University of Southern California Dissertations and Theses

Paleoenvironments and the Precambrian-Cambrian transition in the southern Great Basin: Implications for microbial mat development and the Cambrian radiation

(USC Thesis Other)

Paleoenvironments and the Precambrian-Cambrian transition in the southern Great Basin: Implications for microbial mat development and the Cambrian radiation

PDF

Download a page range

Download transcript

Copy asset link

Request this asset

Transcript (if available)

Content PALEOENVIRONMENTS AND THE PRECAMBRIAN-CAMBRIAN TRANSITION
IN THE SOUTHERN GREAT BASIN: IMPLICATIONS FOR MICROBIAL MAT
DEVELOPMENT AND THE CAMBRIAN RADIATION
by
Scott Andrew Mata
________________________________________________________________________
A Dissertation Presented to the
FACULTY OF THE USC GRADUATE SCHOOL
UNIVERSITY OF SOUTHERN CALIFORNIA
In Partial Fulfi llment of the
Requirements for the Degree
DOCTOR OF PHILOSOPHY
(GEOLOGICAL SCIENCES)
August 2012
Copyright 2012 Scott Andrew Mata
ii
ACKNOWLEDGEMENTS
First I would like to thank my advisor Dave Bottjer for his encouragement
throughout the years and his never-ending resilience, trust, and patience for adapting to
my ever-changing and evolving research interests. In spite of the shifting focus of my
projects, Dave was always accommodating and understanding, stressing that the key to
achieving success in research is to pursue methods and projects that I would actually fi nd
enjoyable. Dave went above and beyond his role as my research advisor and I can fi nd
no greater compliment than to say that ever since arriving at USC, Dave has always felt
more like a collaborator, a friend, and an equal than just a research advisor.
Second I would like to thank Frank Corsetti for all his advice and guidance on my
research projects, as well as being a role model for my professional development. The
merit and worth of a professor’s advice can be measured by how often they are sought
out by students. With how many people seek out Frank on a given day, he should have
his own receptionist and appointment book. No matter what I do in science, I know that
Frank will always give me a truthful opinion about its content and always provide helpful
suggestions on how to make it even better.
I would also like to acknowledge my committee members for my qualifying
exams and dissertation: Will Berelson, Donn Gorsline, and Wiebke Ziebis. Each was so
integral to shaping my project and being good enough to kindly point out the portions that
veered a bit too far from practicality or possibility. I would like to thank all the friends
who have been through the Paleolab during my time at USC: Rowan Martindale, Lydia
Tackett, Kathleen Ritterbush, Sarah Greene, Carlie Pietsch, Liz Petsios, Kirk Domke, and
Alyssa Bell. Special thanks go to those who accompanied me on my fi eld work to look at
rocks, even though they didn’t have many fossils in them.
iii
Lastly, I would like to thank all my friends and family who were there for me
in non-academic ways. The support and encouragement of my parents paved the way
for my decision to pursue graduate school. I also appreciate that they still press me to
explain my research to them, even though they may not necessarily understand it. I
would also like to thank my brother and sisters and my nephew and nieces for supplying
a pleasant weekend distraction. I also want to express my deepest gratitude to all my
friends that spent time with me during weekday work hours when I needed a break: Adam
Ianno, Katy Johanesen, Katie Harazin, and Emily Mortazavi. It takes a special kind of
friend to be willing to turn Wednesday coffee hour into a two-hour event.

iv
TABLE OF CONTENTS
ACKNOWLEDGEMENTS ii
TABLE OF CONTENTS iv
LIST OF FIGURES v
ABSTRACT ix
CHAPTER I: Introduction 1
CHAPTER II: The self-destructive prograding barrier island model: Evidence 31
from the Precambrian-Cambrian transition of the southern Great Basin, United
States
TABLE 2.1. Sedimentary facies of the Precambrian-Cambrian Transition 39
in the southern Great Basin
CHAPTER III: Lower Cambrian Grand Cycles of the southern Great Basin: 74
Implications for mixed carbonate-siliciclastic systems
CHAPTER IV: Facies control on lower Cambrian wrinkle structure development 91
and paleoenvironmental distribution, southern Great Basin, United States
CHAPTER V: Early Cambrian microbial mat-associated sediment mounds and 128
fun nels from the middle member of the Wood Canyon Formation, southern
Great Basin, United States
CHAPTER VI: Early Cambrian anemone burrows from the upper member of the 155
Wood Canyon Formation, Death Valley region, United States: Paleoecological
and paleoenvironmental signifi cance
CHAPTER VII: Conclusions 187
REFERENCES 193
APPENDIX 221
Locality information 221
v
LIST OF FIGURES
FIGURE 1.1. Illustration of the changes associated with the Cambrian substrate 4
revolution
FIGURE 1.2. Time-environment diagram for the Cambrian-Ordovician 6
FIGURE 1.3. Illustration of the changes associated with the agronomic 7
revolution
FIGURE 1.4. Model of the rifting of western North America in the late 9
Precambrian
FIGURE 1.5. Interfi ngering Precambrian-Cambrian successions of the 10
southern Great Basin
FIGURE 1.6. Map of mountain ranges within the southern Great Basin 12
FIGURE 1.7. Map showing the localities addressed in this dissertation 14
FIGURE 1.8. Lithologic correlation of Precambrian-Cambrian strata in the 21
southern Great Basin
FIGURE 1.9. Current Cambrian chronostratigraphic subdivisions 23
FIGURE 1.10. Biostratigraphy of the upper member of the Wood Canyon 26
Formation
FIGURE 1.11. Biostratigraphic correlation of Precambrian-Cambrian strata in 29
the southern Great Basin
FIGURE 2.1. Depositional model for a prograding barrier island 32
FIGURE 2.2. The Galveston Island prograding barrier island model 34
FIGURE 2.3. Stratigraphic intervals examined for the self-destructive prograding 36
barrier island model
FIGURE 2.4. Stratigraphic section through the upper member of the Wood 47
Canyon Formation
vi
FIGURE 2.5. Facies from the upper member of the Wood Canyon Formation 48
FIGURE 2.6. Stratigraphic section through the Hines Tongue of the Reed 50
Dolomite
FIGURE 2.7. Facies from the Hines Tongue of the Reed Dolomite 51
FIGURE 2.8. Stratigraphic section through the lower member of the Wood 53
Canyon Formation
FIGURE 2.9. Facies from the lower member of the Wood Canyon Formation 54
FIGURE 2.10. Stratigraphic section through the Lower Member of the Deep 56
Spring Formation
FIGURE 2.11. Facies from the Lower Member of the Deep Spring Formation 57
FIGURE 2.12. Stratigraphic section through the Middle and Upper Members of 60
the Poleta Formation
FIGURE 2.13. Facies from the Middle and Upper Members of the Poleta 61
Formation
FIGURE 2.14. The self-destructive prograding barrier island model 66
FIGURE 2.15. Different progradational pathways for barrier island successions 71
FIGURE 3.1. Grand Cycle boundaries in the southern Great Basin 75
FIGURE 3.2. Stratigraphy and depositional environments of Grand Cycle A 78
FIGURE 3.3. Outcrop photographs of the Montenegro bioherm 80
FIGURE 3.4. Sedimentary facies of Grand Cycle A 82
FIGURE 3.5. Shifting archaeocyath paleoecology during Grand Cycle A 84
FIGURE 3.6. Sequence stratigraphy of the Poleta Formation at Poleta Folds 87
FIGURE 4.1. Stratigraphic intervals examined for the wrinkle structures 95
vii
FIGURE 4.2. Stratigraphic section through the middle member of the Wood 97
Canyon Formation showing the distribution of wrinkle structures
FIGURE 4.3. Sedimentary facies of the middle member of the Wood Canyon 99
Formation
FIGURE 4.4. Wrinkle structures from the middle member of the Wood 100
Canyon Formation
FIGURE 4.5. Stratigraphic section through the Campito-Poleta Formation 102
transition showing the distribution of wrinkle structures
FIGURE 4.6. Sedimentary facies and wrinkle structures from the Montenegro 103
Member of the Campito Formation
FIGURE 4.7. Stratigraphic section through the Middle Member of the Poleta 106
Formation showing the distribution of wrinkle structures
FIGURE 4.8. Sedimentary facies and wrinkle structures from the Middle Member 107
of the Poleta Formation
FIGURE 4.9. Stratigraphic section through the Harkless Formation showing the 109
distribution of wrinkle structures
FIGURE 4.10.Sedimentary facies and wrinkle structures from the Harkless 110
Formation
FIGURE 4.11. Facies model for wrinkle structure formation and preservation 120
FIGURE 5.1. Stratigraphic interval examined for sediment mounds and funnels 130
FIGURE 5.2. Stratigraphic section through the middle member of the Wood 132
Canyon Formation showing the location of the examined sediment mound
bedding plane
FIGURE 5.3. Bedding plane of the middle member of the Wood Canyon 134
Formation showing wrinkle structures and sediment mounds
FIGURE 5.4. Bedding plane exposures showing sediment mound and funnel 136
morphologies
viii
FIGURE 5.5. Polished slabs and x-radiographs of isolated funnels and sediment 138
mounds
FIGURE 5.6. Photomicrographs of isolated funnels and sediment mounds 140
FIGURE 5.7. Methods by which sediment mounds can form 142
FIGURE 5.8. Model for sediment mounds within shallow depressions 147
FIGURE 5.9. Interpretations of the sediment mounds within shallow depressions 148
FIGURE 5.10. Ichnotaxonomy of burrows from the middle member of the Wood 151
Canyon Formation
FIGURE 6.1. Stratigraphy of the Death Valley succession and locality map for 157
upper member of the Wood Canyon Formation
FIGURE 6.2. Stratigraphic section through the upper member of the Wood 160
Canyon Formation showing the location of the studied anemone burrows
FIGURE 6.3. Siliciclastic facies of the upper member of the Wood Canyon 161
Formation
FIGURE 6.4. Carbonate facies of the upper member of the Wood Canyon 163
Formation
FIGURE 6.5. Paleoenvironmental reconstruction of upper member Wood 166
Canyon sedimentary facies
FIGURE 6.6. Outcrop photographs of the trace fossil Bergaueria 168
FIGURE 6.7. Outcrop photographs of the trace fossil Conichnus 170
FIGURE 6.8. Outcrop photographs of the trace fossil Dolopichnus 172
FIGURE 6.9. Anemone burrow margins in thin section 174
FIGURE 6.10. Burrowing behavior represented by the anemone burrows of the 178
upper member of the Wood Canyon Formation
ix
ABSTRACT
The Precambrian-Cambrian transition (~542 Ma) is a pivotal time in Earth’s
history and is notable for marking the fi rst skeletonized occurrences of many major
groups of marine organisms. Concurrent with the appearance of these taxa are signifi cant
changes to the marine substrate, shifting from a seafl oor dominated by microbial mats
to one dominated by metazoan bioturbation. In the southern Great Basin, United
States, these changes, however, appear gradual and the early Cambrian is unlike either
the Precambrian or the subsequent Phanerozoic in terms of the unusual coexistence
between seafl oor microbial mats and marine organisms. Essential to understanding
these interactions is a strong paleoenvironmental framework to work within to examine
which environments microbial mats occur in and whether these are or are not the same
environments that metazoan fossils and trace fossils occur in. The purpose of this
dissertation is to develop a rigorous depositional model of Precambrian-Cambrian strata
within the southern Great Basin to document the paleoenvironmental distribution of fossil
taxa, trace fossils, and microbial mat structures to examine the nature of this coexistence
during the early Cambrian. This may shed light on whether metazoans and microbial
mats existed in similar environments or if each occupied a preferred environmental niche
apart from the other.
The Precambrian-Cambrian transition of the southern Great Basin, United
States, consists of mixed-carbonate siliciclastic sedimentary rocks. Prior studies have
emphasized that carbonate rocks were deposited offshore to siliciclastic rocks; however,
reevaluation shows that these mixed carbonate-siliciclastic sedimentary rocks refl ect the
interaction of shallow shelf and shoreface siliciclastic sediments with marginal marine
carbonate sediments of a barrier island coastline. The highest abundance of skeletal
x
material in the form of shell beds, bioherms, and reefs occurs within the nearshore
carbonates of these deposits. These mixed carbonate-siliciclastic strata are arranged into
depositional units termed ‘Grand Cycles’, consisting of a siliciclastic half-cycle overlain
by a carbonate half-cycle. The siliciclastic-carbonate transitions of these depositional
units refl ect a shift in lateral depositional environments, rather than a complete change
in the depositional system, and the progradation events recorded by these Grand Cycles
allow for assessment of paleoecological patterns over short timescales. Notably, the
early Cambrian archaeocyathans expand their ecological niche from their fi rst occurrence
as level bottom forms to bioherms to reefs over the course of a single progradational
episode—less than 6 million years in duration—despite remaining in an onshore position,
within the shoreface or landward of it.
Lower Cambrian wrinkle structures within siliciclastic strata show a preference
for occurring within heterolithic deposits, primarily of offshore transition and mixed tidal
fl at environments. These microbial mat features refl ect the absence of a surface mixed
layer, which is in accord with previous observations that early Cambrian bioturbation was
shallow-penetrating and not extensive within siliciclastic shelf environments. In contrast,
siliciclastic tidal fl at environments from the middle member of the Wood Canyon
Formation record deep-penetrating burrows associated with wrinkle structures; however,
bioturbation levels are generally low (ii = 2). Lagoonal environments of the upper
member of the Wood Canyon Formation record similar deep-penetrating burrows that are
attributable to anemones and exhibit higher levels of bioturbation (ii = 4).
The nature of the lower Cambrian trace fossil and microbial mat record shows
that nearshore environments were characterized by deep-penetrating burrows in
lagoonal and tidal fl at environments, while shallow shelf environments exhibit shallow-
penetrating levels of bioturbation and a prevalence of microbial mat structures. These
distributions—coupled with the high abundance of shell beds, bioherms, and reefs
xi
in onshore environments—parallel previous observations that evolutionary novelties
within Cambrian-Ordovician communities show a preference for onshore origination.
Paleoenvironmental distributions of microbial mats and metazoans in the earliest
Cambrian in the southern Great Basin suggest that metazoan taxa and deep-penetrating
bioturbation dominated in the onshore position throughout much of this interval,
while microbial mat substrates and shallow penetrating bioturbation were typical for
shelf environments, a distribution opposite to that found across shelf and nearshore
environments today.
1
CHAPTER I
Introduction
Purpose of Study
The Precambrian-Cambrian transition marks a very unique interval in Earth’s
history and records an unusual juxtaposition of microbial-infl uenced and metazoan-
infl uenced ecosystems within shallow shelf and nearshore settings. This contrasts
with the preceding Precambrian, which was notably microbial mat-dominated, and
the remainder of the Phanerozoic, which is characterized by a surface mixed layer
perpetuated by metazoan bioturbation in most oxygenated shallow shelf environments.
