towhat
is observed.
To explain the observed Zr and Hf zoning, fraction-
ation of a Zr- and Hf-rich phase must be called on.
Furthermore, in each system studied, we found that the
Zr/Hf ratio decreases from core to rim (Fig. 2). Zircon is
the only common phase capable of causing this de-
crease (Bea et al., 2006). Effects of zircon fractionation
(lowered Zr/Hf) and accumulation (increased Zr/Hf)
have been documented in plutonic (e.g. Claiborneetal.,
2006) and volcanic (e.g. Reidetal., 2010) systems. Thus,
although individual zoning patterns of Zr and Hf in Hbl
are strong indicators of zircon fractionation, we suggest
that the onset of decreasing Zr/Hf ratios is even stronger
evidence, and therefore conclude that in each of these
plutonic systems, Hbl and zircon co-precipitated.
If the preceding conclusion is correct, then it is pos-
sible to estimate temperatures at which Hbl þ zircon
crystallized. We refer to these temperatures as
JournalofPetrology, 2020, Vol. 0, No. 0 5
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
257
estimates, because each sample displays slightly differ-
ent trends with regard to Zr, Hf, and Ti behavior (e.g.
Fig. 2a). If decreasing Zr is assumed to characterize zir-
con fractionation, then the maximum hornblende tem-
peratures obtained may be used as (minimum)
indicators of zircon stability. These T values are c.
890
C for the English Peak pluton, c. 800
C for the
Wooley Creek batholith, and c. 825
C for the Tuolumne
complex (Fig. 2). A more conservative approach is to
identify the temperature at which the Zr/Hf ratios begin
to decrease. This T varies from 780 to 890
C for the
English Peak pluton, depending upon the sample. We
chose 820
C as typical of the point at which Zr/Hf
decreases in the majority of samples. For the other two
systems, estimated T at which Zr/Hf decreases is the
same as the T at which Zr decreases: c. 800
C for the
Wooley Creek batholith, and c. 825
C for the Tuolumne
complex (Fig. 2). These temperatures are all higher than
zircon saturation temperatures (Fig. 1).
If these temperature estimates are correct, then they
should agree, within uncertainty, with temperatures cal-
culated using the Ti-in-zircon thermometer (Ferry &
Watson, 2007). Trace element analyses of zircons used
for dating late-stage rocks of the English Peak pluton
and upper zone rocks of the Wooley Creek batholith
(see Coint et al., 2013b; Ernst et al., 2016) yield Ti con-
tents of 48–143 ppm and 25–136 ppm, respectively.
Both plutons contain scant ilmenite and magnetite and
40
80
2
4
6
8
30
35
20
25
Hf (ppm)
Zr/Hf
120
160
200
12000 14000 16000
0
5
10
15
2000 4000 6000 8000 10000
Ti (ppm)
Zr (ppm)
English Pk pluton
late stage
border unit
Yellowjacket Ridge unit
Chimney Rock unit
~820°C ~780°C ~890°C
2
4
0
5
10
15
20
25
30
35
2000 4000 6000 8000 10000 12000 14000 16000
Ti (ppm)
Zr/Hf
upper Wooley Creek batholith
Hbl-bearing roof dikes
late-stage intrusion
40
80
120
160
200
6
8
Zr (ppm)
Hf (ppm)
~800°C
20
25
30
35
2
4
6
0
5
10
15
Zr/Hf
2000 4000 6000 8000 10000 12000 14000 16000
Ti (ppm)
20
40
60
80
100
Zr (ppm)
8
Hf (ppm)
Kuna Crest
Tuolumne complex
equigranular Half Dome
porphyritic Half Dome
Cathedral Peak
Johnson Granite
southern Half Dome lobe
~825°C
~850°C
~790°C
AB C
Fig.2. Variation of Zr, Hf, and Zr/Hf in hornblende as a function of Ti. Error bars represent 1r uncertainties. On a sample-by-sample
basis, Zr and Hf abundances decrease monotonically with decreasing Ti. In contrast, the Ti value at which Zr/Hf ratios decrease
varies from sample to sample. Vertical blue lines indicate approximate temperature, as calculated from Putirka (2016). (a) English
Peak pluton; (b) Wooley Creek batholith; (c) Tuolumne Intrusive Complex.
6 JournalofPetrology, 2020, Vol. 0, No. 0
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
258
neither contains magmatic titanite; we therefore used
activities of SiO
2
and TiO
2
of 10 and 05, respectively
(see Methods section). Resulting temperatures range
from 766 to 855
C in two samples from the late-stage
English Peak and from 686 to 797
C in seven samples
from the upper Wooley Creek batholith. Thus, for both
of these plutons, Ti-in-zircon T overlaps T estimates for
zircon fractionation made on the basis of amphibole
zoning.
Ti-in-zircon temperatures of 780–640
C were
reported on two samples from the Tuolumne complex
(Matzeletal., 2007), somewhat lower thanT determined
for zircon fractionation on the basis of Hbl zoning (c.
825
C). These calculations used TiO
2
activity of 075
owing to the presence of magmatic titanite. It is unclear
whether analysis of additional samples would expand
the range of Ti-in-zircon temperatures.
Zrcontentsofmelts
The trends of decreasing bulk-rock Zr and hornblende
Zr/Hf in all three systems (Figs 1 and 3) clearly indicate
that zircon was a fractionating phase. Therefore, if one
assumes that bulk-rock compositions represent the
melt compositions from which zircon crystallized, the
necessary Zr content for zircon stability may be calcu-
lated for a given temperature. For this calculation, we
use the approximate temperature at which Zr/Hf
decreases in amphibole. The results are plotted against
bulk-rock Zr contents in Fig. 3. For the Boehnke et al.
(2013) zircon saturation algorithm, the amount of Zr ne-
cessary to saturate zircon is tens to hundreds of ppm
higher than present in the rocks hosting the zircon. The
algorithm of Watson & Harrison (1983) results in a
smaller discrepancy between saturation concentrations
and bulk-rock contents; nevertheless, all but one sam-
ple from the Wooley Creek batholith require higher Zr
for zircon saturation than is present in the rock. [It
should be noted that use of the Gervasoni et al. (2016)
saturation equation results in melt Zr contents higher
than are determined using the Boehnke et al. (2013)
equation.]
The differences between measured bulk-rock Zr and
the amounts of Zr necessary to saturate melts assumed
to have bulk-rock compositions are difficult to explain in
a closed-system setting. Some of these differences may
be explained by fractional crystallization (see Harrison
et al., 2007). However, for many samples this approach
requires at least 50 % crystallization (Watson & Harrison
algorithm) and as much as 75 % crystallization
(Boehnke et al. algorithm). Such extreme degrees of
fractionation are unlikely considering the textural equi-
librium displayed among plagioclase and Hbl and the
fact that Hbl and zircon coprecipitated. Instead, we sug-
gest that (1) the Zr contents of melts in these systems
were necessarily higher than in bulk-rock samples, and
(2) some portion of these higher-Zr melts was lost dur-
ing crystallization, such that bulk-rock compositions
reflect a magma that consisted of cumulate phases and
the remaining interstitial melt.
300
400
500
600
700
800
900
Zr (ppm) for saturation at 820°C
Zr (bulk-rock, ppm)
Heiney Bar pluton
English Peak pluton
1 : 1
1 : 1
1 : 1
late-stage English Peak pluton
border unit
Yellowjacket Ridge unit
Chimney Rock unit
Boehnke et al.
Boehnke et al.
Watson & Harrison
Watson & Harrison
200
300
400
500
600
700
800
Zr (ppm) for satn 800°C
central zone
upper zone
late-stage intrusion
roof-zone dikes
intra-pluton felsic dikes
Wooley Creek batholith
200
300
400
500
600
700
Zr (ppm) for satn at 825°C
0 50 100 150 200 250
Boehnke et al.
Watson & Harrison
b
b
Kuna Crest
Tuolumne complex
equigranular Half Dome
porphyritic Half Dome
Cathedral Peak
Johnson Granite
southern Half
Dome lobe
Fig. 3. Zr saturation temperatures calculated assuming bulk-
rock compositions are equivalent to melt compositions for
samples with 20M 09. Enclosed fields identify results
from the Watson & Harrison (1983) and Boehnke et al. (2013)
equations.