The temporal overlap of these two ecosystems is best represented in the southern Great
Basin where mixed carbonate-siliciclastic sedimentary rocks provide an excellent record
of latest Precambrian and earliest Cambrian body fossils, trace fossils, and microbial mat
structures. While the sedimentary successions of the southern Great Basin have been
studied in detail with regard to stratigraphy (e.g., Nelson, 1962; Stewart, 1970; Nelson,
1978; Corsetti and Hagadorn, 2003), trace fossils (Nelson, 1978; Langille, 1974; Corsetti
and Hagadorn, 2003), extent of bioturbation (Droser and Bottjer, 1989; Marenco and
Bottjer, 2008), and the distribution of microbial mat structures (Hagadorn and Bottjer,
1997; 1999), there has yet to be a high-resolution paleoenvironmental analysis capable of
uniting the current understanding of microbial mat structures, bioturbation, and emerging
metazoan taxa into a comprehensive depositional and stratigraphic framework. Such a
framework would allow for the examination of the interactions, or lack thereof, between
microbial mats and metazoans during the earliest Cambrian, and may provide new
insights into the physical and biological controls on microbial mat distribution.
Mixed carbonate-siliciclastic sedimentation is poorly represented in modern
settings. Previous interpretations of the depositional environments of the Precambrian-
2
Cambrian transition in the southern great Basin have been hindered by a lack of adequate
depositional models capable of explaining the mixing of carbonate and siliciclastic
sediments within these strata. Prior depositional models restricted siliciclastic
sedimentation to onshore environments landward of offshore carbonate sedimentation.
Recent studies (e.g., Mata et al., in review), however, provide a new solution to the
problem and reverse this pattern, placing siliciclastic sediments in shallow shelf and
shoreface environments and carbonate sediments within back-barrier environments of
a barrier island coastline. Utilizing this new template for mixed systems, depositional
environments of the southern Great Basin can be reinterpreted and their constituent body
fossils, trace fossils, and microbial mat structures can be more accurately contextualized.
This study uses high-resolution sedimentology and stratigraphy to redefi ne late
Precambrian and early Cambrian mixed carbonate-siliciclastic systems with the purpose
of examining the spatial relationship between microbial mat structures, body fossils, and
trace fossils during the early Cambrian. To this aim, this chapter presents an overview of
the concepts, geological framework, and age constraints upon which subsequent chapters
are based. It will review the notable physical and biological events of the Precambrian-
Cambrian transition and provide a general outline of the geological setting of the southern
Great Basin, its regional stratigraphy, a brief history of study in the region, and the
constraints upon which the events of the Precambrian-Cambrian transition are correlated
across the region.
The Precambrian-Cambrian Transition and Major Events
The Precambrian-Cambrian transition (~542 Ma) is a signifi cant interval
in Earth’s history and is hallmarked by notable changes in the nature of the marine
environment from Precambrian seafl oors dominated by microbial mats to seafl oors
dominated by metazoan bioturbation in the later Cambrian and subsequent Phanerozoic
3
(Fig. 1.1) (Droser and Bottjer, 1989; Seilacher and Pfl üger, 1994; Hagadorn and Bottjer,
1997; 1999; Bottjer et al., 2000). This fundamental shift in the nature of the marine
substrate during this interval has been termed the “agronomic revolution” (Seilacher and
Pfl üger, 1994). The transition between these seafl oor ecosystems is best represented
in the early Cambrian, when deeper-tiered (> 6 cm depth) burrowing had yet to make
a signifi cant impact in marine shelf settings, and would not do so until the Ordovician
(Bottjer and Ausich, 1986).
Associated with the changes in the marine substrate through the Precambrian-
Cambrian transition were changes in ecology and the emergence and diversifi cation
of a wide range of skeletonized metazoans. The early Cambrian saw the emergence
of calcareous body fossils for a wide range of phyla, including sponges (e.g.,
archaeocyathans), mollusks, brachiopods, echinoderms, and arthropods. The appearance
of these calcareous body fossils, however, likely does not refl ect the emergence of these
phyla, and several major groups are preceded by soft bodied forms or by trace fossils.
Mollusk shells fi rst appeared in the earliest Cambrian, while the phylum may have
been preceded by Ediacaran soft-bodied precursors, such as the body fossil Kimberella,
which is associated with purported radular scratch marks of the trace fossil Radulichnus
(Seilacher and Hagadorn, 2010). Trilobites, the earliest skeletonized arthropods, are
preceded by trace fossils with distinctive bi-lobed morphologies and scratch marks,
which hint at their presence prior to the preservation of their body fossils (Crimes, 1992).
The emergence of each of these phyla as calcareous body fossils is not synchronous, but
rather is staggered throughout the early Cambrian, so much so that the appearance of
certain groups has been used to mark boundaries between Cambrian Series and Stages
(e.g., Babcock and Peng, 2007). Of note are the fi rst occurrence of archaeocyathans and
the fi rst occurrence of trilobites, which mark the start of the second and third stages of the
Cambrian, respectively.
4
FIGURE 1.1. Diagramatic reconstruction of the changes in the marine substrate from
the latest Precambrian through the post-Cambrian Phanerozoic and the prevailing
processes associated with this transition. Note that there is a dramatic swing from latest
Precambrian substrates in which microbial processes yield to metazoan processes as the
major biological infl uence on the marine substrate. Modifi ed from Bottjer et al. (2000).
5
The use of high resolution paleoenvironmental analysis has been applied
previously to the Cambrian radiation of trace fossil diversity (e.g., MacNaughton and
Narbonne, 1999; McIlroy and Logan, 1999) and extent of bioturbation (e.g., Droser and
Bottjer, 1993), as well as the emergence of evolutionary novelties (e.g., Jablonski et
al., 1983). Of note is that many evolutionary novelties and new styles of bioturbation
show an onshore origination followed by offshore expansion. Time-environment
analysis of Cambrian-Ordovician marine benthic communities by Jablonski et al. (1983)
showed that evolutionary novelties and ecological innovations show a preference for
onshore origination followed by offshore expansion (Fig. 1.2). This pattern is echoed
in Cambrian-Ordovician extent of bioturbation, measured by ichnofabric indices (sensu
Droser and Bottjer, 1986), within carbonate ramp settings of the Great Basin, USA
(Droser and Bottjer, 1993). Ichnofabric indices increase in all shelf environments
through the Cambrian and Ordovician; however, inner shelf settings consistently have
higher ichnofabric indices than middle or outer shelf settings. The overall pattern
refl ects offshore expansion of extensive bioturbation, and offshore environments trail the
increasing ichnofabric index trend of onshore environments.
Matground ecosystems of the latest Precambrian and earliest Cambrian are
notable for their effects on the ecology of inhabitant organisms. The Ediacaran biota has
been suggested to contain several microbial matground-based ecologies including mat
encrusters that lived attached to the mat; matstickers that lived embedded within the mat,
possibly growing upward with it; mat scratchers that grazed the mat without destroying
it; and undermat miners that burrowed beneath the mat, possibly consuming degrading
organic matter (Fig. 1.3) (Seilacher, 1999). These ecologies become scarce or non-
existent following the development of extensive vertical bioturbation in the Cambrian and
Ordovician. Earliest Cambrian animals still exhibit adaptations to a microbial mat-based
lifestyle, including several forms of echinoderms that likely lived as sediment stickers
6
FIGURE 1.2. Time-environment diagram for the Cambrian-Ordovician showing
the temporal range and environmental distribution of marine shelf communities. A)
The vertical ranges of the boxes in the diagram represent the temporal scale, and the
horizontal ranges represent the environmental distribution. Boxes are subdivided into 4
clusters representing similar ordinal level faunas. Note that these clusters of faunas show
an onshore origination followed by offshore expansion. B) Representative faunas that
comprise the four clusters found in A. Modifi ed from Jablonski et al. (1983) and Sepkoski
and Sheehan (1983).
7
FIGURE 1.3. The widespread development of seafl oor microbial mats led to unique
ecologies during the Ediacaran. These ecologies became scarce or non-existent followed
the development of deep-penetrating vertical bioturbation during the Cambrian, which
resulted in the demise of matgrounds by the Ordovician. Modifi ed from Seilacher (1999).
8
within a fi rm substrate (Bottjer et al., 2000). By the late Cambrian, however, these
adaptations had largely disappeared and were replaced by adaptations for softground
substrates. In the case of the echinoderms, lifestyles included either living attached to
hard substrates or utilizing root-like holdfasts for anchoring in soft substrates (Bottjer et
al., 2000).

Precambrian-Cambrian Tectonic Setting and Paleogeography of the Southern Great Basin
Proterozoic-Cambrian strata in east-central California were deposited on a
thermally subsiding passive margin that formed following late Neoproterozoic rifting and
existed seaward of a stable cratonic platform (Stewart, 1970; Armin and Mayer, 1983;
Picha and Gibson, 1985). This rifting led to the reshaping of the continental margin and
the development of a laterally extensive miogeocline that extended from Sonora, Mexico
to the Yukon Territory, Canada (Fig. 1.4) (Stewart and Suczek, 1977; Bond et al., 1985).
In the southern Great Basin, these sediments deposited across this subsiding margin
were preserved as a succession of strata that thicken from the southeast to the northwest.
This northwestward thickening is attributable to asymmetric subsidence across the rifted
margin (Fedo and Cooper, 2001). Precambrian-Cambrian strata of the southern Great
Basin are divided into four interfi ngering successions that span the ancient craton to shelf,
and included from onshore to offshore are the Craton, Mojave (Craton Margin), Death
Valley, and White-Inyo successions (Fig. 1.5) (Nelson, 1978; Corsetti and Hagadorn,
2000; Fedo and Cooper, 2001). Death Valley and White-Inyo successions are thickest
because they represent deposition within the rapidly subsiding miogeocline, while
craton and craton margin successions are generally thinner and much of the Precambrian
sedimentary record is not preserved. These successions are distributed throughout the
southern Great Basin as a laterally extensive Neoproterozoic-Cambrian outcrop belt that
is represented at numerous mountain ranges (Fig. 1.6). Focus is placed primarily on the
9
FIGURE 1.4. Diagramatic model of the rifting and reshaping of the western continental
margin of North America. Rifting began in late Precambrian time and led to the
development of a thermally-subsiding passive margin across which the Precambriam-
Cambrian succession of western North America was deposited. Modifi ed from Stewart
and Suczek (1977).
10
100 km
Las
Vegas
CA
NV
Bishop
White
Inyo
Death
Valley
Craton Margin
Cratonal
FIGURE 1.5. Map of the southern Great Basin showing the four interfi ngering
successions deposited during the Precambrian-Cambrian transition. The Cratonal
succession represent the most onshore locality, and the White-Inyo succession represents
the most offshore locality. Modifi ed from Corsetti and Hagadorn (2000).
11
Death Valley and White-Inyo successions that represent the paleogeographic proximal
shelf and proximal-middle shelf, respectively (Corsetti and Hagadorn, 2000; 2003).
Figure 1.7 shows the distribution of specifi c localities examined for this study.
Southern Great Basin Stratigraphy
The following sections outline the regional sedimentology and stratigraphy of
the Precambrian-Cambrian transition in the southern Great Basin to provide a broad
overview of the rock units discussed in subsequent chapters. Also is detailed the
constraints on the lithologic and biostratigraphic correlation of strata across the region.
Death Valley Stratigraphy
The Proterozoic to lower Cambrian succession of the Death Valley region is
comprised of, in stratigraphic order, the Pahrump Group, Noonday Dolomite, Johnnie
Formation, Stirling Quartzite, Wood Canyon Formation, Zabriskie Quartzite, and Carrara
Formation. The Pahrump Group represents the lowermost Neoproterozoic strata in
the region and rests atop crystalline basement rock. This basement rock consists a of
gneiss and schist metamorphic complex that was last metamorphosed and deformed
approximately 1.8-1.7 Ga, with quartz monzonite and granite intruding the complex
around 1.4 Ga (Wasserburg et al., 1959; Lanphere, 1964; Labotka, 1978; Labotka et
al., 1980). The Pahrump Group is subdivided into the Crystal Spring Formation, Beck
Spring Dolomite, and the Kingston Peak Formation. The Pahrump Group formed
largely prior to late Neoproterozoic rifting of the western margin of North America and
represents a different tectonic setting than overlying strata. Subsequent strata formed on
a thermally subsiding passive margin, the Cordilleran Miogeocline of Stewart (1972),
resulting in a westward-thickening package of sedimentary rocks that extends laterally
along the western margin of North America.
12
CA
NV
118°W
116°W
36°N
40 km
Owens Valley
Valley
Death
Sho-
shone
Beatty
Salt Spring
Hills
Saline
Panamint
Range
Resting
Spring
Range
Nopah
Range
Inyo Mountains
Black
Mts.
Grapevine Mts.
Funeral
Mts.
Kingston
Range
Last Chance Range
Valley
White Mountains
Silver
Peak
Range
Weepah
Hills
Mt. Dunfee
Bare
Mt.
Eagle
Mt.
Saddle Peak
Hills
FIGURE 1.6. Map showing the approximate locations of Precambrian-Cambrian
outcrop belt in the southern Great Basin (shaded in gray) and the names of the respective
mountain ranges in which they are found (after Mount et al., 1991; Corsetti and
Hagadorn, 2003).
13
Within the Pahrump Group, the Crystal Spring Formation is a very heterolithic
unit and consists of quartzite, conglomerate, shale, chert, and stromatolitic limestone
and dolostone (Stewart, 1970; Awramik et al., 2000). A diabase sill intrudes the Crystal
Spring Formation and has a radiometric age of 1.08 Ga (Heaman and Grotzinger, 1992);
however, this age is likely not to be representative of the entire formation.
The Beck Spring Dolomite is comprised predominantly of dolostone, with lesser
chert, and has traditionally been subdivided into three informal members: the lower
laminated member, the middle oolitic member, and the upper stromatolitic member
(e.g., Gutstadt, 1968). Recent work by Harwood and Sumner (2011) based on new
observations proposed revisions to the names of the members and suggested a lower
laminated member, a middle thrombolitic/breccia member, and an upper cherty member.
The overlying Kingston Peak Formation consists of interbedded sandstone
and siltstone with prominent and distinctive diamictite deposits found within the
lower portion and upper portion of the formation (Prave, 1999). These diamictites are
comprised of fi ne-grained mudstone with carbonate clasts, large olistoliths, and sediment
clasts from of the underlying Crystal Spring Formation and diabase sill found within it
(Prave, 1999).
The Noonday Dolomite overlies the Pahrump Group and consists of coarsely
crystalline dolostone with a lower ‘algal member’ and an upper ‘sandy member’, with
the lower member containing stromatolitic megadomes with tube structures (Cloud et al.,
1974; Stewart, 1970; Corsetti and Grotzinger, 2005). The upper member contains similar
stromatolitic textures, but smaller domes. These two members are separated across an
erosional unconformity and sequence boundary marked by sand-fi lled grikes (Corsetti
and Kaufman, 2005). This sequence boundary is overlain by pebble conglomerates of the
transient Ibex Formation that yields upward into the upper sandy member of the Noonday
Dolomite, forming potentially a depositional sequence (Corsetti and Kaufman, 2005).