JournalofPetrology, 2020, Vol. 0, No. 0 7
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
259
If Zr-rich melts were indeed lost from these systems,
an obvious question follows: what were the composi-
tions of those melts, and in particular, what were their
Zr contents? We used two approaches to calculate Zr
contents in such melts. In the first, the chemographic
equations of Zhang et al. (2017) were used to calculate
major oxide melt compositions in equilibrium with the
highest-T Hbl. Zhang et al. (2017) did not develop an
equation to estimate Na
2
O contents, so a fixed value of
Na
2
O (melt) of 35 wt% was used. Variation of assumed
Na
2
O from 30to 40 wt% has only a minor effect on the
calculated Zr abundance. Thus, once the melt compos-
ition was calculated, zircon saturation equations were
used to calculate the Zr content necessary for zircon sat-
uration in the calculated melt at the temperature deter-
mined from the Putirka (2016) Hbl thermometer. The
second approach calculated melt Zr contents using Hbl/
melt partition coefficients (d
Zr
).
These two approaches were applied to data from the
English Peak pluton and Tuolumne complex and the
results are summarized in Fig. 4. (It should be noted
that these calculations cannot be done for the Wooley
Creek batholith because too few coincident microprobe
and laser ablation data are available.) The plotted sym-
bols represent amphibole interior zones and the arrow
indicates decreasing T. The saturation equation used
was from Watson & Harrison (1983); this equation
yields the lowest Zr (melt) contents necessary for satur-
ation. Three different d
Zr
values were used: 10 [shown
as the colored symbols and similar to values from
Nandedkar et al., 2016)], 08 and 03 (shown as shaded
symbols and similar to the majority of publishedd
Zr
val-
ues). The closest agreement between these types of cal-
culations (Fig. 4) is ford
Zr
values of 03–04 (e.g. Sisson,
1994; Padilla & Gualda, 2016), rather than the higherd
Zr
values of Nandedkaretal. (2016).
Ifmeltwaslost,howmuchandwhereto?
We have calculated Zr contents necessary to saturate
zircon at 800–825
C in three different ways. Of these
three methods, the one that most closely approaches
bulk-rock Zr contents is the use of chemographic calcu-
lations to determine melt composition followed by use
of the Watson & Harrison (1983) equation for zircon sat-
uration. Nevertheless, each method leads to the same
conclusions: melts were Zr-enriched compared with
bulk-rock compositions and therefore the bulk-rock val-
ues represent partial cumulates. The fact that melt was
lost leads to the question: how much?
A simple answer to this question consists of a com-
parison of the Zr content needed for saturation with the
bulk-rock Zr content. The normalized difference be-
tween these two values should represent the proportion
of melt lost during accumulation; that is,
Zr deficit in rock¼ Zr(melt) – Zr(rock). Therefore, the
approximate proportion of melt lost should be ¼
[Zr(melt) – Zr(rock)]/Zr(melt). In this comparison, we (1)
restricted Hbl compositions to those of core or near-
core compositions and (2) used Zhang et al. (2017)
equations to calculate melt in equilibrium with Hbl (with
Na
2
O content set as 35 wt%) at the temperature calcu-
lated from the Putirka (2016) Hbl thermometer.
For the English Peak pluton, the Boehnkeetal. (2013)
and Watson & Harrison (1983) equations indicate that
bulk-rock samples lost 50–70 % and 35–65 % melt, re-
spectively (Fig. 5a and b). Similarly, for the Tuolumne
complex, percentage of melt loss was calculated as 0–
58 % (Boehnke et al., 2013) and 0–45 % (Watson &
Harrison, 1983)(Fig. 5c and d). These estimates are
within the same range as those made for other intrusive
suites by previous workers: 20–40 % (Deering &
Bachmann, 2010), 28–40 % (Lee & Morton, 2015), and c.
50 % (Schaenetal., 2017).
100
200
300
400
melt Zr (ppm) to saturate calculated melt (W&H)
0 100 200 300 400 500
predicted melt Zr (ppm) from partition coefficents
Hbl/melt
d = 1.0
Hbl/melt
d = 1.0
Hbl/melt d= 0.8
Hbl/melt d= 0.8
1 : 1
1 : 1
Hbl/melt d = 0.3
Hbl/melt
d = 0.3
c
c
oc
Chimney Rock unit
Yellowjacket Ridge unit
border unit
English Peak pluton
100
200
300
0 100 200 300
Tuolumne complex
equigranular Half Dome
porphyritic Half Dome
Cathedral Peak
c
Fig.4. Comparison of Zr contents calculated usingd
Zr
between Hbl and melt (using threed
Zr
values) and Zr calculated by applying
the Watson & Harrison (1983) saturation equation to melt compositions calculated according the Zhangetal. (2017). Colored sym-
bols represent calculated melt using d
Zr
¼ 10. Gray-shaded symbols represent calculated melt using d
Zr
¼ 03. Arrows indicate
decreasing temperature as calculated from the Hbl thermometer of Putirka (2016).
8 JournalofPetrology, 2020, Vol. 0, No. 0
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
260
In the English Peak and Tuolumne magmatic sys-
tems, the proportion of melt remaining after extraction
is in the range of 30–65 % and 100–42 %, respectively.
Assuming that the melt fraction remaining approxi-
mates the magma porosity (Lee et al., 2015), we inter-
pret that melt loss resulted in magmas that ranged from
low-porosity crystal-rich mushes, where melt extraction
was efficient, to highly porous melt-rich magmas. This
spectrum is represented within individual units of the
plutons (it is particularly pronounced in the Tuolumne
Intrusive Complex samples) and suggests that melt ex-
traction was spatially and temporally non-uniform with-
in the magma reservoirs. One consequence of this
variation is that the remaining magmas have distinct
physical properties, which could lead to greater struc-
tural and compositional diversity within the magma
body.
If Zr-rich melts were lost from the systems, where
are they? We suggest that these ‘lost’ melts were
extracted while the magmatic systems were still rela-
tively crystal poor. Unfortunately, as is common with
mid- to upper-crustal plutons, coeval volcanic rocks are
not preserved. Nevertheless, it is instructive to compare
compositions of rhyolitic to dacitic arc rocks with the
plutons under study. In a plot of bulk-rock Zr versus
SiO
2
(Fig. 6), samples from the Tuolumne complex,
late-stage English Peak, and upper Wooley Creek batho-
lith are compared with rhyolite and dacite from five
young arc volcanic complexes and with a suite of
Cretaceous volcanic rocks from the Sierra Nevada. The
Zr contents of the majority of silicic volcanic rocks are
higher, and in some cases significantly higher, than Zr
contents of the plutons.
We recognize that this indirect comparison is imper-
fect. However, it supports studies that concluded that
many granitic plutons do not represent pure melt com-
positions, but instead are products of crystal accumula-
tion or in some instances accumulation of residual
phases (e.g. Bachmann & Bergantz, 2004; Hildreth,
2004, 2007; Bacon & Lowenstern, 2005; Bacon et al.,
2007; Bachmann et al., 2014; Graeter et al., 2015;
Bachmann & Huber, 2016). If the chemical compositions
of many plutonic rocks reflect crystal accumulation,
then statistical analyses comparing plutonic rocks with
volcanic rocks will necessarily fail to find perfect correl-
ation because such analyses are comparing melt-rich
volcanic rocks with cumulate-rich plutonic ones [for ex-
ample, compare statistical approaches (e.g. Kelleretal.,
2015; Glazner et al., 2015) with process-oriented
approaches (e.g. Gelmanetal., 2014)].
Whereisthemissingzircon?
Data and models discussed here indicate that zircon
was a fractionating phase as Hbl crystallized. However,
unlike some plutonic systems (e.g. Spirit Mountain,
Claiborne et al., 2006; Searchlight, Bachl et al., 2001;
Risco Bayo–Huemul, Schaenetal., 2017, 2018) very few
rocks in the plutons studied display evidence of zircon
accumulation (Fig. 1). This means that some of the zir-
con involved in fractionating the melt was lost from
these systems. We suggest two possibilities. The first is
simple: some zircon crystals were small enough to be
removed with escaping melt. The second is more com-
plicated and much more speculative. We consider
40
60
80
Boehnke et al.
Watson & Harrison
40
60
80
% melt lost for zircon saturation of
calculated melt at T(Hbl)
100 120 140 160 180 200
Zr, bulk rock (ppm)
English Peak A
B
C
D
border unit
Yellowjacket Ridge unit
Chimney Rock unit
Boehnke et al.
Watson & Harrison
10
20
30
10
20
30
40
50
% melt lost for zircon saturation of
calculated melt at T(Hbl)
0
80 100
40
50
60
120 140 160 180 200
Zr, bulk rock (ppm)
220
Kuna Crest
Tuolumne complex
equigranular Half Dome
porphyritic Half Dome
Cathedral Peak
Fig. 5. Percentage of melt lost, on the basis of calculated Zr
(melt) compared with bulk-rock Zr. Results are shown for
Boehnkeetal. (2013) and Watson & Harrison (1983) equations.
(See text for details.)
JournalofPetrology, 2020, Vol. 0, No. 0 9
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
261
magmatic systems that are repeatedly rejuvenated by
influxes of new magma. Increases in T associated with
magma influx may be sufficient to dissolve small zircon
grains, but not sufficient to dissolve larger Hbl crys-
tals—whose trace element compositions have already
recorded zircon fractionation. If melt escapes as a result
of the rejuvenation process, the Zr-rich melt is lost but
the associated evidence (Hbl) remains in the mush. If re-
juvenation is repeated, the process could efficiently de-
plete the mush in Zr and Hf.
CONCLUSIONSANDIMPLICATIONS
Major and trace element compositions of magmatic Hbl
in granitic plutons provide important information about
the conditions of amphibole crystallization and about
the coexisting phases. In particular, this study indicates
that zoning patterns of high field strength elements
such as Zr and Hf provide information about the timing
and temperature of zircon saturation, and indicate that
zircon was a stable phase at higher temperatures than
suggested by zircon saturation thermometry.
Even with the use of a conservative criterion such as
decreasing Zr/Hf in Hbl as a marker of zircon saturation,
the data are consistent with zircon saturation at
T 800
C in each of the three systems examined, and
this temperature range agrees with Ti-in-zircon therm-
ometry. Moreover, even with use of the most conserva-
tive zircon saturation algorithm (Watson & Harrison,
1983), the data indicate that Zr (melt) was higher, to
much higher, than Zr (rock). This discrepancy is most
readily explained if a Zr-rich melt was lost prior to so-
lidification; thus the bulk-rock is a cumulate. This con-
clusion was also reached by Werts et al. (2020) for the
Tuolumne complex and by Barnes et al. (2016a) for the
Wooley Creek batholith on the basis of Fe–MgK
D
values
between mafic minerals and host-rock compositions.
We therefore agree with many previous researchers
that bulk-rock compositions of granitic plutons com-
monly do not represent melt compositions, but instead
represent cumulates from which a significant propor-
tion of rhyolitic magma has escaped. Moreover, for Hbl-
bearing granitoids, the method presented here provides
a means by which proportions of lost melt may be
estimated.
This study calls into question the utility of zircon sat-
uration thermometry on plutonic rocks when consider-
ing the survivability of inherited, xenocrystic, or
antecrystic zircons (Harrison et al., 2007; Sie ´gel et al.,
2018). Each of the plutonic systems considered here
contains, at the very least, zircon antecrysts (Memeti
et al., 2010b; Coint et al., 2013b and unpublished; Ernst
et al., 2016), despite the fact that bulk-rock zircon satur-
ation temperatures suggest that such zircons should
have dissolved in the melt. This observation does not
detract from the utility of zircon saturation, rather it
emphasizes the cumulate nature of many coarse-
grained granitic rocks.
These rocks are cumulates as a result of having lost
a relatively Zr-rich rhyolitic melt. Simple mass balance
indicates that all of the analyzed samples of the English
Peak and Wooley Creek systems are cumulate as are
many from the Tuolumne complex. This result means
that use of bulk-rock compositions in thermodynamic
models, such as rhyolite-MELTS (Gualdaetal., 2012), to
model crystallization is problematic, because the rock
compositions reflect some amount of cumulate miner-
als, not a melt. It also means that direct comparison be-
tween bulk compositions of volcanic and potentially
consanguineous plutonic rocks may be complicated by
the presence of plutonic cumulates (e.g. Keller et al.,
2015).
Our results nevertheless indicate that large, incre-
mentally emplaced plutonic complexes are logically the
unerupted products of long-lived, crustal-scale mag-
matic systems (e.g. Bacon & Lowenstern, 2005;
Hildreth, 2007; Lipman, 2007; Reid et al., 2010;
Bachmann et al., 2014; Gelman et al., 2014; Graeter
et al., 2015; Bachmann & Huber, 2016; and many
others). Given the levels of emplacement of all three
plutons studied here (c. 8–12 km), there was ample
room for aggregation and eruption of voluminous rhyo-
litic magma. In the case of the Tuolumne Complex,
0
100
200
300
400
500
600
Zr (ppm)
45 50 55 60 65 70 75 80
SiO
2
(wt%)
late-stage English Peak &
upper Wooley Creek
Mount St. Helens
glass
Taupo (Okataina)
Tuolumne
Fish Canyon
Kamchatka (Karymshina)
Kamchatka (marine tephra)
Chile (Laguna del Maule)
Sierra Nevada
volcanic rocks
Crater Lake climactic eruption
Fig.6. Comparison of Zr and SiO
2
contents in the three pluton-
ic suites studied with arc-related rhyolite and andesite. Data
sources for volcanic rocks: Sierra Nevada from Peck & van
Kooten (1983), Lowe (1995) and present study; Fish Canyon
from Bachmann and Dungan (2002); Taupo from Deering et al.
(2011); Kamchatka from Bindeman et al. (2019); Chile from
Andersen et al. (2017); Crater Lake from C. Bacon (personal
communication, 2017); Mount St Helens from Blundy et al.
(2008).
10 JournalofPetrology, 2020, Vol. 0, No. 0
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
262
from which as much as 50 % melt could have been lost
(Fig. 5c), there is potential for several supervolcano-size
rhyolitic eruptions during the c. 10 Myr history of the
system.
ACKNOWLEDGEMENTS
We thank Calvin Miller and Rita Economos for thor-
ough, thoughtful reviews, Charlotte Allen, Carol Frost,
and Scott Paterson for helpful discussions and reviews
of manuscript drafts, and Adam Kent for editorial han-
dling of the paper.
FUNDING
This research was supported by National Science
Foundation grants EAR-0838342 and 1550969 to C.G.B.
and EAR-1550935 and EAR-1624854 to V.M.
SUPPLEMENTARYDATA
Supplementary data for the paper are available at
JournalofPetrology online.
REFERENCES
Andersen, N. L., Singer, B. S., Jicha, B. R., Beard, B. L.,
Johnson, C. M. & Licciardi, J. M. (2017). Pleistocene to
Holocene growth of a large upper crustal rhyolitic magma
reservoir beneath the active Laguna del Maule volcanic
field, central Chile.JournalofPetrology58, 85–114.
Ague, J. J. & Brimhall, G. H. (1988). Magmatic arc asymmetry
and distribution of anomalous plutonic belts in the batho-
liths of California: Effects of assimilation, crustal thickness,
and depth of crystallization. Geological Society of America
Bulletin100, 912–927, doi:10.1130/0016 7606.
Anderson, J. L. (1996). Status of thermobarometry in granitic
batholiths. Earth and Environmental Science Transactions
oftheRoyalSocietyofEdinburgh87, 125–138.
Anderson, J. L., Barth, A. P., Wooden, J. L. & Mazdab, F. (2008).
Thermometers and thermobarometers in granitic systems.
In: Putirka, K. D., & Tepley, F. J., III (eds)Minerals,Inclusions
and Volcanic Processes. Mineralogical Society of America
and Geochemical Society, Reviews in Mineralogy and
Geochemistry69, 121–142, doi:10.2138/rmg.2008.69.4.
Andersen, D. J. & Lindsley, D. H. (1988). Internally consistent
solution models for Fe–Mg–Mn–Ti oxides: Fe–Ti oxides.
AmericanMineralogist73, 714–726.
Ardill, K., Paterson, S. & Memeti, V. (2018). Spatiotemporal
magmatic focusing in upper–mid crustal plutons of the
Sierra Nevada arc.Earth and Planetary ScienceLetters498,
88–100.
Bachl, C., Miller, C., Miller, J. & Faulds, J. (2001). Construction
of a pluton: Evidence from an exposed cross section of the
Searchlight pluton, Eldorado Mountains, Nevada.
GeologicalSocietyofAmericaBulletin113, 1213–1228.
Bachmann, O. & Dungan, M. A. (2002). Temperature-induced
Al-zoning in hornblendes of the Fish Canyon magma,
Colorado.AmericanMineralogist87, 1062–1076.
Bachmann, O. & Bergantz, G. W. (2004). On the origin of
crystal-poor rhyolites: Extracted from batholithic crystal
mushes.JournalofPetrology45, 1565–1582.
Bachmann, O., Deering, C. D., Lipman, P. W. & Plummer, C.
(2014). Building zoned ignimbrites by recycling silicic
cumulates: insight from the 1,000 km
3
Carpenter Ridge Tuff,
CO. Contributions to Mineralogy and Petrology 167, article
number 1025.