14
CA
NV
118°W
116°W
36°N
40 km
Owens Valley
Valley
Death
Sho-
shone
Beatty
Southern
Salt Spring
Hills
Saline
Valley
Poleta Folds
Cedar
Flat
Emigrant
Pass
Chicago
Pass
Hines
Ridge
Study localities
FIGURE 1.7. Precambrian-Cambrian outcrop belt in the southern Great Basin (shaded in
gray) and the distribution of localities examined as part of this study (after Mount et al.,
1991; Corsetti and Hagadorn, 2003).
15
The overlying Johnnie Formation is a very heterolithic succession consisting of
interbedded siltstone, sandstone, limestone, and dolostone with stromatolitic intervals
throughout. Siliciclastic strata consist of hummocky cross-stratifi ed and quasi-planar
laminated sandstone interbedded with siltstone. An approximately 2 m thick distinctive
oolitic grainstone occurs near toward the top of the Johnnie Formation, termed the
‘Johnnie oolite’, and serves as a distinctive marker bed that can be traced throughout the
Death Valley region (Stewart, 1970; Summa, 1993). Carbonate seafl oor precipitates also
occur within the upper Johnnie Formation, and are found above the Johnnie oolite (Pruss
et al., 2008). A broad incised valley occurs within the uppermost Johnnie Formation,
showing up to 120 meters of relief, and terminates at the base of the overlying Stirling
Quartzite (Clapham and Corsetti, 2005).
The Stirling Quartzite consists of interbedded quartzite and conglomeratic
quartzite with lesser siltstone and dolostone (Stewart, 1970). The Stirling Quartzite has
been divided into fi ve members by Stewart (1966), members A, B, C, D, and E. Wertz
(1982) notes that the arrangement of members within the Stirling Quartzite is such that
the upper and lower portions of the formation are approximate mirror images of each
other with respect to rock characteristics. Members A and E consists of coarse grained
quartz sandstone, while members B and D consist of interbedded sandstone and siltstone
that is transitional with member C. Of particular note for biostratigraphy is member D.
Member D is the only member to contain signifi cant amounts of dolostone and the only
member to yield body fossils, which were originally documented by Langille (1974)
and subsequently reinterpreted as potentially being the terminal Ediacaran body fossil
Cloudina by Corsetti and Hagadorn (2000).
The Wood Canyon Formation has been informally divided into a lower, middle,
and upper member by Stewart (1966) and consists of interbedded siltstone, sandstone,
conglomerate, and dolostone. The lower member consists of interbedded siltstone and
16
quartz sandstone with three prominent dolostone subunits (Diehl, 1979). The trace fossil
Treptichnus pedum and the Precambrian-Cambrian boundary (~542 Ma) fall between
the second and third of these dolostone subunits (Corsetti and Hagadorn, 2000). The
earliest arthropod trace fossils occur at the very top of the lower member, above the
third dolostone subunit. The middle member contains a generally upward-fi ning trend
throughout and its lower portion consists of conglomerate and conglomeratic quartzite
with interbedded siltstone, and its upper portion consists primarily of interbedded
quartzite and siltstone. The upper member is similar to the lower member and consists
of interbedded siltstone and sandstone with prominent dolostone subunits (Stewart,
1966; Diehl, 1979). In the more onshore southern Nopah Range, only a single dolostone
subunit exists, while in more offshore localities there are additional dolostone subunits
that are separated by siliciclastic facies. Trilobites mark their fi rst occurrence within
siliciclastics of the upper member, below the fi rst dolostone subunit. Archaeocyaths also
occur in the upper member and are found within the lowermost dolostone subunit.
The Zabriskie Quartzite is divided into two members, the Resting Springs
Member and the Emigrant Pass Member. The Resting Springs Member consists of a
general upward-coarsening succession that transitions upward from interbedded siltstone
and sandstone into pebbly sandstone (Prave, 1992). The Emigrant Pass member also
consists of interbedded sandstone and siltstone, but is lacking in pebbly sandstone (Prave,
1992). The Zabriskie Quartzite is notable because it contains the fi rst and only extensive
development of Skolithos pipe rock in the region (e.g., Droser, 1991).
The Carrara Formation has been divided by Palmer and Halley (1979) into nine
members, although only the lowermost fi ve are lower Cambrian in age. The Carrara
Formation members alternate between predominantly shale and predominantly limestone.
Limestone members consist of bioclastic, oncoidal, and pelletal carbonates. The top of
17
the lower Cambrian is found within the Pyramid Shale Member, the fi fth member from
the base. The remainder of the Carrara is middle Cambrian in age (Palmer and Halley,
1979).

White-Inyo Stratigraphy
The Proterozoic to lower Cambrian succession of the White-Inyo region consists
of, in stratigraphic order, the Wyman Formation, Reed Dolomite, Deep Spring Formation,
Campito Formation, Poleta Formation, Harkless Formation, Saline Valley Formation, and
Mule Spring Limestone. The Wyman Formation is comprised of sandstone and siltstone,
with minor oolitic carbonate horizons scattered throughout the upper portion of the unit
(Nelson, 1962; Stewart, 1970; Corsetti and Hagadorn, 2003).
The Reed Dolomite is divided into a Lower Member, Hines Tongue Member,
and an Upper Member. The Lower Member consists of coarsely crystalline dolostone
with interbedded oolite (Nelson, 1962). The Hines Tongue of the Reed Dolomite is a
southward-thickening wedge of mixed siliciclastic-carbonate sedimentary rocks that
contrasts with the predominantly dolostone lithology of the Lower and Upper Members
(Corsetti and Hagadorn, 2003). The Upper Member consists of massive dolostone and
contains the Neoproterozoic body fossil Cloudina (Corsetti and Hagadorn, 2003), which
is known to be latest Ediacaran in age (Grant, 1990).
The Deep Spring Formation is divided into three members, the Lower Member,
Middle Member, and Upper Member. In the White-Inyo Mountains, the Precambrian-
Cambrian boundary is found within the Upper Member and is marked by the fi rst
occurrence of the trace fossil Treptichnus pedum (Corsetti and Hagadorn, 2003).
Each member of the Deep Spring Formation consists of mixed siliciclastic-carbonate
sedimentary rocks, with shallow shelf and shoreface siliciclastics typically being overlain
18
by peritidal carbonates within each member (Mount and Signor, 1985; Corsetti and
Hagadorn, 2003).
The Campito Formation is divided into a lower Andrews Mountain Member
and an upper Montenegro Member (Nelson, 1962). The Andrews Mountain Member
consists primarily of quartz sandstone interbedded with lesser siltstone and shale.
The Montenegro Member consists primarily of siltstone and shale with lesser
quartz sandstone. Near the top of the Montenegro Member are found interbedded
archaeocyathan limestone layers and lenses (Nelson, 1962).
The Poleta Formation was originally subdivided by Nelson (1962) into a Lower
Member and Upper Member; however, McKee and Moila (1962) and subsequent workers
split the Upper Member of Nelson (1962) into a Middle Member and Upper Member.
The Lower Member of the Poleta Formation consists of calcimicrobial-archaeocyathan
limestone and oolite with lesser siltstone. The Middle Member consists of interbedded
shale, siltstone, and quartz sandstone with lesser discrete limestone intervals. The Upper
Member consists of silty and cherty limestones, similar to those of the Middle Member,
and is distinguished from the Lower Member by the absence of archaeocyathans (McKee
and Moila, 1962).
The Harkless Formation consists of shale, siltstone, and quartz sandstone with
interbedded archaeocyathan limestone lenses near its base and top (Nelson, 1962). It
is differentiated from the Middle Member of the Poleta Formation by the presence of
these archaeocyathan limestone lenses and the presence of pisolitic limestone beds
(McKee and Moila, 1962). The Harkless Formation is overlain by the formally defi ned
Saline Valley Formation; however, in Esmeralda County, Nevada, the Saline Valley
Formation is not recognized and the Harkless Formation is directly overlain by the Mule
Spring Limestone. In the White-Inyo Mountains, the Saline Valley Formation consists
of interbedded quartz sandstone, shale, and limestone. The Mule Spring Limestone is
19
the uppermost lower Cambrian unit in the White-Inyo region and consists of bedded to
massive limestone, dolostone, and lesser shale (Nelson, 1962).
A Brief History of Stratigraphic Nomenclature in the Southern Great Basin
While many early workers took note of the Precambrian-Cambrian succession
of the southern Great Basin, the fi rst study to recognize the signifi cance of these rocks
was that of Walcott (1908), which described a lower Cambrian set of strata from
Waucoba Springs in Esmeralda County, Nevada. Edwin Kirk was the fi rst to begin the
naming of rock units within the White-Inyo Mountains and described and named the
Reed Dolomite, Deep Spring Formation, and Campito Formation in Knopf (1918). The
Reed Dolomite was named after its exposures near Reed Flat in the Blanco Mountain
quadrangle, the Deep Spring Formation after exposures within Deep Spring Valley,
and the Campito Formation from exposures on Campito Mountain (Knopf, 1918). The
underlying Wyman Formation, which forms the oldest sedimentary unit in the White-Inyo
Mountain succession, was described and named by Maxson (1935) for exposures inside
Wyman Canyon within the Blanco Mountain quadrangle. The uppermost portion of the
lower Cambrian succession in the White-Inyo Mountains was described and named by
Nelson (1962), which defi ned the Poleta Formation, Harkless Formation, Saline Valley
Formation, and Mule Spring Limestone, and it was recognized that these units overly the
previously defi ned Campito Formation.
The succession in the Death Valley region was known about by early workers;
however, some of the earliest descriptions were those of Nolan (1929), which described
and named the Johnnie Formation, Stirling Quartzite, and Wood Canyon Formation
from exposures near the Johnnie Mine in the Spring Mountains, Nevada. The Johnnie
Formation was named after Johnnie Wash, located near the Johnnie mine; the Stirling
Quartizite was named for exposures on Mount Stirling, approximately 8 kilometers east
20
of the Johnnie mine; and the Wood Canyon Formation was named after Wood Canyon,
approximately 6.5 kilometers south of Crystal Springs within the western portion of the
Spring Mountains. Hazzard (1937) correlated these strata from the Spring Mountains
to the Nopah Range and defi ned a Zabriskie Quartzite Member within the upper portion
of the Wood Canyon Formation, with additional overlying strata being placed within
the Wood Canyon Formation. This member was later raised by Wheeler (1948) to be
its own formation, thus redefi ning the top of the Wood Canyon Formation at the base
of the Zabriskie Quartzite. Removing the Zabriskie Quartzite Member from the Wood
Canyon Formation of Hazzard (1937) left additional strata above the Zabriskie Quartzite
that no longer belonged to the Wood Canyon Formation. These strata were subsequently
reexamined by Cornwall and Kleinhampl (1961), combined with the Cadiz Formation of
Hazzard (1937), and renamed the Carrara Formation.
Lithostratigraphic Correlation
Lithostratigraphic correlation between the Death Valley and White-Inyo
regions is facilitated primarily through the tracing of prominent and distinct carbonate
units across the southern Great Basin (Fig. 1.8). Lithologic correlations below the
Precambrian-Cambrian boundary remain unclear, however, the lithostratigraphic
correlations within the lower Cambrian are generally agreed upon. The second and
third prominent dolostone subunits within the lower member of the Wood Canyon
Formation are lithologically correlated to the limestone units of the Middle Member and
Upper Member of the Deep Spring Formation, respectively (e.g., Stewart, 1982). An
unconformity is found at the base of the Campito Formation and this same unconformity
can be traced to the base of the middle member of the Wood Canyon Formation, marking
a chronostratigraphic surface (Mount, 1982; Mount and Bergk, 1998). The middle
member and lower siliciclastic portion of the upper member of the Wood Canyon
21
Johnnie
Stirling
W ood Canyon
Zabriskie
Carrara
lower
middle
upper
Noonday
Wyman
Reed
Deep
Spring
Campito
Poleta
Harkless
Mule
Spring
ANDREWS
MOUNT AIN
ECHO
CANY ON
NOP AH
RANGE
Crystalline Basement
Base Not
Exposed
km
0
1
Saline V alley
Lower
Middle
Upper
Lower
Hines
Upper
Andrews Mountain
Montenegro
Lower
Middle
Upper
WHITE-INY O SUCCESSION
DEA TH V ALLEY SUC CESSION
?
?
?
?
limestone
shale
dolostone
conglomerate
quartzite
int. sand/siltstone
Correlative carbonate unit
Correlative unit boundary
Correlation uncertain
archaeocyaths
hyoliths
FIGURE 1.8. Lithologic correlation of Precambrian-lower Cambrian strata in the
southern Great Basin (after Nelson, 1962, 1976; Stewart, 1982; Corsetti and Hagadorn,
2000; 2003). Correlation is based upon contacts between formations and carbonate
marker beds. The lowermost dolostone of the upper member of the Wood Canyon in Echo
Canyon is correlated to the single dolostone in the Nopah range based upon the presence
of a distinctive fauna of archaeocyahs and hyolithids. Correlations after Stewart (1982).
22
Formation correlate to the entirety of the Campito Formation. The lowermost dolostone
subunit within the upper member of the Wood Canyon Formation, which contains the
fi rst occurrence of archaeocyathans in the Death Valley region (Stewart, 1966; Stewart,
1970; Mount et al., 1991), correlates to the archaeocyathan-bearing limestones of
the Lower Member of the Poleta Formation. Subsequent upper member dolostone
subunits, lacking in archaeocyathans and found primarily in more offshore localities of
the Death Valley succession, likely correlate to the limestone beds within the Middle
Member and the limestones of the Upper Member of the Poleta Formation that are also
lacking in archaeocyathans. The uppermost siliciclastic facies of the upper member
of the Wood Canyon and the Zabriskie Quartzite correlate roughly to the Harkless
Formation and Saline Valley Formation. The Mule Spring Limestone, which terminates
at approximately the lower-middle Cambrian boundary, correlates to the lower Cambrian
portion of the Carrara Formation, which consists of intertonguing shale and limestone.
Biostratigraphic Correlation
Biostratigraphic correlation throughout the southern Great Basin is based on
several fossil groups due to the staggered emergence of key taxonomic groups (Fig. 1.9).
The latest Ediacaran Period is marked by the occurrence of the earliest biomineralized
body fossil Cloundina, which is thought to be an index fossil for the terminal Ediacaran
(Grant, 1990). The Precambrian-Cambrian Boundary is marked by the fi rst occurrence
of the trace fossil Treptichnus (Phycodes) pedum (Narbonne et al., 1987), which occurs
near the base of the Upper Member of the Deep Spring Formation within the White-Inyo
Mountains (Corsetti and Hagadorn, 2003) and near the base of the third carbonate-capped
parasequence of the lower member of the Wood Canyon Formation in the Death Valley
region (Corsetti and Hagadorn, 2000).
23
Furongian
Series
Cambrian
Series 3
(Undefined)
Cambrian
Series 2
(Undefined)
Terreneuvian
Series
Fortunian
Drumian
Guzhangian
Paibian
Tremadocian
Cambrian Stage 10
(Undefined)
Cambrian Stage 9
(Undefined)
Cambrian Stage 5
(Undefined)
Cambrian Stage 4
(Undefined)
Cambrian Stage 3
(Undefined)
Cambrian Stage 2
(Undefined)
Lower
SERIES STAGES
Cambrian
Ordovician
SYSTEMS
BOUNDARY HORIZONS (GSSPs)
OR PROVISIONAL STRATIGRAPHIC TIE POINTS
Ediacaran
FAD of Iapetognathus fluctivagus (GSSP)
FAD of Lotagnostus americanus
FAD of Agnostotes orientalis
FAD of Glyptagnostus reticulatus (GSSP)
FAD of Lejopyge laevigata
FAD of Ptychagnostus atavus
?FAD of Oryctocephalus indicus
?FAD of Olenellus or Redlichia
FAD of trilobites
?FAD of SSF or archaeocyathan species
?FAD of Treptichnus pedum (GSSP)
488 Ma
542 Ma
FIGURE 1.9. Chart showing the current Cambrian chronostratigraphic subdivisions.
Four series are defi ned and subdivided into ten stages. At presernt, not all series or stages
have been defi ned and are currently tentatively named by numerical order. Modifi ed from
Babcok and Peng (2007).
24
The Cambrian Period is currently subdivided into four series that are subdivided
into a total of ten stages (e.g., Babcock and Peng, 2007). The fi rst four stages are roughly
equivalent to the previously defi ned Lower Cambrian. The fi rst series, consisting of the
fi rst two stages, is termed the Terrenuvian. The fi rst stage of the Cambrian, the Fortunian,
begins at the fi rst occurrence of Treptichnus pedum, which also marks the Precambrian-
Cambrian boundary. The second stage begins at the fi rst occurrence of small shelly
fossils or archaeocyathans. Cambrian Series 2, which consists of the third and fourth
stages, begins at the fi rst occurrence of trilobites, which is within the Andrews Mountain
Member of the Campito Formation in the White-Inyo region (Nelson, 1978) and the
upper member of the Wood Canyon Formation in the Death Valley region (Mount et al.,
1991). The boundary between stages 3 and 4 occurs at the fi rst occurrence of the trilobite
genera Olenellus or Redlichia. Of these two taxa, only Olenellus occurs, and is found
fi rst near the top of the Middle Member of the Poleta Formation in the White-Inyo region
(Nelson, 1978) and the top of the upper member of the Wood Canyon Formation (Mount
et al., 1991). The end of Cambrian Series 2, which roughly equates to the Lower-Middle
Cambrian Boundary, begins at the fi rst occurrence of the trilobite species Oryctocephalus
indicus. While this species of trilobite has not been recognized in the southern Great
Basin, the genus does occur within the Monola Formation (Nelson, 1978), well above the
Harkless Formation and overlying Mule Spring Limestone in the White-Inyo Mountains,
and within the upper portion of the Pyramid Shale Member of the Carrara Formation in
the Death Valley region (Palmer and Halley, 1979).
For higher resolution biostratigraphy of the southern Great Basin strata, trilobite
biozones are utilized. The trilobite biozones used for southern Great Basin strata were
originally defi ned by Fritz (1972) for the MacKenzie Mountains of northwestern Canada,
but were later correlated to the remainder of the North American Cordillera by Fritz
(1975). Three lower Cambrian biozones are recognized and span the extent of Cambrian
25
Series 2. From oldest to youngest are the Fallotaspis biozone, Nevadella biozone, and
Bonnia-Olenellus biozone. Fritz (1975) was the fi rst to apply these biozones to the
White-Inyo Mountains; however, Nelson (1976) modifi ed their placement, as well as
documenting additional biostratigraphic constraints. The defi ned Fallotaspis biozone
of Nelson (1976) within the White-Inyo Mountains extends from the middle Andrews
Mountain Member to the middle Montenegro Member of the Campito Formation.
Hollingsworth (2005) revised the placement of the base of this Fallotaspis biozone,
based on a reevaluation of prior trilobite claims, and placed it close to the contact
between the Andrews Mountain Member and the Montenegro Member. The Nevadella
biozone extends from middle Montenegro Member to the upper Middle Member of the
Poleta Formation. The Bonnia-Olenellus biozone includes the remainder of the Poleta
Formation, the entirety of the Harkless Formation, and terminates at the top of the Mule
Spring Limestone (Nelson, 1978).
In the Death Valley region Hunt (1990) and Mount et al. (1991) placed the start
of the Fallotaspis biozone within the upper member of the Wood Canyon Formation
(Fig. 1.10). In the Grapevine Mountains the upper member of the Wood Canyon
Formation consists of predominantly siliciclastic sedimentary rocks with two prominent
dolostone intervals, a lower and upper, that are found within subunits B-C and E of
Hunt (1990), respectively. The base of the Fallotaspis biozone of Hunt (1990) was
placed within siliciclastics below the lower dolostone interval (subunits B-C), but above
the fi rst occurrence of trilobites; these siliciclastics included the identifi ed trilobite
Fallotaspis. The boundary between the Fallotaspis and Nevadella biozones was placed
within siliciclastics above the lower dolostone interval, just above the occurrence of
Acephalacanthus, but below the fi rst occurrence of Nevadella. The boundary between
the Nevadella biozone and Bonnia-Olenellus biozone was placed within the uppermost
siliciclastics of the upper member, marked by the fi rst occurrences of Olenellus and
26
subunit A
subunit B
subunit C
subunit D
subunit E
subunit F
0 m
100
200
300
Zabriskie
{Nevadella} (upper Nevadella biozone)
[Cirquella] (lower Nevadella biozone) {formerly Acephalacanthus}
[New unnamed genus] (upper Fallotaspis biozone) {formerly Fallotaspis}
{Trilobite fragments}
Archaeocyathans
Hyoliths
Nevadella Biozone
{Olenellus-Wanneria} (lower-middle Bonnia-Olenellus biozone)
[ ] Reassignments of Fritz (1993) and Hollingsworth (2005)
Fallotaspis Biozone
Trilobite biozone boundary
Uncertain trilobite biozone boundary
Bonnia-Olenellus Biozone
Base of member not exposed
Biozone indeterminate
{ } Original identifications of Hunt (1990)
( ) Stratigraphic range of identified trilobites
FIGURE 1.10. Biostratigraphy of the upper member of the Wood Canyon Formation,
Grapevine Mountains, Death Valley, California. Trilobite occurrences and biozones
are based on Hunt (1990), with reidentifi cations by Fritz (1993) and Hollingsworth
(2005). If biozones are defi ned at the fi rst occurrence of a representative trilobite, then
the Fallotaspis-Nevadella biozone boundary is placed near the base of subunit D. It
is plausible that the boundary could also occur within the interval termed ‘Biozone
indeterminate’ because trilobites have not been recovered from this part of the succession.
27
Wanneria. The Nevadella biozone includes the upper dolostone interval found within
subunit E of Hunt (1990). The Bonnia-Olenellus biozone includes the uppermost
siliciclastic facies of the upper member of the Wood Canyon Formation (Hunt, 1990;
Mount et al., 1991), the entirety of the Zabriskie Quartzite, and extends into the Pyramid
Shale Member of the Carrara Formation (Palmer and Halley, 1979).
Fritz (1993) reevaluated the trilobite identifi cations of Hunt (1990) and noted that
the indentifi ed Acepahalcanthus trilobite used to defi ne the top of the Fallotaspis biozone
is actually the trilobite Cirquella, which is found in the lower Nevadella zone (Palmer
and Repina, 1993). This reassignment shifted the base of the Nevadella biozone in the
Grapevine Mountains to a horizon below this occurrence of Cirquella. Hollingsworth
(2005) also reevaluated the trilobite identifi cations of Hunt (1990) and found that the
identifi ed Fallotaspis is most likely an unnamed new genus that has been recognized
previously in the Montezuma Range of Nevada within the upper Fallotaspis biozone.
Based on these reevaluations, if the base of each biozone is placed at the fi rst occurrence
of a representative trilobite from that zone, as established originally by Fritz (1972) when
defi ning these zones, this would place the Fallotaspis-Nevadella biozone boundary just
above the lowermost dolostone interval, below the occurrence of Cirquella. Utilizing
the fi rst occurrence may not, however, be completely accurate because there is an
approximately 100 meter interval separating the uppermost Fallotaspis biozone trilobite
from the lowermost Nevadella biozone trilobite occurrence. The boundary between
the Fallotaspis biozone and the Nevadella biozone could, therefore, actually occur at
some horizon above the occurrence of the unnamed trilobite genus of the Fallotaspis
biozone, but below the occurrence of Cirquella of the lower Nevadella biozone. This
zone of uncertainty includes the lower dolostone interval and siliciclastic strata just
above and just below it. In spite of this, for clarity, all subsequent discussion will follow
28
the methodology of Fritz (1972) and consider the base of each biozone to be at the fi rst
occurrence of a representative trilobite from that biozone.
The upper member in the southern Nopah Range contains only one dolostone
interval, while offshore localities to the northwest, such as the Funeral Mountains and
Grapevine Mountains, contain at least two prominent dolostone intervals within a
thicker upper member (e.g., Stewart, 1966; Mount and Rowland, 1981; Mount et al.,
1991). The single dolostone interval in the southern Nopah range has a very distinctive
biota of archaeocyaths, hyoliths, and echinoderm ossicles. This dolostone subunit is
typically correlated to the lower dolostone interval in more offshore localities, such as
the Grapevine Mountains, that also contains archaeocyaths, hyoliths, and echinoderm
ossicles (Fig. 1.11) (e.g., Stewart, 1966; Mount and Rowland, 1981; Hunt, 1990; Mount
et al., 1991). The upper dolostone interval from offshore localities appears to pinch-
out to the southeast (e.g., Stewart, 1966) and may not have a correlative unit within the
Nopah Range. In the more onshore southern Nopah Range, the Fallotaspis biozone
is defi ned by Mount et al. (1990) to cover an interval lithologically correlative to the
Fallotaspis biozone in the Grapevine Mountains. This includes the single dolostone
interval in the southern Nopah Range and siliciclastic strata just above and below it.
The Nevadella bizone in the southern Nopah Range is restricted to a thin interval within
siliciclastic facies above the single dolostone interval of the upper member (Mount et al.,
1990). This may be due to low rates of deposition in this more onshore locality, while
more offshore localities have a thicker upper member and a more developed Nevadella
biozone. The base of the Bonnia-Olenellus biozone is traced by Mount et al. (1991) from
the Grapevine Mountains to the uppermost siliciclastic strata within the upper member of
the Wood Canyon Formation.
29
Johnnie
Stirling
W ood Canyon
Zabriskie
Carrara
lower
middle
upper
Noonday
Wyman
Reed
Deep
Spring
Campito
Poleta
Harkless
Mule
Spring
ANDREWS
MOUNT AIN
ECHO
CANY ON
NOP AH
RANGE
Base Not
Exposed
km
0
1
Saline V alley
Lower
Middle
Upper
Lower
Hines
Upper
Andrews Mountain
Montenegro
Lower
Middle
Upper
WHITE-INY O SUCCESSION
DEA TH V ALLEY SUC CESSION
lower Cambrian
Neoproterozoic
Fallotaspis
Nevadella
Bonnia-Olenellus
Crystalline Basement
T. pedum
Cloudina
limestone
shale
dolostone
conglomerate
quartzite
int. sand/siltstone
lower Cambrian
Neoproterozoic
lower Cambrian
middle Cambrian
Nevadella
Bonnia-Olenellus
FIGURE 1.11. Biostratigraphic correlation of Precambrian-lower Cambrian strata in the
southern Great Basin (after Nelson, 1962, 1976; Stewart, 1982; Corsetti and Hagadorn,
2000; 2003). Correlation is based upon trace fossils, body fossils, and trilobite biozones.
Correlations after Hunt, 1990; Mount et al., 1991; Corsetti and Hagadorn, 2003.
30
Dissertation Summary
This study utilizes sedimentology and stratigraphy, trace fossils, extent of
bioturbation, and body fossils to examine the relationships between the Cambrian
radiation and the distribution of microbial mat structures in siliciclastic settings during the
early Cambrian. First is presented a new prograding barrier island model that establishes
a new set of depositional environments for mixed carbonate-siliciclastic strata of the
southern Great Basin. This new model is then utilized to examine the nature of lower
Cambrian Grand Cycles that represent mixed carbonate-siliciclastic sedimentary cycles
and what they reveal about the paleoenvironmental distribution of emerging Cambrian
taxa. This new depositional model is also used to examine the paleoenvironmental
distribution of lower Cambrian wrinkle structures to examine the roles that physical and
biological processes play in restricting the environments in which microbial mats can be
formed and preserved. With an understanding of the physical controls on microbial mat
distribution, the implications of two unique types of bioturbation are examined. First is
documented an early Cambrian tidal fl at system in which vertical burrows and burrow
mounds are associated with wrinkle structures, similar to modern burrows found in
association with microbial mats. Second is documented early Cambrian large vertical
burrows formed within a back-barrier lagoonal environment and their implications for
the emerging locus of vertical bioturbation in the earliest Cambrian. All these factors
together will show that earliest Cambrian environments show a high degree of niche
development for microbial mats and metazoan organisms, and that this environmental
restriction may have played a signifi cant role in allowing for both to coexist in the early
Cambrian.
31
CHAPTER II
The Self-Destructive Prograding Barrier Island Model: Evidence from the Precambrian-
Cambrian Transition of the Southern Great Basin, United States
Introduction
Prograding barrier island systems are rare in the modern ocean. With Holocene
sea level rise, most barrier island coastlines are generally transgressive in nature and
migrate landward (e.g., Fischer, 1961; Penland and Boyd, 1981; Niedoroda et al., 1985),
although there are exceptions (e.g., Bernard et al., 1962; Leatherman, 1983). While
transgressive barrier island examples in the modern ocean outweigh those of prograding
barriers, the reverse is more accurate for ancient examples, and prograding barrier island
deposits dominate (e.g., Kerr, 1977; Ryer, 1977; Plint, 1984; Plint and Walker, 1987). In
the rock record, a typical prograding barrier island deposit is comprised of an upward-
coarsening offshore to shoreface succession that is capped by foreshore deposits and is
abruptly overlain by lagoonal deposits across a sharp contact (e.g., Kerr, 1977; Ryer,
1977; Plint, 1984; Plint and Walker, 1987).
Prograding barrier island models that are based in the modern predict the offshore
to shoreface succession and the capping foreshore deposits (Bernard et al., 1962; Davies
et al., 1971), but fail to explain the superposition of lagoonal deposits over foreshore
deposits without intervening barrier island sediments from backshore dunes and fl ats.