Bachmann, O., Dungan, M. A. & Bussy, F. (2005). Insights into
shallow magmatic processes in large silicic magma bodies:
the trace element record in the Fish Canyon magma body,
Colorado. Contributions to Mineralogy and Petrology 149,
338–349.
Bachmann, O., Dungan, M. & Lipman, P. (2002). The Fish
Canyon magma body, San Juan Volcanic Field, Colorado:
Rejuvenation and eruption of an upper-crustal batholith.
JournalofPetrology43, 1469–1503.
Bachmann, O. & Huber, C. (2016). Silicic magma reservoirs in
the Earth’s crust.AmericanMineralogist101, 2377–2404.
Bacon, C. R. & Lowenstern, J. B. (2005). Late Pleistocene grano-
diorite source for recycled zircon and phenocrysts in rhyo-
dacite lava at Crater Lake, Oregon. Earth and Planetary
ScienceLetters233, 277–293.
Bacon, C. R., Sisson, T. W. & Mazdab, F. K. (2007). Young cumu-
late complex beneath Veniaminof caldera, Aleutian arc,
dated by zircon in erupted plutonic blocks. Geology 35,
491–494.
Barnes, C. G. (1983). Petrology and upward zonation of the
Wooley Creek batholith, Klamath Mountains, California.
JournalofPetrology24, 495–537.
Barnes, C. G., Allen, C. M., Hoover, J. D. & Brigham, R. H.
(1990). Magmatic components of a tilted plutonic system,
Klamath Mountains, California. In: Anderson, J. L. (ed.) The
Nature and Origin of Cordilleran Magmatism. Geological
SocietyofAmerica,Memoirs174, 331–346.
Barnes, C. G., Allen, C. M. & Saleeby, J. B. (1986). Open- and
closed-system characteristics of a tilted plutonic system,
Klamath Mountains, California. Journal of Geophysical
Research91, 6073–6090.
Barnes, C. G., Berry, R., Barnes, M. A. & Ernst, W. G. (2017).
Trace element zoning in hornblende: tracking and modeling
the crystallization of a calc-alkaline arc pluton. American
Mineralogist102, 2390–2405.
Barnes, C., Burton, B., Burling, T., Wright, J. & Karlsson, H.
(2001). Petrology and geochemistry of the Late Eocene
Harrison Pass pluton, Ruby Mountains core complex, north-
eastern Nevada.JournalofPetrology42, 901–929.
Barnes, C. G., Coint, N. & Yoshinobu, A. (2016). Crystal accumu-
lation in a tilted arc batholith. American Mineralogist 101,
1719–1734.
Barnes, C. G., Ernst, W. G., Berry, R. & Tsujimori, T. (2016).
Petrology and geochemistry of an upper crustal pluton: a
view into crustal-scale magmatism during arc to retro-arc
transition.JournalofPetrology57, 1361–1388.
Barnes, C. G., Memeti, V. & Coint, N. (2016). Deciphering mag-
matic processes in calc-alkaline plutons using trace element
zoning in hornblende.AmericanMineralogist101, 328–342.
Bateman, P. C. (1992).PlutonisminthecentralpartoftheSierra
Nevada Batholith, California. US Geological Survey,
Professional Papers 1483, 186 pp.
Bateman, P. C. & Chappell, B. W. (1979). Crystallization, frac-
tionation, and solidification of the Tuolumne Intrusive
Series, Yosemite National Park, California. Geological
SocietyofAmericaBulletin90, 465–482.
Bateman, P. C. & Nokleberg, W. J. (1978). Solidification of the
Mount Givens granodiorite, Sierra Nevada, California.
JournalofGeology86, 563–579.
Bea, F., Montero, P. & Ortega, M. (2006). A LA-ICP-MS evalu-
ation of Zr reservoirs in common crystal rocks: Implications
for Zr and Hf geochemistry, and zircon-forming processes.
CanadianMineralogist44, 693–714.
JournalofPetrology, 2020, Vol. 0, No. 0 11
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
263
Bindeman, I. N., Leonov, V. L., Colo ´ n, D. P., Rogozin, A. N.,
Shipley, N., Jicha, B., Loewen, M. W. & Gerya, T. V. (2019).
Isotopic and petrologic investigation, and a thermomechani-
cal model of genesis of large-volume rhyolites in arc envi-
ronments: Karymshina Volcanic Complex, Kamchatka,
Russia.FrontiersinEarthScience6, 238.
Blundy, J., Cashman, K. & Berlo, K. (2008). Evolving magma
storage conditions beneath Mount St. Helens inferred from
chemical variations in melt inclusions from the 1980–1986
and current (2004–2006) eruptions. In: Sherrod, D. R., Scott,
W. E., & Stauffer, P. H. (eds) A Volcano Rekindled: The
Renewed Eruption of Mount St. Helens, 2004–2006. US
GeologicalSurvey,ProfessionalPapers1750, 755–790.
Boehnke, P., Watson, E. B., Trail, D., Harrison, T. M. & Schmitt,
A. K. (2013). Zircon saturation re-revisited. Chemical
Geology351, 324–334.
Borisov, A. & Aranovich, L. (2019). Zircon solubility in silicate
melts: New experiments and probability of zircon crystalliza-
tion in deeply evolved basic melts. Chemical Geology 510,
103–112.
Burgess, S. D. & Miller, J. S. (2008). Construction, solidification
and internal differentiation of the large felsic arc pluton:
Cathedral Peak granodiorite, Sierra Nevada Batholith. In:
Burgess, S. D., & Miller, J. S. (eds) Dynamics of Crustal
Magma Transfer, Storage and Differentiation. Geological
Society,London,SpecialPublications304, 203–233.
Cao, W., Paterson, S. R., Memeti, V., Mundil, R., Anderson, J. L.
& Schmidt, K. (2015). Tracking paleodeformation fields in
the Mesozoic central Sierra Nevada arc: Implications for
intra-arc cyclic deformation and arc tempos. Lithosphere 7,
296–320, doi:10.1130/L389.1.
Chappell, B. W., Bryant, C. J., Wyborn, D., White, A. J. R. &
Williams, I. S. (1998). High- and low-temperature I-type gran-
ites.ResourceGeology48, 225–235.
Chappell, B. W., White, A. J. R. & Wyborn, D. (1987). The import-
ance of residual source material (restite) in granite petrogen-
esis.JournalofPetrology28, 1111–1138.
Claiborne, L. L., Miller, C. F., Walker, B. A., Wooden, J. L.,
Mazdab, F. K. & Bea, F. (2006). Tracking magmatic processes
through Zr/Hf ratios in rocks and Hf and Ti zoning in zircons:
an example from the Spirit Mountain batholith, Nevada.
MineralogicalMagazine70, 517–543.
Coint, N., Barnes, C. G., Yoshinobu, A. S., Barnes, M. A. & Buck,
S. (2013). Use of trace element abundances in augite and
hornblende to determine the size, connectivity, timing, and
evolution of magma batches in a tilted pluton.Geosphere9,
1747–1765.
Coint, N., Barnes, C. G., Yoshinobu, A. S., Chamberlain, K. R. &
Barnes, M. A. (2013). Batch-wise assembly and zoning of a
tilted calc-alkaline batholith: field relations, timing, and com-
positional variation.Geosphere9, 1729–1746.
Coleman, D. S., Bartley, J. M., Glazner, A. F. & Pardue, M. J.
(2012). Is chemical zonation in plutonic rocks driven by
changes in source magma composition or shallow-crustal
differentiation?Geosphere8, 1568–1587.
Coleman, D. S., Gray, W. & Glazner, A. F. (2004). Rethinking the
emplacement and evolution of zoned plutons:
Geochronologic evidence for incremental assembly of the
Tuolumne Intrusive Suite, California.Geology32, 433–436.
Deering, C. D. & Bachmann, O. (2010). Trace element indicators
of crystal accumulation in silicic igneous rocks. Earth and
PlanetaryScienceLetters297, 324–331.
Deering, C. D., Bachmann, O., Dufek, J. & Gravley, D. M. (2011).
Rift-related transition from andesite to rhyolite volcanism in
the Taupo Volcanic Zone (New Zealand) controlled by crys-
tal–melt dynamics in mush zones with variable mineral
assemblages.JournalofPetrology52, 2243–2263.
Economos, R. C., Memeti, V., Paterson, S. R., Miller, J. S.,
Erdmann, S. &
Za ´k, J. (2009). Causes of compositional diver-
sity in a lobe of the Half Dome granodiorite, Tuolumne
Batholith, central Sierra Nevada, California. Earth and
EnvironmentalScienceTransactionsoftheRoyalSocietyof
Edinburgh100, 173–183.
Ernst, W. G., Gottlieb, E. S., Barnes, C. G. & Hourigan, J. K.