Rather, most prograding barrier island models based in the modern only suggest that
foreshore deposits should be overlain by backshore dunes, and do not explain the
relationship of this progradational barrier island succession to the adjacent lagoonal
succession (Fig. 2.1). It is obvious from the rock record that lagoonal deposits do indeed
overlie foreshore and shoreface sediments, yet there is limited evidence for this on
modern barrier island coastlines. Evidence from the rock record may, however, provide
32
Lower
shoreface
Upper
shoreface
Foreshore
Backshore
Dunes
FWWB
MLWL
MHWL
REGRESSIVE (PROGRADING) BARRIER ISLAND
Hummocky cross-stratified sandstone
Trough cross-stratified sandstone
Planar laminated sandstone
Tabular cross-stratified sandstone
Large-scale trough cross-stratified
sandstone
FIGURE 2.1. Depositional model for a regressive (prograding) barrier island. In this
model, shoreface deposits transition upward into foreshore deposits, which are capped
by backshore deposits and eolian dunes (after Reinson, 1992). While this model predicts
the distribution and stacking of facies on the seaward side of the barrier island, it does
not explain how these deposits relate to those of back-barrier environments, including
lagoons and tidal fl ats. FWWB = fair-weather wave base; MLWL = mean low water
level; MHWL = mean high water level.
33
new insight into the dynamics of barrier island progradation that are not fully represented
in the modern environment.
Since its inception, the Galveston Island prograding barrier island model—based
on Galveston Island, Texas—has become the standard for interpreting regressive barrier
island deposits in the rock record (Fig. 2.2) (e.g., Bernard et al., 1962; Davies et al.,
1971; Reinson, 1979). It predicts that the preserved deposits of a prograding barrier
island in stratigraphic order will be offshore, shoreface, foreshore, and backshore dunes,
with the corollary that lagoonal deposits cap this succession (e.g., Reinson, 1979). No
aspect of the model, however, explains the common superposition of lagoonal deposits
over shoreface and foreshore deposits. Also implicit in the model is that backshore dune
deposits will be preserved capping foreshore deposits. In the rock record, however, this
is rarely ever the case, and in most instances lagoonal deposits abruptly overly foreshore
deposits, without any preservation of backshore dunes (e.g., Kerr, 1977; Ryer, 1977;
Plint, 1984; Plint and Walker, 1987). The Galveston Island prograding barrier island
model, while appropriate for strand-plain systems, is therefore not suffi cient to explain
the stratigraphy or dynamics of barrier island progradation.
The prevalence of transgressive barrier islands in the modern, and the ubiquity of
prograding barrier island deposits in the ancient, has led to a disconnection between the
modern and the ancient, and most prograding barrier island deposits in the rock record
have no notable modern analogue. Rather, understanding ancient prograding barrier
island deposits requires detailed observations from the rock record, coupled with an
understanding of modern barrier island processes.
The Precambrian-Cambrian transition in the southern Great Basin is a mixed
carbonate-siliciclastic succession comprised of nearshore carbonates and offshore
siliciclastics of a continental shelf setting and barrier island coastline (Corsetti and
Hagadorn, 2000; 2003; Mata et al., in review). Siliciclastic sedimentary rocks consist
34
WEATHERED SAND
LAMINATED SAND
BURROWED
SAND
INTERBEDDED SAND AND SILTY CLAY
SILTY CLAY
AND SAND
PLEISTOCENE
BEDROCK
3500 YRS
2000 YRS
1200 YRS
800 YRS
0 1 2 3 km
GALVESTON ISLAND BARRIER ISLAND MODEL
0
5
10
15
m
BARRIER ISLAND SUCCESSION
LAGOONAL
SUCCESSION
GULF OF MEXICO
FIGURE 2.2. Diagrammatic representation of the Galveston Island prograding barrier
island model (after Bernard et al., 1962; McCubbin, 1982). Galveston Island and related
deposits rest unconformably upon Pleistocene bedrock. Dashed lines with ages are
timelines that represent isochronous surfaces. Dashed boxes highlight the stacking pattern
of facies that would be observed in a vertical core through the barrier island and the
back-barrier lagoon. This model illustrates that a prograding barrier island will develop
an upward-coarsening barrier island succession that results from the seaward migration of
facies. A separate lagoonal succession is developed within the back-barrier environment,
and consists primarily of silty clay and sand that exhibit no discernible progradational
stacking pattern. While these two successions are known to develop within the same
barrier island complex, it is unclear how these two successions can stack during
progradation of the barrier island.
35
primarily of storm-dominated, shallow marine successions deposited in environments
from offshore to shoreface, while carbonate sedimentary rocks were deposited primarily
in tidal inlet, lagoon, and tidal fl at environments (Moore, 1976; Mount and Signor,
1985; Mount et al., 1991; Mata et al., in review). Siliciclastic and carbonate facies
are typically combined into carbonate capped parasequences, with upward-coarsening
siliciclastic shallow shelf and shoreface deposits sharply overlain by peritidal carbonates
(e.g., Corsetti and Hagadorn, 2000; Mata et al., in review). Mixed carbonate-siliciclastic
sedimentary rocks of the southern Great Basin have been recently interpreted as a
prograding barrier island coastline, with siliciclastic sedimentary rocks being deposited
across the shoreface and shelf, while carbonates were deposited within tidal inlet and
back-barrier environments (Mata et al., in review). In many Precambrian-Cambrian
successions of the region, however, tidal inlet deposits are absent and back-barrier
deposits directly overlie those of the shoreface across a sharp contact. The purpose of
this study is to detail the transition between shoreface siliciclastic sedimentary rocks
and back-barrier carbonate sedimentary rock of the Precambrian-Cambrian transition of
the southern Great Basin to examine the nature of the contact between these contrasting
lithologies to explain the common superposition of lagoonal deposits over shoreface and
foreshore deposits in the rock record.
Geological Setting and Methods
Precambrian-Cambrian strata of the southern Great Basin are comprised of
a northwestward-thickening succession consisting of mixed siliciclastic-carbonate
sedimentary rocks that were deposited across a thermally subsiding passive margin that
developed following Neoproterozoic rifting of western North America (Stewart, 1970;
Stewart and Poole, 1974; Stewart and Suczek, 1977). Two interfi ngering successions
were examined, the White-Inyo succession and the Death Valley succession (Fig. 2.3).
36
FIGURE 2.3. Generalized regional stratigraphy of the Neoproterozoic-Cambrian
succession in the southern Great Basin showing the stratigraphic intervals examined in
this study (after Nelson, 1978; Stewart, 1982; Corsetti and Hagadorn, 2000; 2003). 1 =
Hines Tongue of the Reed Dolomite; 2 = Lower Member of the Deep Spring Formation;
3 = Middle-Upper members of the Poleta Formation; 4 = lower member of the Wood
Canyon Formation; 5 = upper member of the Wood Canyon Formation.
White-Inyo Mountains
Southern Nopah Range
WHITE-INYO REGION
DEATH VALLEY REGION
Neoproterozoic
Cambrian
3
2
1
5
4
W ood Canyon
Zabriskie
Carrara
lower
middle
upper
Reed
Deep
Spring
Campito
Poleta
Harkless
km
0
1
Lower
Middle
Upper
Lower
Hines
Upper
Andrews Mountain
Montenegro
Lower
Middle
Upper
Conglomerate
Sandstone
Siltstone and sandstone
Limestone
Dolostone
37
From the White-Inyo succession this study examines the Hines Tongue of the Reed
Dolomite, the Lower Member of the Deep Spring Formation, and the Poleta Formation.
From the Death Valley succession, this study examines the lower and upper members of
the Wood Canyon Formation.
The Reed Dolomite is Neoproterozoic in age and is divided into three members,
the Lower Member, the Hines Tongue, and the Upper Member, and was examined at
Hines Ridge in the White-Inyo Mountains (Fig. 1.7). The Lower Member consists of
coarsely crystalline dolostone with interbedded oolite (Nelson, 1962). The Hines Tongue
of the Reed Dolomite is a southward-thickening wedge of mixed siliciclastic-carbonate
sedimentary rocks that contrasts with the predominantly dolostone lithology of the Lower
and Upper Members (Corsetti and Hagadorn, 2003). The Upper Member consists of
massive dolostone and contains the Neoproterozoic body fossil Cloudina (Corsetti and
Hagadorn, 2003), which is known to be latest Neoproterozoic in age (Grant, 1990).
The Deep Spring Formation is divided into three members, the Lower Member,
Middle Member, and Upper Member, and was also examined at Hines Ridge (Fig. 1.7).
At this locality, the Precambrian-Cambrian boundary is found within the Upper Member
and is marked by the fi rst occurrence of the trace fossil Treptichnus pedum (Corsetti
and Hagadorn, 2003). Each member of the Deep Spring Formation consists of mixed
siliciclastic-carbonate sedimentary rocks, with shallow shelf and shoreface siliciclastics
typically being overlain by peritidal carbonates within each member (Mount and Signor,
1985; Corsetti and Hagadorn, 2003). The Lower Member examined in this study occurs
above the fi rst occurrence of Cloudina and below the Precambrian-Cambrian boundary,
and is therefore latest Neoproterozoic in age.
The Poleta Formation is divided into a Lower, Middle, and Upper Member. The
Lower Member consists of calcimicrobe-archaeocyathan boundstone, skeletal limestones,
and oolite, and represents an early Cambrian reef complex (Rowland, 1984; Rowland
38
and Gangloff, 1988). The Middle Member is comprised of shallow shelf and shoreface
siliciclastics interbedded with rare carbonates (Mount and Signor, 1985). The Upper
Member consists primarily of peritidal interbedded cherty limestones and oolite (Moore,
1976). The Middle and Upper members of the Poleta Formation were examined at the
northeast end of the Poleta folds region of the White-Inyo Mountains (Fig. 1.7).
The Wood Canyon Formation is divided into three informal members: the lower
member, middle member, and upper member. This study focuses on the lower and upper
members, which are similar in lithology and consist of carbonate-capped siliciclastic
units (Stewart, 1966; Stewart, 1970; Diehl, 1974; 1979; Corsetti and Hagadorn, 2000).
The Precambrian-Cambrian boundary, marked by the fi rst occurrence of T. pedum, falls
within siliciclastic facies of the uppermost carbonate-capped unit of the lower member
(Corsetti and Hagadorn, 2000). The lower member was examined at Chicago Pass in
the northern Nopah Range. The upper member was observed at Emigrant Pass in the
southern Nopah Range (Fig. 1.7).
Partial stratigraphic sections were measured through these formations at each
locality and facies were defi ned based upon lithology, sedimentary structures, and fossil
content examined at the outcrop, polished slab, and microfacies scale. Upper and lower
contacts of these partial stratigraphic columns are terminated by extensive covered
intervals that limit their extent.
Sedimentary Facies
Ten sedimentary facies are recognized within the stratigraphic successions
examined from the White-Inyo and Death Valley regions (Table 2.1). These facies were
grouped into facies associations that represent similar depositional settings. The two
recognized facies associations are siliciclastic shelf and shoreface and carbonate back-
39
40
barrier, both of which are related to a Precambrian-Cambrian passive margin, barrier
island coastline system (e.g., Mata et al., in review).
Sedimentary Facies - Siliciclastic Shelf and Shoreface Facies Association
Siliciclastic facies encountered include facies A, B, C, and D that defi ne
environments from below storm wave base to above fair-weather wave base. Facies A
consists of quartz siltstone interbedded with thin (< 2cm thick) very fi ne quartz sandstone
or calcareous sandstone layers, laminae, and lenses. Siltstones are planar laminated
to massively bedded, while sandstone beds are planar laminated, low-angle cross-
laminated, or normally graded. Upper and lower boundaries of sandstones layers are
typically sharp. This facies is found within the Middle Member of the Poleta Formation.
Facies A was deposited below storm wave base, but within the zone affected by storm-
generated currents. Siltstones were likely deposited from suspension during fair-weather
conditions, while planar laminated and graded sandstones represent graded rhythmites
that result from sediment gravity fl ows initiated and propagated by storm-generated
currents (Reineck and Singh, 1972; Nelson, 1982).
Facies B is comprised of quartz siltstone interbedded with fi ne to very fi ne
sandstone beds. Siltstones are massively bedded to planar laminated. Sandstone beds
range in thickness from 2 cm to 20 cm and are non-amalgamated. Internal structure of
the sandstone beds consists of hummocky cross-stratifi cation and quasi-planar laminae
(sensu Arnott, 1993), and upper and lower bedding contacts are sharp. Beds typically
thicken and increase in frequency upsection. This facies is found within the Lower
Member of the Deep Spring Formation, the Middle Member of the Poleta Formation, and
the upper member of the Wood Canyon Formation, and typically conformably overlies
sandy siltstones of facies A. Facies B was deposited within the offshore transition,
between fair-weather wave base and storm wave base. Hummocky cross-stratifi cation
41
and quasi-planar lamination result from strong oscillatory motion or combined-fl ow
conditions, respectively, during high-energy storm events (Harms et al., 1975; Duke et
al., 1991; Arnott, 1993). These events erode sand from the shoreface and transport it
offshore (Duke et al., 1991). Storm sands deposited below fair-weather wave base are
commonly non-amalgamated and are interbedded with silt and mud that settle out of
suspension during fair-weather conditions (Harms et al., 1975; Aigner and Reineck, 1982;
Arnott et al., 1995).
Facies C consists of quartz sandstone with hummocky cross-stratifi cation
and quasi-planar laminae, with rare swaley cross-stratifi cation. Beds are frequently
amalgamated and individual hummocky cross-stratifi ed or quasi-planar laminated
beds range up to 1 meter in thickness. This facies is found within the Hines Tongue
of the Reed Dolomite, the Lower Member of the Deep Spring Formation, the Middle
Member of the Poleta Formation, and the lower and upper members of the Wood Canyon
Formation, and typically conformably overlies interbedded sandstone and siltstone of
facies B. Facies C was deposited within the lower shoreface, above fair-weather wave
base within the zone of shoaling swells. Hummocky cross-stratifi cation and quasi-planar
lamination sandstone beds can form in a wide range of environments during storm events,
where strong oscillatory or combined-fl ow conditions are present. They are, however,
most commonly preserved as amalgamated units within the lower shoreface, above fair-
weather wave base, where fi ne grained silts and muds are less likely to be deposited or
preserved (Harms et al., 1975; Arnott et al., 1995).
Facies D is comprised of medium sandstone with small-scale trough cross-
stratifi cation. Sets of cross-stratifi cation are typically less than 10 cm in thickness
and are commonly bidirectional or multidirectional with abundant scour surfaces.
Interbedded sets of planar laminae are sporadically distributed throughout and form a
lesser component within this facies. This facies is found within the Hines Tongue of
42
the Reed Dolomite and the lower and upper members of the Wood Canyon Formation,
and typically conformably overlies hummocky cross-stratifi ed sandstones of facies C.