(2016). Zircon U–Pb ages and petrologic evolution of the
English Peak granitic pluton: Jurassic crustal growth in
northwestern California.Geosphere12, 1422–1436.
Ernst, W. & Liu, J. (1998). Experimental phase-equilibrium
study of Al- and Ti-contents of calcic amphibole in
MORB—A semiquantitative thermobarometer. American
Mineralogist83, 952–969.
Ferry, J. M. & Watson, E. B. (2007). New thermodynamic mod-
els and revised calibrations for the Ti-in-zircon and
Zr-in-rutile thermometers. Contributions to Mineralogy and
Petrology154, 429–437.
Fiedrich, A. M., Bachmann, O., Ulmer, P., Deering, C. D., Kunze,
K. & Leuthold, J. (2017). Mineralogical, geochemical, and
textural indicators of crystal accumulation in the Adamello
Batholith (Northern Italy). American Mineralogist 102,
2467–2483.
Fiske, R. S. & Tobisch, O. T. (1994). Middle Cretaceous ash-flow
tuff and caldera-collapse deposit in the Minarets Caldera,
east–central Sierra Nevada, California.GeologicalSocietyof
America Bulletin 106, 582–593, doi:10.1130/0016-7606
(1994)106.
Flood, R. H. & Shaw, S. E. (1979). K-rich cumulate diorite at the
base of a tilted granodiorite pluton from the New England
batholith, Australia.JournalofGeology87, 417–425.
Fuhrman, M. L. & Lindsley, D. H. (1988). Ternary-feldspar mod-
eling and thermometry.AmericanMineralogist73, 201–215.
Gelman, S. E., Deering, C. D., Bachmann, O., Huber, C. &
Gutie ´rrez, F. (2014). Identifying the crystal graveyards
remaining after large silicic eruptions. Earth and Planetary
ScienceLetters403, 299–306.
Gervasoni, F., Klemme, S., Rocha-Ju ´ nior, E. R. V. & Berndt, J.
(2016). Zircon saturation in silicate melts: a new and
improved model for aluminous and alkaline melts.
ContributionstoMineralogyandPetrology171, 21.
Glazner, A. F., Coleman, D. S. & Mills, R. D. (2015). The volca-
nic–plutonic connection. In: Breitkreuz, C. & Rocchi, S. (eds)
Physical Geology of Shallow Magmatic System. Advances
in Volcanology. Cham: Springer, pp. 59–82.
Graeter, K. A., Beane, R. J., Deering, C. D., Gravley, D. M. &
Bachmann, O. (2015). Formation of rhyolite at the Okataina
Volcanic Complex, New Zealand: new insights from analysis
of quartz clusters in plutonic lithics. American Mineralogist
100, 1778–1789.
Gray, W., Glazner, A. F., Coleman, D. S. & Bartley, J. M. (2008).
Long-term geochemical variability of the Late Cretaceous
Tuolumne Intrusive Suite, central Sierra Nevada, California.
In: Annen, C., & Zellmer, G. F. (eds) Dynamics of Crustal
Magma Transfer, Storage and Differentiation. Geological
Society,London,SpecialPublications304, 183–201.
Greene, D. C. & Schweickert, R. A. (1995). The Gem Lake Shear
Zone—Cretaceous dextral transpression in the northern
Ritter Range pendant, eastern Sierra Nevada, California.
Tectonics14, 945–961, doi:10.1029/95TC01509.
Gualda, G. A. R., Ghiorso, M. S., Lemons, R. V. & Carley, T. L.
(2012). Rhyolite-MELTS: a modified calibration of MELTS
optimized for silica-rich, fluid-bearing magmatic systems.
JournalofPetrology53, 875–890.
Hanchar, J. M. & Watson, E. B. (2003). Zircon saturation therm-
ometry. In: Hanchar, J. M., & Hoskin, P. W. O. (eds) Zircon.
12 JournalofPetrology, 2020, Vol. 0, No. 0
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
264
MineralogicalSocietyofAmericaandGeochemicalSociety,
ReviewsinMineralogyandGeochemistry53, 89–112.
Harper, B. E., Miller, C. F., Koteas, G. C., Cates, N. L., Wiebe, R.
A., Lazzareschi, D. S. & Cribb, J. W. (2004). Granites, dynam-
ic magma chamber processes and pluton construction: the
Aztec Wash pluton, Eldorado Mountains, Nevada, USA.
Earth andEnvironmental ScienceTransactionsof theRoyal
SocietyofEdinburgh95, 277–295.
Harrison, T. M., Watson, E. B. & Aikman, A. B. (2007).
Temperature spectra of zircon crystallization in plutonic
rock.Geology35, 635–638.
Hildreth, W. (2004). Volcanological perspectives on Long
Valley, Mammoth Mountain, and Mono Craters: several con-
tiguous but discrete systems. Journal of Volcanology and
GeothermalResearch136, 169–198.
Hildreth, W. (2007). Quaternary magmatism in the
Cascades—geologic perspectives. US Geological Survey,
Professional Papers 1744, 125 pp.
Huber, N. K., Bateman, P. C. & Wahrhaftig, C. (1989). Geologic
map of Yosemite National Park and vicinity, California. US
Geological Survey, Miscellaneous Investigations Series
I-1874.
Humphreys, M. C. S., Cooper, G. F., Zhang, J., Loewen, M.,
Kent, A. J. R., Macpherson, C. G. & Davidson, J. P. (2019).
Unravelling the complexity of magma plumbing at Mount
St. Helens: a new trace element partitioning scheme for
amphibole.Contributionsto Mineralogy and Petrology174,
article number 9,
Keller, C. B., Schoene, B., Barboni, M., Samperton, K. M. &
Husson, J. M. (2015). Volcanic–plutonic parity and the differ-
entiation of the continental crust.Nature523, 301–307.
Kerrick, D. M. (1970). Contact metamorphism in some areas of
the Sierra Nevada, California.GeologicalSocietyofAmerica
Bulletin81, 2913–2938.
Kistler, R. W., Chappell, B. W., Peck, D. L. & Bateman, P. C.
(1986). Isotopic variation in the Tuolumne Intrusive Suite,
central Sierra Nevada, California. Contributions to
MineralogyandPetrology94, 205–220.
Klein, M., Stosch, H.-G. & Seck, H. (1997). Partitioning of high
field-strength and rare-earth elements between amphibole
and quartz-dioritic to tonalitic melts: an experimental study.
ChemicalGeology138, 257–271.
Lee, C.-T. A. & Morton, D. M. (2015). High silica granites: termin-
al porosity and crystal settling in shallow magma chambers.
EarthandPlanetaryScienceLetters409, 23–31.
Lee, C.-T. A., Morton, D. M., Farner, M. J. & Moitra, P. (2015).
Field and model constraints on silicic melt segregation by
compaction/hindered settling: the role of water and its latent
heat release.AmericanMineralogist100, 1762–1777.
Lipman, P. W. (2007). Incremental assembly and prolonged
consolidation of Cordilleran magma chambers: evidence
from the Southern Rocky Mountain volcanic field.
Geosphere3, 42–70.
Lowe, T. K. (1995). Petrogenesis of the Minarets and Merced
Peak volcanic–plutonic complexes, Sierra Nevada,
California. PhD dissertation, Stanford University, 157 pp.
Mahood, G. & Cornejo, P. (1992). Evidence for ascent of differ-
entiated liquids in a silicic magma chamber found in a gran-
itic pluton. Earth and Environmental Science Transactions
oftheRoyalSocietyofEdinburgh83, 63–69.
Matzel, J., Miller, J., Mundil, R., Wooden, J., Mazdab, F.,
Burgess, S., Paterson, S. & Memeti, V. (2007). Growth of the
Tuolumne Batholith: Zircon crystallization temperature, age
and trace element data. American Geophysical Union, Fall
Meeting 2007, Abstract V42C-08.
McCarthy, T. S. & Groves, D. I. (1979). The Blue Tier Batholith,
northeastern Tasmania. A cumulate-like product of
fractional crystallization. Contributions to Mineralogy and
Petrology71, 193–209.
Memeti, V., Gehrels, G. E., Paterson, S. R., Thompson, J. M.,
Mueller, R. M. & Pignotta, G. S. (2010). Evaluating the
Mojave–Snow Lake fault hypothesis and origins of central
Sierran metasedimentary pendant strata using detrital zir-
con provenance analyses. Lithosphere 2, 341–360, doi:
10.1130/L58.1.
Memeti, V., Paterson, S., Matzel, J., Mundil, R. & Okaya, D.
(2010). Magmatic lobes as “snapshots” of magma chamber
growth and evolution in large, composite batholiths: an ex-
ample from the Tuolumne intrusion, Sierra Nevada,
California. Geological Society of America Bulletin 122,
1912–1931, doi:10.1130/B30004.1.
Memeti, V., Paterson, S. & Mundil, R. (2014). Day 4: Magmatic
evolution of the Tuolumne Intrusive Complex. In: Memeti,
V., Paterson, S. R., & Putirka, K. D. (eds) Formation of the
Sierra Nevada Batholith: Magmatic and Tectonic Processes
and Their Tempos. Geological Society of America Field
Guide34, 43–74, doi:10.1130/2014.0034(04).
Miller, J. S., Matzel, J. E. P., Miller, C. F., Burgess, S. D. & Miller,
R. B. (2007). Zircon growth and recycling during the assem-
bly of large composite arc plutons. Journal of Volcanology
andGeothermalResearch167, 282–299.
Miller, C. F., McDowell, S. M. & Mapes, R. W. (2003). Hot and
cold granites?: Implications of zircon saturation tempera-
tures and preservation of inheritance.Geology31, 529–532.
Miller, C. F. & Miller, J. S. (2002). Contrasting stratified plutons
exposed in tilted blocks, Eldorado Mountains, Colorado
River rift, NV, USA.Lithos61, 209–224.
Nandedkar, R. H., Hu ¨ rlimann, N., Ulmer, P. & Mu ¨ ntener, O.
(2016). Amphibole–melt trace element partitioning of frac-
tionating calc-alkaline magmas in the lower crust: an experi-
mental study. Contributions to Mineralogy and Petrology
171, article number 71,
Otten, M. T. (1984). The origin of brown hornblende in the
Artfja ¨llet gabbro and dolerites. Contributions to Mineralogy
andPetrology86, 189–199.
Padilla, A. J. & Gualda, G. A. R. (2016). Crystal–melt element
partitioning in silicic magmatic systems: an example from
the Peach Spring Tuff high-silica rhyolite, Southwest USA.
ChemicalGeology440, 326–344.
Paterson, S. R. (2009). Magmatic tubes, pipes, troughs, diapirs,
and plumes: Late-stage convective instabilities resulting in
compositional diversity and permeable networks in
crystal-rich magmas of the Tuolumne batholith, Sierra
Nevada, California.Geosphere5, 496–527.
Paterson, S. R. & Memeti, V. (2014). Day 5: Mesozoic volcanic
rocks of the Central Sierra Nevada Arc. In: Memeti, V.,
Paterson, S. R., & Putirka, K. D. (eds)FormationoftheSierra
Nevada Batholith: Magmatic and Tectonic Processes and
Their Tempos. Geological Society of America Field Guide
34, 75–85, doi:10.1130/2014.0034(05).
Paterson, S., Memeti, V., Mundil, R. &
Za ´ k, J. (2016). Repeated,
multiscale, magmatic erosion and recycling in an
upper-crustal pluton: Implications for magma chamber dy-
namics and magma volume estimates. American
Mineralogist101, 2176–2198.
Peck, D. L. (1980). Geologic map of the Merced Peak quadran-
gle, centralSierraNevada,California. USGeologicalSurve,
GeologicQuadrangleMapGQ-1531, scale 1:62,500.
Peck, D. L. & van Kooten, G. K. (1983).MercedPeakquadrangle,
central Sierra Nevada, California—Analytic data. US
GeologicalSurvey,ProfessionalPapers1170-D.
Putirka, K. (2016). Amphibole thermometers and barometers
for igneous system, and some implications for eruption
JournalofPetrology, 2020, Vol. 0, No. 0 13
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
265
mechanisms of felsic magmas at arc volcanoes. American
Mineralogist101, 841–858.
Ratschbacher, B. C., Keller, C. B., Schoene, B., Paterson, S. R.,
Anderson, J. L., Okaya, D., Putirka, K. & Lippoldt, R. (2018). A
new workflow to assess emplacement duration and melt
residence time of compositionally diverse magmas
emplaced in a sub-volcanic reservoir. Journal of Petrology
59, 1787–1810.
Reid, M. R., Vazquez, J. A. & Schmitt, A. K. (2010). Zircon-scale
insights into the history of a supervolcano, Bishop Tuff,
Long Valley, California, with implications for the Ti-in-zircon
geothermometer. Contributions to Mineralogy and
Petrology159, 313–314,
Ridolfi, F. & Renzulli, A. (2012). Calcic amphiboles in
calc-alkaline and alkaline magmas: thermobarometric and
chemimetric empirical equations valid up to 1,130
C and
2.2 GPa. Contributions to Mineralogy and Petrology 163,
877–895.
Ridolfi, F., Renzulli, A. & Puerini, M. (2010). Stability and chem-
ical equilibrium of amphibole in calc-alkaline magmas: an
overview, new thermobarometric formulations and applica-
tions to subduction-related volcanoes. Contributions to
MineralogyandPetrology160, 45–66.
Schaen, A. J., Cottle, J. M., Singer, B. S., Keller, C. B., Garibaldi,
N. & Schoene, B. (2017). Complementary crystal accumula-
tion and rhyolite melt segregation in a late Miocene Andean
pluton.Geology45, 835–838.
Schaen, A. J., Singer, B. S., Cottle, J. M., Garibaldi, N., Schoene,
B., Satkoski, A. M. & Fournelle, J. (2018). Textural and min-
eralogical record of low pressure melt extraction and silicic
cumulate formation in the late Miocene Risco Bayo–Huemul
plutonic complex, southern Andes.JournalofPetrology59,
1991–2016.
Schiller, D. & Finger, F. (2019). Application of Ti-in-zircon therm-
ometry to granite studies: problems and possible solutions.
ContributionstoMineralogyandPetrology174, 51.
Sie ´gel, C., Bryan, S. E., Allen, C. M. & Gust, D. A. (2018). Use
and abuse of zircon-based thermometers: A critical review
and a recommended approach to identify antecrystic zir-
cons.Earth-ScienceReviews176, 87–116.
Sisson, T. W. (1994). Hornblende–melt trace-element partition-
ing measured by ion microprobe. Chemical Geology 117,
331–344.
Srogi, L. & Lutz, T. M. (1996). The role of residual melt migration
in producing compositional diversity in a suite of granitic
rocks.EarthandPlanetaryScienceLetters144, 563–576.
Srogi, L. & Lutz, T. M. (1997). Chemical variation in plutonic
rocks caused by residual melt migration: implications for
granite petrogenesis. In: Sinha, A. K., Whalen, J. B., &
Hogan, J. P. (eds) The Nature of Magmatism in the
Appalachian Orogen. Geological Society of America,
Memoirs191, 309–335.
Turnbull, R., Weaver, S., Tulloch, A., Cole, J., Handler, M. &
Ireland, T. (2010). Field and geochemical constraints on
mafic–felsic interactions, and processes in high-level arc
magma chambers: an example from the Halfmoon Pluton,
New Zealand.JournalofPetrology51, 1477–1505.
Vernon, R. H. & Paterson, S. R. (2008). Mesoscopic structures
resulting from crystal accumulation and melt movement in
granites. Transactions of the Royal Society of Edinburgh:
EarthSciences97, 369–381.
Walker, B. A., Jr, Miller, C. F., Claiborne, L. L., Wooden, J. L. &
Miller, J. S. (2007). Geology and geochronology of the Spirit
Mountain batholith, southern Nevada: implications for time-
scales and physical processes of batholith construction.
Journal of Volcanology and Geothermal Research 167,
239–262.
Watson, E. B. & Harrison, T. M. (1983). Zircon saturation revis-
ited: temperature and composition effects in a variety of
crustal magma types. Earth and Planetary Science Letters
64, 295–304.
Weinberg, R. F. (2006). Melt segregation structures in granitic
plutons.Geology34, 305–308.
Werts, K. R., Barnes, C. G., Memeti, V., Paterson, S.,
Ratschbacher, B. & Williams, D. R. (2018). Volcanic vs plu-
tonic hornblende: tools for discerning crystal accumulation
and melt loss in plutons. Abstract. Goldschmidt Conference,
Boston, MA, August 13–17.
Werts, K., Barnes, C. G., Memeti, V., Paterson, S., Ratschbacher,
B. & Williams, D. (2020). Hornblende as a tool for assessing
mineral–melt equilibrium and recognition of crystal accu-
mulation.AmericanMineralogist105, 77–91.
Wiebe, R. A. (1996). Mafic–silicic layered intrusions: the role of
basaltic injections on magmatic processes and the evolution
of silicic magma chambers. Earth and Environmental
Science Transactions of the Royal Society of Edinburgh 87,
233–242.