Facies D was deposited within the upper shoreface or within a wave-dominated ebb-
tidal delta, likely within the build-up and surf zone (e.g., Clifton et al., 1971). Trough
cross-stratifi cation with bidirectional or multidirectional palaeocurrents is suggestive of
a frequently reworked substrate characterized by highly varied fl ow conditions. This
is most typical of the high-energy upper shoreface in which wave activity within the
build-up and surf zone promote onshore and offshore migrating bedforms that deposit
bidirectional or multidirectional cross-stratifi cation through traction (e.g., Clifton et al.,
1971; Greenwood and Mittler, 1985). The rare occurrences of planar laminae within this
facies may be the result of high-energy breaking waves that generate a planar bed in the
outer surf zone or may be the result of swash activity within the foreshore (Clifton, 1969;
Clifton et al., 1971). In the context of a barrier island coastline, this facies also may
represent the ebb-tidal delta of a wave-dominated tidal inlet. Wave-dominated ebb-tidal
deltas are small-scale sand bodies that develop on the seaward side of tidal inlets and are
characterized by small to large-scale bidirectional cross-stratifi cation with planar laminae
locally developed within swash bars (Hayes, 1980).
Sedimentary Facies - Carbonate Back-Barrier Facies Association
Carbonate facies encountered consist of facies E, F, G, H, I, and J that represent
back-barrier environments of an ancient barrier island coastline. Facies E consists of
medium-scale trough and tabular cross-stratifi ed limestone or dolostone. Allochems
include peloids, ooids, or echinoderm fragments. Beds are 10-30 cm thick and have a
complex internal structure consisting of bidirectional cross-stratifi cation with reactivation
surfaces bounded by sharp contacts. Planar laminae within this facies usually comprise
a minor component, but can occasionally make up nearly half this facies in parts, as is
43
the case in the Reed Dolomite. Within the Lower Deep Spring Formation, this facies
contains rare micritic intraclasts near it base. This facies is found within the Hines
Tongue of the Reed Dolomite, the Lower Member of the Deep Spring Formation, and the
upper member of the Wood Canyon Formation, and can conformably or disconformably
overlie facies C or D. Disconformities at the base of this facies are marked by sharp
irregular scours or by large-scale incision into underlying facies. Facies E was deposited
in back-barrier tidal channels, possibly within a fl ood-tidal delta, landward of a tidal
inlet. The medium-scale cross-stratifi cation with reactivation surfaces found within this
facies is typically found within tidal inlets and channels in which alternating ebb and
fl ood currents result in bidirectional paleocurrents (Hubbard et al., 1979; Hayes, 1980).
The clean carbonate lithology suggests that this facies was spatially separated from the
siliciclastic facies of the shoreface and shelf by a barrier (Mata et al., in review).
Facies F consists of massively bedded micritic limestone or dolostone with rare
peloids or ooids. Primary sedimentary structures are generally lacking in Precambrian-
age deposits of this facies, although extensive mottling due to bioturbation is developed
within Cambrian-age deposits. This facies is found within the lower and upper
members of the Wood Canyon Formation, and typically conformably overlies facies E
or disconformably overlies facies D. Facies F was deposited in a back-barrier lagoonal
environment. The absence of physical structures suggests a quiet-water environment
in which physical reworking is negligible and carbonate sediment is deposited from
suspension or precipitated in situ. The common association of this facies with back-
barrier deposits of facies E suggests that it also represents deposition within a back-
barrier environment, most likely a protected lagoon in which currents and waves are
weakly developed.
Facies G is comprised of mud-chip intraclastic limestones. Beds are
approximately 10 cm in thickness and are frequently amalgamated with irregular upper
44
and lower bedding surfaces. Mud-chip intraclasts are mostly 1 mm in their thinnest
dimension and range in length from 1-2 cm. Intraclasts are normally graded within
each bed, and are densely concentrated near the base as a grainstone lag overlain by
intraclastic packstone and wackestone. This facies is found within the Lower Member of
the Deep Spring Formation and conformably overlies trough cross-stratifi ed limestone of
facies E. Facies G was deposited on tidal channel levee and back-levee tidal fl ats. Tidal
channel levees are typically covered with laminated microbial mats that can be ripped
up to form chip-like intraclasts (e.g., Shinn et al., 1969; Hardie and Ginsburg, 1977;
Ginsburg and Planavsky, 2008). When fl ooding tide water overfl ows tidal channel banks
it erodes levee sediments, including microbial mat chips, and deposits them as normally
graded, intraclast-rich deposits that thin away from the channel margin (Shinn et al.,
1969; Hardie and Ginsburg, 1977). Also, levee deposits are commonly irregular and
discontinuous, pinching out laterally (Shinn et al., 1969), similar to the irregularly bedded
intraclastic layers found within this facies.
Facies H consists of thinly interbedded limestone and cherty siltstones.
Limestone makes up the dominant fraction of this facies and limestone beds are organized
into 1-2 cm thick irregular layers and lenses. Siltstone layers are typically mm-scale,
but may range up to 2 cm in thickness. These two lithologies are arranged into fl aser
and wavy bedded sets, with the siltstone forming drapes over the limestone. This facies
is found within the Middle Member of the Poleta Formation where it disconformably
overlies hummocky cross-stratifi ed and quasi-planar laminated sandstones of facies
C. Facies H was deposited on an intertidal fl at within a back-barrier environment.
Flaser and wavy bedding are the result of changes in fl ow strength associated with tidal
currents (Reineck, 1967). Tidal currents transport coarser sediment by traction across
the substrate, while fi ne sediment is held in suspension. During slack-water, transport
through traction ceases and fi ne sediment held in suspension is deposited as drapes
45
(Reineck, 1967). The segregation of limestone and siltstone within this facies is likely
due to silt-sized particles being carried into the back-barrier environment in suspension
with the fl ood tide, getting partially deposited during slack-water, and being carried out in
suspension with the ebb tide.
Facies I is comprised of cherty limestones with micritic pebbles forming an
oligomict conglomerate. Chert layers, lenses, and veins dissect the carbonate matrix
and typically parallel bedding, although sometimes do cut across it in parts. Limestone
and chert occur in approximately equal abundance. Micritic pebbles range in size from
several millimeter to several centimeters in their longest dimension. Pebbles vary in
degree from angular to rounded, and are generally poorly sorted. Most pebbles occur
within a chert matrix. Irregular syndepositional folding occurs throughout this facies,
and is most extensively developed in association with large irregularly shaped micrite
or chert-fi lled cavities. This facies is found within the Upper Member of the Poleta
Formation where it disconformably overlies quasi-planar laminated sandstones of facies
C. Facies I represents a solution collapse breccia, likely formed within a supratidal fl at
environment. Solution collapse breccias form due to post-depositional dissolution of
evaporite minerals formed primarily within evaporative supratidal settings. Dissolution
is accompanied by collapse, which can promote syndepositional folding (Lock and
Roberts, 1999), and brecciation of the host sediments (Middleton, 1961; Clifton, 1967).
The abundance of chert within this facies may be due to its replacement of former
mineralogies, as is the case in many Cambrian and Precambrian evaporite facies in which
silicifi cation serve as common replacement process (e.g., Tucker, 1976; Young, 1979;
Nicolaides, 1995; Maliva et al., 2005). The presence of chert veins dissecting this facies
is likely the result of fracturing during brecciation (Clifton, 1967).
Facies J consists of limestone with mm-thick chert seams that are arranged into
a cm-scale netted pattern. This pattern is reminiscent of the chicken-wire structures
46
produced during the crystallization of evaporative mosaic gypsum or anhydrite that
displaces the host sediment (e.g., Riley and Byrne, 1961; Dean et al., 1975). Carbonate
pseudomorphs after gypsum or anhydrite are implied based on the morphology of the
fabrics. This replacement process can occur during early stage diagenesis soon after
deposition, or can occur later during deep burial diagenetic processes (e.g., Pierre and
Rouchy, 1988). The presence of chert seams, rather than clay or silt seams, may be
due to the presence of evaporative silica within Precambrian-Cambrian evaporative
environments that was responsible for extensive silicifi cation of marginal marine
deposits (e.g., Tucker, 1976; Young, 1979; Nicolaides, 1995; Maliva et al., 2005). Facies
J is interpreted to represent deposition within a supratidal fl at environment in which
evaporative processes dominated. The close association of this facies with solution
collapse breccias of facies I reinforces the supratidal and evaporative nature of this facies.

Tidal Inlet Successions and Evidence for a Barrier Island Coastline - Upper Member
Wood Canyon Formation
The upper member of the Wood Canyon Formation at Emigrant Pass, Death
Valley region, is comprised of facies B, C, D, E, and F. The base of the member is not
exposed at this locality and the measured section begins within siliciclastic facies below
the prominent carbonate unit of the member, passes though the siliciclastic-carbonate
transition, and terminates within the carbonate unit of the member (Fig. 2.4). Two
types of conformable successions occur within the upper member. One consists of the
upward-coarsening succession of interbedded siltstone and very fi ne sandstone of facies
B, amalgamated hummocky cross-stratifi ed fi ne sandstone of facies C, and small-scale
trough cross-stratifi ed medium sandstone of facies D (Fig. 2.5A-B). This succession
represents progradation of the shoreline from offshore transition to lower shoreface, and
capped by ebb-tidal delta or upper shoreface deposits.
47
Lower
shoreface
Ebb-tidal
delta
Ebb-tidal
delta
Ebb-tidal
delta
Lower
shoreface
Offshore
transition
VC
C
M
F
VF
Silt
Shale
Lithology
Lower
shoreface
FS
FS
FS
Ebb-tidal
delta
FS
FS
Lagoon
Flood-tidal
delta
Flood-tidal
delta
Flood-tidal delta
Ebb-tidal
delta
Flood-tidal
delta
Siltstone and hummocky
cross-stratified sandstone
Hummocky cross-stratified
sandstone
Trough cross-stratified sandstone
+/- herringbone cross-stratification
Planar laminated sandstone
Small-scale trough cross-
stratified sandy dolostone
Medium-scale trough cross-
stratified oolitic dolostone
Dolostone with sandy
layers and vertical fabric
0
2
4
6
8
10
12
14
16
18
20
22
24
26
28
30
1
3
5
7
9
11
13
15
17
19
21
23
25
27
29
31
meters
Erosional unconformity
upper member Wood Canyon
FIGURE 2.4. Partial stratigraphic section measured through the upper member of the
Wood Canyon Formation at Emigrant Pass, southern Nopah Range, California and
interpreted depositional settings. Base of section begins at N 35° 53.577 ′, W 116°
04.800′. FS = fl ooding surface.
48
FIGURE 2.5. Sedimentary facies from the upper member of the Wood Canyon
Formation, Emigrant Pass locality. (A) Outcrop photograph of hummocky cross-stratifi ed
fi ne sandstone; lower shoreface. Staff scale is in decimeters. (B) Outcrop photograph of
bidirectional small-scale trough cross-stratifi ed and herringbone cross-stratifi ed medium
sandstone; ebb-tidal delta. Pen is 15 cm long. (C) Outcrop photograph of medium-scale
trough cross-stratifi ed oolite with reactivation surfaces; fl ood tidal delta. Staff scale is
in decimeters. (D) Outcrop photograph of medium-scale trough cross-stratifi ed oolite
deposited within channel incised into underlying bidirectional cross-stratifi ed ebb-tidal
delta deposits. Overlying the oolite is bioturbated sandy dolostone; lagoon.
49
The second conformable succession consists of medium-scale trough and
tabular cross-stratifi ed oolitic dolostones with reactivation surfaces of facies E
overlain by massively bedded and heavily bioturbated sandy micritic dolostones with
rare echinoderm fragments of facies F (Fig. 2.5C). Ooids within the cross-stratifi ed
dolostones can range up to coarse sand in size. These two facies together are interpreted
as fl ood-tidal delta and lagoonal environments, respectively. The contact between ebb-
tidal delta deposits of facies D and fl ood-tidal delta deposits of facies E, marking the
transition between the two conformable successions, is generally sharp and planar. This
contact may be conformable in parts, although carbonate facies E and F do also occur
within a broad channel incised into ebb-tidal delta deposits (Fig. 2.5D). This erosional
unconformity likely represents the throat of a tidal inlet.
A complete upward-shallowing succession consists of upward-coarsening
siliciclastic deposits of the offshore transition and shoreface overlain successively by
fl ood-tidal and lagoonal deposits of a carbonate back-barrier setting. Upward-shallowing
successions at this locality show a progradational stacking pattern, with each succession
representing a shallower transect than the previous.
Tidal Inlet Successions and Evidence for a Barrier Island Coastline - Hines Tongue Reed
Dolomite
The Hines Tongue of the Reed Dolomite at Hines Ridge is comprised of facies
C, D, E, and F. The measured section encompasses the lowermost 34 meters exposed
near the base of the member (Fig. 2.6). The facies successions preserved within the
Hines Tongue are generally similar to those within the upper member of the Wood
Canyon Formation. Amalgamated quasi-planar laminated fi ne sandstones of facies C are
consistently and conformably overlain by small-scale trough cross-stratifi ed and planar
laminated medium sandstones of facies D, representing deposition within the lower
50
Flood-tidal
delta
Trough cross-stratified
sandstone
Quasi-planar laminated
sandstone
Lower
shoreface
1
0
2
4
3
5
7
6
8
10
9
11
VC
C
M
F
VF
Silt
Shale
Lithology
FS
12
14
13
15
16
18
17
19
20
21
23
22
Flood-tidal
delta
Lower shoreface
Ebb-tidal delta
FS
Tidal
delta
25
24
26
27
Medium-scale trough
cross-stratified dolostone
29
28
30
31
32
33
Lower
shoreface
Ebb-tidal
delta
Ebb-tidal
delta
FS
Planar laminated dolostone
Planar laminated sandstone
meters
Hines Tongue Reed Dolomite
Erosional unconformity
FIGURE 2.6. Partial stratigraphic section measured near the base of the Hines Tongue
of the Reed Dolomite at Hines Ridge in the White-Inyo Mountains and interpreted
depositional settings. Base of section begins at N 37° 06.317 ′, W 118° 05.446 ′. FS =
fl ooding surface.
51
10 cm
FIGURE 2.7. Sedimentary facies from the Hines Tongue of the Reed Dolomite, Hines
Ridge locality. (A) Outcrop photograph of bidirectional small-scale trough cross-stratifi ed
medium sandstone; ebb-tidal delta. Hammer is 30 cm long. (B) Outcrop photograph of
medium-scale trough and tabular cross-stratifi ed dolostone with reactivation surfaces;
fl ood tidal delta. Staff scale is in decimeters. (C) Outcrop photograph of the sharp and
erosive contact between underlying planar laminated sandy ebb-tidal delta deposits and
overlying tabular cross-stratifi ed dolostone fl ood-tidal delta deposits. Hammer is 30 cm
long.