Wiebe, R. & Collins, W. (1998). Depositional features and strati-
graphic sections in granitic plutons: implications for the em-
placement and crystallization of granitic magma.Journalof
StructuralGeology20, 1273–1289.
Wiebe, R. A., Jellinek, A. M. & Hodge, K. F. (2017). New insights
into the origin of ladder dikes: Implications for punctuated
growth and crystal accumulation in the Cathedral Peak
granodiorite.Lithos277, 241–258.
Zhang, J., Humphreys, M. C. S., Cooper, G. F., Davidson, J. P. &
Macpherson, C. G. (2017). Magma mush chemistry at sub-
duction zones, revealed by new melt major element inver-
sion from calcic amphiboles. American Mineralogist 102,
1353–1367.
14 JournalofPetrology, 2020, Vol. 0, No. 0
Downloaded from https://academic.oup.com/petrology/article-abstract/doi/10.1093/petrology/egaa008/5803082 by USC Law user on 23 April 2020
266
Appendix B: Supplementary data to Chapter 2
Appendix B contains the datasets compiled and presented in Chapter 2. It can also be accessed via the link
to the published article: Ardill et al. (2018), EPSL (https://doi.org/10.1016/j.epsl.2018.06.023).
Supplementary Figure 1 is a map of the central Sierra Nevada color coded by age (color-coding shown in
Figure 2.3 within the manuscript (p. 16). The star symbols show the location of U-Pb zircon ages used to
estimate focusing patterns and rates, the color of the symbol indicates the method. Supplementary Figure 2
summarizes major element trends within the Cretaceous compilation, color coded by age according to the
legend. These correspond to patterns identified by Moore (1959) and Bateman (1992), described in Chapter
2, section 2 (p. 14). Supplementary Table S1 includes a list of references to published data included in the
compilation and a reference key to the map Supplementary Figure 1. Supplementary Table S2 is a table
summarizing new LA-ICP-MS U-Pb zircon ages. Supplementary Table S3 includes a list of references to
published geochemical datasets compiled in this study and shown in Figure 2.5, 2.6, and Supplementary
Figure 2. Supplementary Table S4 summarizes new major and trace element and isotopic data collected in
this study. Supplementary Table S5 presents the results of the pluton area and volume estimates for the
central Sierra Nevada, building on the compilation of Karlstrom et al. (2017).
267
0 8 km
96
100
96
109
102
120
104
118
123
95
100
90
94
94
97
99
95.2±0.2
97
99
99
95 ±2
97.4 ±0.4
117.4 ±2.1
101.8 ±0.2
103-107
102-103
102-103
102-104
100.32±0.3
100.63±0.31
98.-101
118.4
107.3
119.7
117.4
95
97
93
96
96
96
93
93
98
97
86?
98
98
100
97 112
93?
97?
114 114?
117
103
96?
103
119
116
116
113
108
114
Stern et al., 1981
Published, unknown
method
Bulk zircon fractions
Single zircon ages
(LA-ICP-MS/ TIMS)
Single zircon ages
(LA-ICP-MS)
Unpublished new ages
(LA-ICP-MS)
U-Pb zircon ages
87.4 ±0.4
85.1 ±0.9
83.9 ±0.3
88.5 ±0.12
90.6 ±0.2
92.9±0.11
94.4 ±0.3
93.6 ±0.4
87.7±1.1
89.7 ±0.2
90.1±0.1
94.9 ±0.3
91.1 ±0.2
90.6 ±0.2
89.8±0.2
87.0±0.7
87.3±0.7
89.9 ±0.2
90.2 ±0.2
86.2 ±0.1
87.3±0.2
92.8±0.4
91.5±0.1
93.5
>0.706
<0.706
268
K
2
O
Position along migration transect (reprojected E)
0 8 5 - < 9 0
0 9 0 - < 9 5
0 9 5 - < 1 0 0
1 0 0 - < 1 0 2
1 0 2 - < 1 0 5
1 0 5 - < 1 1 0
1 1 0 - < 1 1 5
1 1 5 - < 1 2 0
1 2 0 - < 1 2 5
1 2 5 - < 1 3 0
-120.2 -120.0 -119.8 -119.6 -119.4 -119.2 -119.0
0
1
2
3
4
5
6
7
0 8 5 - < 9 0
0 9 0 - < 9 5
0 9 5 - < 1 0 0
1 0 0 - < 1 0 2
1 0 2 - < 1 0 5
1 0 5 - < 1 1 0
1 1 0 - < 1 1 5
1 1 5 - < 1 2 0
1 2 0 - < 1 2 5
1 2 5 - < 1 3 0
SiO
2
Position along migration transect (reprojected E)
-120.2 -120.0 -119.8 -119.6 -119.4 -119.2 -119.0
35
40
45
50
55
60
65
70
75
80
0
2
4
6
8
10
12
14
16
CaO
Position along migration transect (reprojected E)
-120.2 -120.0 -119.8 -119.4 -119.2 -119.0 -119.6
0 8 5 - < 9 0
0 9 0 - < 9 5
0 9 5 - < 1 0 0
1 0 0 - < 1 0 2
1 0 2 - < 1 0 5
1 0 5 - < 1 1 0
1 1 0 - < 1 1 5
1 1 5 - < 1 2 0
1 2 0 - < 1 2 5
1 2 5 - < 1 3 0
0 8 5 - < 9 0
0 9 0 - < 9 5
0 9 5 - < 1 0 0
1 0 0 - < 1 0 2
1 0 2 - < 1 0 5
1 0 5 - < 1 1 0
1 1 0 - < 1 1 5
1 1 5 - < 1 2 0
1 2 0 - < 1 2 5
1 2 5 - < 1 3 0
12
14
16
18
20
22
24
26
Al
2
O
3
Position along migration transect (reprojected E)
-120.2 -120.0 -119.8 -119.4 -119.2 -119.0 -119.6
269
Supplementary Table S1: U-Pb zircon ages
References for data used in Figures 2 and 3
Source U-Pb zircon Method Map Reference
Published compilations
Bateman, 1992
Bateman, P.C., 1992. Plutonism in the central part of the Sierra Nevada
batholith, California, U.S. Geological Survey Open-File Report (No.
1483).
mixed g
Chapman et al., 2012 (Table SD-T4)
Chapman, A.D., Saleeby, J.B., Wood, D.J., Piasecki, A., Kidder, S.,
Ducea, M.N. and Farley, K.A., 2012. Late Cretaceous gravitational
collapse of the southern Sierra Nevada batholith,
California. Geosphere, 8(2), pp.314-341.
mixed g
Irwin and Wooden., 2001
Irwin, W.P. and Wooden, J.L., 2001. Map showing plutons and
accreted terranes of the Sierra Nevada, California with a tabulation of
U/Pb isotopic ages: U.S. Geological Survey Open-File Report 01-229
mixed g
Additional sources
Bracciali et al., 2008
Bracciali, L., Paterson, S., Memeti, V., Rocchi, S., and Mundil, R.
(2008) Filling the magma chamber of the Tuolumne Batholith, Sierra
Nevada, California: LASIIII conference. Physical Geology of
Subvolcanic Systems, Laccolith, Sills, and Dykes, Elba Island.
ID-TIMS single zircon h
Burgess and Miller 2008
Burgess, S.D. and Miller, J.S., 2008. Construction, solidification and
internal differentiation of a large felsic arc pluton: Cathedral Peak
granodiorite, Sierra Nevada Batholith. Geological Society, London,
Special Publications, 304(1), pp.203-233.
ID-TIMS single zircon e
Burgess et al., 2009
Burgess, S.D., Bowring, S.A., Petsche, J., Miller, R.B., and Miller, J.S.,
2009, High precision U-Pb CA-TIMS geochronology for the Sentinel
and Yosemite Creek granodiorites, Sierra Nevada batholith, CA: A
history of punctuated intrusion and protracted crystallization: EOS
(Transactions, American Geophysical Union), v. 52, p. 90.
ID-TIMS single zircon e
Cao, 2015
Cao, W., 2015. Links, Tempos, and Mass Balances of Cyclic
Deformation and Magmatism in Arcs: a Case Study on the Mesozoic
Sierra Nevada Arc Integrating Geological Mapping, Geochronology,
Geobarometry, Strain Analyses and Numerical Simulations, Doctoral
dissertation, University of Southern California. p.410
LA-ICPMS single zircon a
Cao et al., 2015
Cao, W., Paterson, S., Memeti, V., Mundil, R., Anderson, J.L. and
Schmidt, K., 2015. Tracking paleodeformation fields in the Mesozoic
central Sierra Nevada arc: Implications for intra-arc cyclic deformation
and arc tempos. Lithosphere, 7(3), pp.296-320.
ID-TIMS single zircon b
Coleman et al., 2004
Coleman, D.S., Gray, W. and Glazner, A.F., 2004. Rethinking the
emplacement and evolution of zoned plutons: Geochronologic evidence
for incremental assembly of the Tuolumne Intrusive Suite,
California. Geology, 32(5), pp.433-436.
ID-TIMS bulk zircon -
Fiske and Tobisch, 1994
Fiske, R.S. and Tobisch, O.T., 1994. Middle Cretaceous ash-flow tuff
and caldera-collapse deposit in the Minarets Caldera, east-central Sierra
Nevada, California. Geological Society of America Bulletin, 106(5),
pp.582-593.
bulk zircon f
Huber et al., 1989
Huber, N.K., Bateman, P.C. and Wahrhaftig, C., 1989. Geologic map of
Yosemite National Park and vicinity, California (No. 1874)
Map
Lackey et al., 2012
Lackey, J.S., Cecil, M.R., Windham, C.J., Frazer, R.E., Bindeman, I.N.
and Gehrels, G.E., 2012. The Fine Gold Intrusive Suite: The roles of
basement terranes and magma source development in the Early
Cretaceous Sierra Nevada batholith. Geosphere, 8(2), pp.292-313.
LA-ICPMS single zircon d
Matzel et al., 2005
Matzel, J., Mundil, R., Paterson, S., Renne, P., and Nomade, S., 2005,
Evaluating pluton growth models using high resolution geochronology:
Tuolumne intrusive suite, Sierra Nevada, CA: Geological Society of
America Abstracts with Programs, v. 37, no. 7, p. 131
ID-TIMS single zircon g
Matzel et al., 2006b
Matzel, J., Miller, J.S., Mundil, R., and Paterson, S.R., 2006b, Zircon
saturation and the growth of the Cathe dral Peak pluton, CA:
Geochimica et Cosmochimica Acta, v. 70, no. 18, p. A403, doi:
10.1016/ j.gca.2006.06.813.
ID-TIMS single zircon g
McNulty et al., 1996
McNulty, B.A., Tong, W. and Tobisch, O.T., 1996. Assembly of a dike-
fed magma chamber: The Jackass Lakes pluton, central Sierra Nevada,
California. Geological Society of America Bulletin, 108(8), pp.926-940.
zircon fractions d
Memeti et al., 2010a
Memeti, V., Paterson, S., Matzel, J., Mundil, R. and Okaya, D., 2010.
Magmatic lobes as “snapshots” of magma chamber growth and
evolution in large, composite batholiths: An example from the
Tuolumne intrusion, Sierra Nevada, California. Geological Society of
America Bulletin, 122(11-12), pp.1912-1931.
ID-TIMS single zircon a
Memeti et al., 2010b
Memeti, V., Gehrels, G.E., Paterson, S.R., Thompson, J.M., Mueller,
R.M. and Pignotta, G.S., 2010. Evaluating the Mojave–Snow Lake fault
hypothesis and origins of central Sierran metasedimentary pendant
strata using detrital zircon provenance analyses. Lithosphere, 2(5),
pp.341-360.
ID-TIMS single zircon a
Paterson et al., 2016
Paterson, S., Memeti, V., Mundil, R. and Žák, J., 2016. Repeated,
multiscale, magmatic erosion and recycling in an upper-crustal pluton:
Implications for magma chamber dynamics and magma volume
estimates. American Mineralogist, 101(10), pp.2176-2198.
ID-TIMS single zircon f
Ratajeski et al., 2001
Ratajeski, K., Glazner, A.F. and Miller, B.V., 2001. Geology and
geochemistry of mafic to felsic plutonic rocks in the Cretaceous
intrusive suite of Yosemite Valley, California. Geological Society of
America Bulletin, 113(11), pp.1486-1502.
zircon fractions e
Saleeby et al., 1989
Saleeby, J.B., Shaw, H.F., Niemeyer, S., Moores, E.M. and Edelman,
S.H., 1989. U/Pb, Sm/Nd and Rb/Sr geochronological and isotopic
study of northern Sierra Nevada ophiolitic assemblages,
California. Contributions to Mineralogy and Petrology, 102(2), pp.205-
220.
zircon fractions -
Saleeby, 2007
Saleeby, J., 2007. The western extent of the Sierra Nevada batholith in
the Great Valley basement and its significance in underlying mantle
dynamics. In AGU Fall Meeting Abstracts.
single zircon -
Stern et al., 1981
Stern, T.W., Bateman, P.C., Morgan, B.A., Newell, M.F., and Peck,
D.L., 1981, Isotopic U-Pb ages of zircon from the granitoids of the
central Sierra Nevada, California: U.S. Geological Survey Professional
Paper 1185, 17 p.
bulk zircon g
Tobisch et al., 1995
Tobisch, O. T., Saleeby, J. B., Renne, P. R., McNulty, B. A., and Tong,
W., 1995a, Variations in deformation fields during development of a
large volume magmatic arc, central Sierra Nevada, California:
Geological Society of America Bulletin, v. 107, p. 148–166.
zircon fractions c
Tomek et al., 2016
Tomek, F., Žák, J., Verner, K., Holub, F. V., Sláma, J., Paterson, S. R.,
and Memeti, V., 2016, Mineral fabrics in high-level intrusions
recording crustal strain and volcano–tectonic interactions: the
Shellenbarger pluton, Sierra Nevada, California: Journal of the
Geological Society.
LA-ICPMS single zircon c
270
Supplementary Table S2: U-Pb zircon ages
New data: LA-ICPMS single zircon ages, plotted in Figures 2 and 3
Method:
Isotope ratios Apparent ages (Ma)
Analysis U 206Pb U/Th 206Pb* ± 207Pb* ± 206Pb* ± error 206Pb* ± 207Pb* ± 206Pb* ± Best age ± Conc Notes
(ppm) 204Pb 207Pb* (%) 235U* (%) 238U (%) corr. 238U* (Ma) 235U (Ma) 207Pb* (Ma) (Ma) (Ma) (%)
IMP39-A Grizzly Peak porphyritic tonalite
-1MP39A Spot 22 872 36646 1.5 20.4552 1.6 0.106 3.4 0.0157 3 0.89 100.6 3 102.3 3.3 142.4 36.8 100.6 3 NA RIM1
-1MP39A Spot 3 524 17275 2.5 21.0261 1.6 0.104 3.5 0.0159 3.1 0.89 101.5 3.2 100.4 3.4 77.4 37.8 101.5 3.2 NA
-1MP39A Spot 20 942 95748 1.9 20.4431 1.6 0.1072 3.8 0.0159 3.5 0.91 101.7 3.5 103.4 3.8 143.8 37.5 101.7 3.5 NA
-1MP39A Spot 30 989 41407 2.6 20.4938 1.6 0.1072 3.4 0.0159 3 0.88 102 3.1 103.4 3.4 138 38.4 102 3.1 NA
-1MP39A Spot 27 731 29030 1.8 21.0663 2 0.1045 4.4 0.016 3.9 0.89 102.2 4 100.9 4.3 72.9 48.5 102.2 4 NA
-1MP39A Spot 7 816 112532 2 20.7503 1.5 0.1066 2.8 0.016 2.4 0.85 102.6 2.4 102.8 2.8 108.7 35.1 102.6 2.4 NA CORE2
-1MP39A Spot 18 840 18443 3 20.9563 1.5 0.1062 3.3 0.0161 2.9 0.89 103.2 3 102.5 3.2 85.3 36.2 103.2 3 NA
-1MP39A Spot 33 829 55436 3.2 20.9763 1.4 0.1062 3.9 0.0162 3.6 0.93 103.3 3.7 102.5 3.8 83 32.6 103.3 3.7 NA
-1MP39A Spot 19 1323 40452 2.2 20.3854 1.4 0.1093 3.3 0.0162 3 0.91 103.4 3.1 105.4 3.3 150.4 32.2 103.4 3.1 NA RIM4
-1MP39A Spot 17 1180 59920 1.5 21.2515 1.5 0.1049 3.6 0.0162 3.3 0.92 103.4 3.4 101.3 3.5 52 34.8 103.4 3.4 NA
-1MP39A Spot 23 1423 72932 2.4 20.7484 1.6 0.1075 4.1 0.0162 3.8 0.92 103.5 3.9 103.7 4 108.9 38.2 103.5 3.9 NA
-1MP39A Spot 35 561 12157 1.6 20.0635 1.9 0.1112 4.1 0.0162 3.6 0.88 103.6 3.7 107.1 4.2