52
shoreface and ebb-tidal delta, respectively (Fig. 2.7A). Ebb-tidal delta sandstones are
sharply overlain by planar laminated and medium-scale cross-stratifi ed dolostones with
reactivation surfaces of facies E, interpreted as fl ood-tidal delta deposits (Fig. 2.7B).
Whether the contacts between these facies are conformable or not is diffi cult to
determine as most contacts are planar; however, in parts erosional scours may be present,
with fl ood-tidal delta deposits of facies E shallowly incising ebb-tidal deposits of facies
D (Fig. 2.7C). Similar to the facies succession found in the upper member of the Wood
Canyon, the upward-shallowing facies succession of the Hine Tongue represents the
progradation of a barrier island coastline and a shift from shoreface siliciclastics to back-
barrier carbonates by way of a tidal inlet.
Barrier Island Successions without Tidal Inlet Deposits - Lower Member Wood Canyon
Formation
The lower member of the Wood Canyon Formation at Chicago Pass in the
northern Nopah Range consists of mixed siliciclastic-carbonate sedimentary rocks
with three prominent carbonate subunits. The member is arranged into three upward-
shallowing successions that are fl oored by shallow shelf siliciclastic sedimentary rocks
that are capped by shallow marine carbonate sedimentary rocks (Corsetti and Hagadorn,
2000). The Precambrian-Cambrian boundary falls within siliciclastic facies found
between the second and third carbonate units (Corsetti and Hagadorn, 2000) (Fig.
2.3). The measured section begins 5.5 meters below the base of the fi rst carbonate unit,
includes the siliciclastic-carbonate transition, and terminates within the carbonate unit
(Fig. 2.8). All rocks examined are Neoproterozoic in age. The lower member of the
Wood Canyon Formation includes facies C, D, and F.
Siliciclastic facies below the carbonate unit consists of swaley cross-stratifi ed
fi ne sandstone, trough cross-stratifi ed medium sandstone, and parallel laminated medium
53
Upper
shoreface
VC
C
M
F
VF
Silt
Shale
Lithology
Lower
shoreface
Lagoon
Swaley cross-stratified
sandstone
Trough cross-stratified sandstone
+/- herringbone cross-stratification
Planar laminated sandstone
Dolostone with massive
bedding
0
2
4
6
8
10
1
3
5
7
9
meters
Erosional unconformity
lower member Wood Canyon
FIGURE 2.8. Partial stratigraphic section measured through the siliciclastic-carbonate
transition of the fi rst carbonate-capped unit of the lower member of the Wood Canyon
Formation at the Chicago Pass locality and interpreted depositional settings. Base of
section begins at N 36° 08.565 ′, W 116° 09.175 ′.
54
A
B
C
10 cm
FIGURE 2.9. Sedimentary facies from the lower member of the Wood Canyon
Formation, Chicago Pass locality. (A) Outcrop photographs of swaley cross-stratifi ed fi ne
to very fi ne sandstone with large prominent swale in upper left corner. Staff scale is in
decimeters. (B) Outcrop photograph of small-scale herringbone cross-stratifi cation; upper
shoreface. (C) Outcrop photograph of poorly bedded to massively bedded tan dolostone
that makes up the prominent carbonate units within the lower member of the Wood
Canyon Formation. Staff scale is in decimeters.
55
sandstone. These facies are arranged into a single upward-shallowing succession that
leads into the base of the carbonate unit. The basal portion of the siliciclastic succession
consists of swaley cross-stratifi ed fi ne sandstone of facies C (Fig. 2.9A). Swaley cross-
stratifi cation represents the more nearshore equivalent to hummocky cross-stratifi cation
in which aggradation rates are lowered and hummocks are preferentially eroded away
while swales are preferentially preserved (Dumas and Arnott, 2006). Above the swaley
cross-stratifi ed sandstone is preserved a unit of small-scale trough cross-stratifi ed medium
sandstone with herringbone cross-stratifi cation of facies D that is interbedded with
planar laminated medium sandstone (Fig. 2.9B). Trough cross-stratifi ed sandstones are
interpreted as an upper shoreface environment in which the migration of subaqueous
dunes predominates (e.g., Clifton et al., 1971). Interbedded planar laminated sandstones
may represent deposition within the foreshore, due to swash processes (e.g., Clifton,
1969), or may represent planar bedforms that develop and become interspersed with
subaqueous dunes in the outer surf zone (Clifton et al., 1971).
Carbonate units within the lower member of the Wood Canyon are bedded at
the decimeter scale or massively bedded. Lowermost carbonates of the fi rst carbonate
unit consist of massively bedded dolostone of facies F with no prominent sedimentary
structures (Fig. 2.9C). The lack of sedimentary structures within this facies suggests
a quiet-water depositional environment, while the superposition of this facies above
nearshore siliciclastics suggests a lagoonal environment, similar to that present within the
carbonate facies upper member of the Wood Canyon Formation. Unlike the carbonate
facies of the upper member, the carbonate unit of the lower member does not contain
any evidence for tidal activity or for the presence of tidal inlet deposition, and therefore
is likely emplaced atop nearshore siliciclastic facies across an unconformity with barrier
island deposits missing. The precise contact between underlying siliciclastics and
56
Tidal
channel
Trough cross-stratified sandy
packstone and grainstone
Hummocky cross-
stratified sandstone
Siltstone and hummocky
cross-stratified sandstone
Tidal
channel
levee
1
0
2
4
3
5
7
6
8
10
9
11
VC
C
M
F
VF
Silt
Shale
Lithology
FS
12
14
13
15
16
18
17
19
20
21
23
22
Tidal
channel
Tidal
channel
levee
Offshore
transition
Lower
shoreface
FS
Tidal
channel
25
24
26
27
Amalgamated graded
intraclastic limestone
meters
Lower Member Deep Spring Formation
Erosional unconformity
FIGURE 2.10. Partial stratigraphic section measured through the basal siliciclastics and
carbonates of the Lower Member of the Deep Spring Formation at Hines Ridge in the
White-Inyo Mountains and interpreted depositional settings. Base of section begins at N
37° 06.173 ′, W 118° 05.628 ′. FS = fl ooding surface.
57
10 cm
AB
C D
FIGURE 2.11. Sedimentary facies from the Lower Member of the Deep Spring
Formation, Hines Ridge locality. (A) Outcrop photograph of massive siltstone with a
single prominent hummocky cross-stratifi ed very fi ne sandstone bed (center); offshore
transition. Staff scale is in decimeters. (B) Outcrop photographs of hummocky cross-
stratifi ed fi ne to very fi ne sandstone with prominent hummock in top of photo; lower
shoreface. (C) Outcrop photograph of small-scale trough cross-stratifi ed sandy
grainstone; tidal channel fi ll. Scale bar = 4 cm. (D) Outcrop photograph of amalgamated
graded intraclastic limestone beds with erosive bases; tidal channel levee. Scale bar = 4
cm.
58
overlying carbonates is not perfectly exposed at this locality, so it is diffi cult to observe
whether any discernible erosion is present across the siliciclastic-carbonate transition.
Barrier Island Successions without Tidal Inlet Deposits - Lower Member Deep Spring
Formation
The Lower Member of the Deep Spring Formation at Hines Ridge is comprised
of facies B, C, E, and G. The measured section records the lowermost 27 meters
exposed, which starts at approximately the base of the member and includes siltstone,
quartz sandstone, cross-stratifi ed peloidal limestones, and intraclastic limestones (Fig.
2.10) (e.g., Gevirtzman and Mount, 1986; Parsons, 1996). The lowermost 4.3 meters
consists of interbedded siltstone and hummocky cross-stratifi ed sandstones of facies B
(Fig. 2.11A). Sandstone beds thicken upward from the base from 5-10 cm to 10-20 cm
near the top of the facies. The next 3.5 meters is comprised of amalgamated hummocky
cross-stratifi ed sandstones of facies C, with the lowermost contact gradational with
the underlying facies B (Fig. 2.11B). This two facies together represent a prograding
siliciclastic shoreline with facies B and C representing deposition within the offshore
transition and lower shoreface, respectively.
Facies B is sharply overlain by facies E consisting of medium-scale trough cross-
stratifi ed peloidal grainstone and facies G comprised of graded intraclastic limestones
(Fig. 2.11C-D). The couplet of facies E overlain by facies G is interpreted as a migrating
back-barrier tidal channel-levee association that forms an upward-shoaling succession.
The sharp contact between lower shoreface sandstones and back-barrier cross-stratifi ed
grainstones suggests a disconformable contact between these two facies, similar to those
that form within prograding barrier islands in which tidal channel incision and dissection
dominate the back-barrier environment (Duc and Tye, 1987). This association repeats
59
again upsection, shoaling from tidal channel deposits of facies E to tidal channel levee
deposits of facies G.
Barrier Island Successions without Tidal Inlet Deposits - Poleta Formation
The Poleta Formation is exceptionally exposed within a canyon at the
northeastern extent of the Poleta folds region within the White-Inyo Mountains. At
this locality the Middle Member and lowermost Upper Member are well exposed. This
canyon has been studied previously, most notably preserving the large early Cambrian
trace fossil Taphrhelminthopsis, which occurs across extensively exposed bedding planes
approximately 33 and 35 meters above the base of the Middle Member (Hagadorn et
al., 2000). The measured section records the Middle Member and lowermost Upper
Member and is comprised of facies A, B, C, H, I, and J (Fig. 2.12). A complete upward-
shallowing succession, as recognized within siliciclastics of the uppermost Middle
Member and carbonates of the lowermost Upper Member, consists of a basal unit of
interbedded siltstone and planar laminated thin sandstone beds of facies A, which is
overlain by a upward-thickening succession of hummocky cross-stratifi ed and quasi-
planar laminated sandstone beds of facies B that become amalgamated upsection within
facies C (Fig. 2.13A). Hummocky cross-stratifi ed sandstone of facies C is sharply
overlain by fl aser bedded silty carbonates in lower parasequences of the Middle Member
(Fig. 2.13B-C), and at approximately 28.5 meters above the base this facies contact
is notably erosional (Fig. 2.13C). At the Middle-Upper Member contact, uppermost
hummocky cross-stratifi ed sandstones of the Middle Member are capped by an oligomict
carbonate pebble conglomerate of facies I within the lowermost Upper Member (Fig.
13D). This conglomerate grades upwards into cherty limestone with synsedimentary
folding and quartz-fi lled cavities (Fig. 13D). This is capped by limestone with cherty
chicken-wire structures of facies J (Fig. 13E).
60
Siltstone and hummocky
cross-stratified sandstone
Bedded to laminated siltstone
with thin sandstone beds
0
2
6
4
10
8
12
16
14
18
20
24
22
26
28
32
30
34
36
38
42
40
44
46
Lower
shoreface
Offshore
transition
Proximal
offshore
Offshore transition
Lower shoreface
Tidal flat
Lower shoreface
Tidal flat
Proximal
offshore
Offshore
transition
Lower shoreface
Offshore transition
Lower shoreface
Tidal flat
Proximal
offshore
Offshore
transition
Lower
shoreface
Supratidal flat
FS
FS
FS
FS
Micritic limestone
Hummocky cross-stratified
sandstone
Solution collapse
breccia
FS
Middle Member Poleta
Upper
Member
FS
meters
Erosional unconformity
Chert laminae
Chert layers with micritic pebbles
Silt flasers
Micrite or quartz-filled cavities
Chicken wire structures
VC
C
M
F
VF
Silt
Shale
Lithology
Taphrhelminthopsis
FIGURE 2.12. Stratigraphic section measured through the complete Middle Member and
lowermost Upper Member of the Poleta Formation at the Poleta Folds in the White-Inyo
Mountains and interpreted depositional settings. Base of section begins at N 37° 18.877 ′,
W 118° 05.444 ′. FS = fl ooding surface.
61
4 cm
3 cm
surface of erosion
anticlinal
folding
FIGURE 2.13. Sedimentary facies from the Middle Member and Upper Member of the
Poleta Formation, Poleta Folds locality. (A) Outcrop photograph of hummocky cross-
stratifi ed fi ne to very fi ne sandstone; lower shoreface. This sandstone is overlain by fl aser
bedded limestone. Staff scale is in decimeters. (B) Close-up view of the erosive nature of
the contact between underlying hummocky cross-stratifi ed sandstone and overlying fl aser
bedded limestone shown in Fig. 13A. Hammer is 24 cm long. (C) Close-up view of fl aser
bedded limestone and siltstone. (D) Outcrop photograph of oligomict cherty limestone
with syndepositional anticlinal folding (upper left) and cavities fi lled with carbonate
(upper left). Hammer is 24 cm long. E) Outcrop photograph of limestone with cherty
chicken-wire structures.
62
The siliciclastic portion of the parasequences of the Middle Member of the Poleta
Formation is interpreted as an upward-shallowing siliciclastic storm-dominated shelf
and shoreface. Facies A and B record increasingly more proximal sandy storm deposits,
of the proximal offshore and offshore transition, respectively, while facies C records the
amalgamation of these storm deposits within the lower shoreface, above fair-weather
wave base. The observation that these lower shoreface sandstones are sharply overlain by
carbonate tidal fl at deposits of facies H throughout much of the Middle Member suggests
that this contact represents an unconformity and that the sharp lithologic contrast may be
due to the superposition of non-adjacent depositional environments. Further evidence
for this intra-parasequence unconformity is the observation that across the Middle-Upper
Member boundary there is a dramatic shift from lower shoreface deposits to a solution
collapse breccia overlain by carbonate with chicken-wire structures of a supratidal
fl at environment. This superposition implies a signifi cant hiatus between fore-barrier
shoreface deposits and back-barrier supratidal fl at deposits.
Barrier Island Depositional and Erosional Processes
It is generally regarded that during storm surge, most barrier islands experience
erosion on the shoreface, foreshore, and backshore, and deposition on back-barrier fl ats
and in lagoons (e.g., Morton, 2002). This process of storm overwash, in which sand is
transported from the seaward side of the barrier island to the back-barrier environment,
is the primary means by which barrier islands migrate landward over time. This type
of storm overwash is a well-documented process along modern barrier island coastlines
(Kochel and Dolan, 1986). There are, however, processes that can work to erode the
back-barrier environment and may allow for the gradual progradation of the lagoonal
environment in a seaward direction. If coupled with progradation of the shoreface, as is
the case on modern prograding barriers islands such as Galveston Island (e.g., Bernard
63
et al., 1962), this would allow for the barrier island to translate seaward over time as a
discrete entity, without expanding signifi cantly in barrier width.
The primary means by which erosion occurs in back-barrier environments under
fair-weather conditions is through tidal channel incision (e.g., Duc and Tye, 1987;
Pilkey et al., 1989). Tidal channel development is a destructive agent that opposes
the constructional processes of deposition within fl ood-tidal deltas and during storm
overwash (Pilkey et al., 1989). Erosion occurs primarily with each receding tide and
is most extensively developed during large spring tides (Pilkey et al., 1989). These
channels are capable of incising into back-barrier sand fl ats and vegetated dunes, cutting
down to shoreface deposits, and can erode laterally with migration (Duc and Tye, 1987;
Pilkey et al., 1989). On Kiawah Island, South Carolina, back-barrier tidal channel
incision is known to have incised and removed 4.3 meters of underlying foreshore and
shoreface deposits, in addition to overlying storm washover deposits (Duc and Tye,
1987).
Under storm conditions, washover deposits are frequently emplaced within the
back-barrier environment, although erosional processes also operate. In most instances,
erosion in the back-barrier environment during storms is due to the development of large
waves accommodated by a long back-barrier fetch (e.g., Armbruster et al., 1995; Stone
et al., 2004). These waves are most pronounced when the winds parallel the longest axis
of the back-barrier environment (Jackson, 1995). What this implies is that erosion due
to storm waves is most pronounced in large lagoons in which there is suffi cient surface
area for the wind to act upon and for waves to develop. Winter cold fronts passing Santa
Rosa Island, Florida, are capable of removing up to 1.75 meters from the back-barrier
foreshore with a single storm passage (Armbruster et al., 1995). In rare instances when
storm surge is higher in bays and lagoons than in the ocean, storm washout can occur.
Storm washout is opposite that of storm washover and occurs when water and sediment
64
are transported seaward across the barrier island from the back-barrier lagoon into the
ocean (Morton, 2002). This typically leads to barrier island incision, similar to that
generated by storm overwash.
Progradation and seaward building of a barrier island can also lead to erosion
within the back-barrier environment. As a barrier island grows seaward it gets wider
if there are no erosional processes acting on the back-barrier side of the island. When
barrier islands are narrow, storm overwash is capable of transporting sediment to the
back-barrier lagoon where it can be preserved; however, when barrier islands are wide,
storm overwash is less likely to make it to the back-barrier lagoon and is commonly
deposited on back-barrier fl ats (Maynard and Suter, 1983). This is because frictional
dissipation of energy signifi cantly inhibits sediment transport across the barrier, except
in instances in which the barrier is narrow or the surging water level is high (Boyd and
Penland, 1981). Therefore, as a barrier island grows wider with progradation, there
will be a growing starvation of sediment within the back-barrier environment as less
sediment is capable of reaching the back-barrier lagoon. Also, storm overwash deposited
in subaerial environments, such as back barrier fl ats and washover fans, is susceptible
to eolian defl ation and much of the overwash-derived sediment can be redistributed
into backshore dunes and the seaward-facing beach, reducing net accumulation
(Leatherman and Zaremba, 1987). For the back-barrier environment of Galveston
Island, Texas—which is a prograding barrier island—it has been shown that there is a
strong correspondence between back-barrier erosion and sediment starvation, as relative
sea-level rise outpaces sediment accretion rate. Net erosion due to sediment starvation
within back-barrier environments may be a disequilibrium response where post-storm
recovery from barrier island erosion is inhibited due to low sediment supply, and the
shoreline simply fails to rebuild itself (e.g., Xue et al., 2009). The promotion of sediment
starvation and back-barrier erosion with increasing barrier width may actually keep most
65
barrier islands below some critical width above which sediment supply would not be
suffi cient to sustain the barrier and maintain equilibrium.
While the process of storm overwash is widely known to translate barrier islands
landward as discrete entities, there is ample evidence to suggest that similar erosional
and depositional processes may work in concert to translate prograding barrier islands
seaward as discrete entities. Erosion of the back-barrier environment by tidal channel
incision and by storm waves are the primary means by which back-barrier sediments
are eroded, while the act of progradation can lead to increased barrier width, sediment
starvation, and the initiation of back-barrier erosion as well. Implicit in this scenario is
that a prograding barrier island, if coupled with back-barrier erosion, will simply leave
behind a minimal record of the actual subaerial barrier, as it is likely destroyed during
progradation. This has wide-ranging implications for the depositional record left behind
by prograding barrier islands and predicts that with no preserved barrier, foreshore
and shoreface deposits should be directly overlain by back-barrier deposits across an
erosional unconformity, and this appears to be the case for the barrier island systems of
the Precambrian-Cambrian transition of the southern Great Basin.
The Self-Destructive Prograding Barrier Island Model
Prograding barrier island deposits are common in the rock record, yet few modern
coastlines contain deposits that produce analogous sedimentary successions. Ancient
prograding barrier island deposits can occur in both siliciclastic and carbonate settings.
In prograding siliciclastic barrier island deposits of the Upper Eocene of southern
England, lower and upper shoreface deposits are incised by washover deposits, which
are overlain by back-barrier tidal fl at, marsh, and lagoonal deposits (Plint, 1984). The
Cretaceous Frontier Formation of Utah preserves prograding siliciclastic offshore to
foreshore successions that are abruptly capped by siltstones and shales with a brackish
66
BACKSHORE-FORESHORE
TIDAL FLAT
PROGRADATION
Erosional unconformity
Time lines
Conformable contact
Siltstone and hummocky
cross-stratified sandstone
Hummocky cross-stratified
sandstone
Trough cross-stratified
sandstone
Planar laminated and large-scale
cross-stratified sandstone
Massive mudstone
Flaser bedded sanstone
and siltstone
LAGOON
EROSION
UPPER SHOREFACE
PROGRADATION
FWWB
THE SELF-DESTRUCTIVE PROGRADING BARRIER ISLAND MODEL
OFFSHORE
TRANSITION
LOWER
SHOREFACE
UPPER
SHOREFACE
LAGOON
TIDAL FLAT
LOWER SHOREFACE
OFFSHORE TRANSITION
FIGURE 2.14. The self-destructive barrier island model based upon strata from the
southern Great Basin. In this model, progradation of the barrier island coastline is
accompanied by erosion of the backside of the barrier island. These two processes result
in a barrier island that translates seaward over time without growing in width. Based on
these processes, this model predicts that back-barrier deposits should overlie fore-barrier
deposits across an erosional unconformity, with the barrier island eolian backshore
deposits missing. Modifi ed from McCubbin (1982).
67
fauna of a back-barrier lagoon (Ryer, 1977). Strata from the Upper Cretaceous Cardium
Formation of northwestern Alberta, Canada record similar facies of upward-coarsening
shoreface and foreshore deposits abruptly overlain by oyster-rich lagoonal mudstones,
with uppermost foreshore deposits being commonly rooted (Plint and Walker, 1987). In
carbonate facies of the Lower Cretaceous Edwards Formation, tabular cross-stratifi ed
grainstones of the shoreface are overlain by low-angle and planar laminated grainstones
of the foreshore that are capped by a palaeocaliche crust, suggesting prolonged exposure
(Kerr, 1977). These foreshore grainstones are abruptly succeeded by heavily bioturbated
wackestones of a lagoonal environment.
These examples from the rock record are very similar to those recorded from
the Precambrian-Cambrian transition of the southern Great Basin. In all instances, even
though facies associated with barrier islands are present and distinctly fore-barrier and
back-barrier, there is no preservation of barrier island deposits, which should consist
of backshore dunes with large-scale and high-angle cross-stratifi cation. In all these
successions, most upward-shoaling deposits reach no environment shallower than the
foreshore before there is a dramatic and sharp facies shift to back-barrier deposits with no
intervening barrier island dune deposits.
Herein a model is proposed to explain this absence of barrier island facies from
prograding barrier island deposits in the rock record. This model is termed the ‘self-
destructive prograding barrier island model’. Given that erosional processes are indeed
known to operate in back-barrier environments (e.g., Duc and Tye, 1987; Armbruster
et al., 1995), and that barrier islands in the modern ocean are capable of shoreline
progradation (e.g., Bernard et al., 1962), it is highly likely that these two processes can
operate in concert to generate barrier islands that are translational in nature rather than
progradational. In this model, shoreline progradation leads to seaward advancement of
the barrier island, while back-barrier erosion maintains a relatively conservative width of
68
the barrier. This model therefore implies that the translation of the barrier seaward leads
to the accumulation of new sediment on the shoreface and shelf with a contemporaneous
destruction of older barrier island sediments on the back-barrier side.
An analogous process can be found on modern barred coastlines. On barred
coastlines, a linear crested bar of sediment within the shoreface is separated from the
coastline by a longshore trough. This trough is host to longshore currents that scour the
landward side of the bar and can lead to its erosion (Hunter et al., 1979). The predicted
vertical arrangement of facies for a barred coastline is an upward-coarsening and
shallowing succession in which facies on the seaward side of the bar are abruptly overlain
by trough facies across an erosional surface (Hunter et al., 1979). This is because as the
longshore trough and bar translate seaward, the trough erodes the landward side of the
bar, continually removing it from the record as soon as it forms (Hunter et al., 1979).
Therefore, in this situation in which progradation is coupled with erosion, the bar itself
takes on a translational nature, which does not inhibit its perpetuation, but does largely
prohibit its preservation.
In the case of prograding barrier islands, it is therefore expected that if erosion
is suffi cient, translation of the barrier is the result and the barrier facies are removed
from the record soon after they are formed (Fig. 2.14). This model also predicts that
the contact between underlying shoreface and foreshore deposits and overlying back-
barrier deposits represents an erosional surface and an unconformity. In the southern
Great Basin the shift from fore-barrier to back-barrier is marked by a shift in lithology
from siliciclastic to carbonate sedimentary rocks, and in many instances these contacts
are regionally extensive from the Death Valley region to the White-Inyo Mountains.
This would require the progradation of a barrier island system between these regions to
produce this type of extensive contact.
69
In some instances, ancient barrier island systems are known to have prograded
long distances. A barrier island system within the Upper Cretaceous Cardium Formation
of Alberta, Canada exhibits 100 km of seaward progradation (Plint and Walker, 1987),
while barrier island deposits of the Lower Cretaceous Notikewen Member of the Gates
Formation of British Columbia, Canada show 160 km of progradation (Leckie, 1985).
These distances are on the order of progradation required to traverse the Precambrian-
Cambrian distance between the southern Nopah Range and the White-Inyo Range, which,
based on palinspastic reconstructions, were approximately 150 km apart (Levy and
Christie-Blick, 1989; 1991).
It might be expected that barrier island dune deposits should be preserved at the
furthest extent of progradation before subsequent transgression, but this is likely not
the case. Transgression would shift the prograding barrier island into a transgressive
barrier island. Transgressive barrier island models predict the retrograde stacking of
lagoonal deposits, overlain by marsh and washover deposits, and capped by backshore
dune deposits (e.g., Hoyt, 1967; Kraft and John, 1979; Reinson, 1979). This stacking
pattern occurs due to landward migration of the barrier through the combined effects of
storm washover—which supplies sand to the back-barrier environment—and shoreface
retreat (Kraft and John, 1979; Kochel and Dolan, 1986). Ultimately, however, continued
transgression of the shoreline leads to reworking of some or all of the previously
deposited barrier and back-barrier sediments and the development of a transgressive
surface of erosion during shoreface retreat (e.g., Fischer, 1961). This may explain why
prograding barrier islands are so rarely preserved. Back-barrier erosion removes barrier
island deposits during progradation and subsequent transgression removes what remains
of the barrier island at the termination of its progradation.
The types of siliciclastic-carbonate transitions found within Precambrian-
Cambrian strata of the southern Great Basin may refl ect the position along the barrier
70
island across which progradation occurred and where deposition ceased (Fig. 2.15).
Successions in which ebb-tidal and fl ood-tidal delta deposits are found, including the
upper member of the Wood Canyon Formation and Hines Tongue of the Reed Dolomite,
likely passed through a tidal inlet that dissected the barrier island. Successions in which
tidal inlet deposits are not preserved likely passed though and under the barrier island.
In these cases, the type of upward-shallowing succession and siliciclastic-carbonate
transition is refl ected in what environmental setting deposition ceased before the next
upward-shallowing succession ensued. Strata of the lower member of the Wood Canyon
Formation likely formed through translation of a narrow barrier from the shoreface to a
back-barrier lagoon. Strata of the Lower Member of the Deep Spring Formation were
the result of down-cutting of tidal channels into shoreface deposits following barrier
island progradation. The Middle and Upper Poleta Formation were likely the result of
the development of an evaporative tidal fl at setting following exhumation of shoreface
deposits following barrier progradation. This might be expected for a barrier island
setting with extensive back-barrier fl ats, but minimal lagoonal development.
Conclusions
Prograding barrier island systems are common in the rock record, but rare in the
modern ocean. This is because most modern barrier island systems are transgressive in
nature. When barrier island system deposits are found in the rock record they preserve
shelf to shoreface successions that are abruptly overlain by back-barrier lagoonal and
tidal fl at deposits without preserving the intervening barrier. Ideally this barrier would be
represented by large-scale and high-angle cross-stratifi ed sandstones of backshore dunes;
however, more typically shoreface or foreshore deposits are directly overlain by back-
barrier deposits.
71
Lagoon
Tidal
channel
Intertidal
flat
Supratidal
flat
Ebb-tidal
delta
Ebb-tidal
delta
Lagoon
Flood-tidal
delta
Lower
shoreface
Offshore
transition
Upper
shoreface
Lagoon
Lower
shoreface
Offshore
transition
Tidal
channel
levee
Tidal
channel
Lower
shoreface
Offshore
transition
Supratidal
flat
Intertidal flat
Lower
shoreface
Offshore
transition
1 23 4
Solution
collapse
breccia
upper member
Wood Canyon
lower member
Wood Canyon
Lower Member
Deep Spring
Middle-Upper
Poleta
Trough cross-stratified sandy
packstone and grainstone
Hummocky cross-
stratified sandstone
Siltstone and hummocky
cross-stratified sandstone
Amalgamated graded
intraclastic limestone
Erosional unconformity
Micritic limestone
Chert laminae
Chert layers with micritic pebbles
Silt flasers
Micrite or quartz-filled cavities
Chicken wire structures
Medium-scale trough cross-
stratified oolitic dolostone
Dolostone with massive bedding
Trough cross-stratified sandstone
+/- herringbone cross-stratification
Tidal
inlet
Upper shoreface
3
2
1
4
Barrier island
Flood-tidal
delta
FIGURE 2.15. Predicted locations for Neoproterozoic-Cambrian barrier island deposits
within a generalized barrier island coastline. 1) the upper member Wood Canyon
succession results from progradation through a tidal inlet, terminating in a lagoon; 2) the
lower member Wood Canyon succession results from progradation of a narrow barrier
island, terminating in a lagoon; 3) the Lower Member Deep Spring succession results
from barrier island progradation coupled with back-barrier tidal channel incision into
shoreface deposits; 4) the Middle-Upper Poleta succession results from progradation of a
wide barrier island with extensively developed back-barrier intertidal and supratidal fl ats.
72
Precambrian-Cambrian strata of the southern Great Basin preserve mixed
siliciclastic-carbonate prograding barrier island systems. In these systems, siliciclastic
sedimentary rocks were deposited in shallow shelf and shoreface environments, while
carbonate rocks formed in back-barrier lagoons and tidal fl ats. While the Hines Tongue
of the Reed Dolomite and the upper member of the Wood Canyon Formation preserve
evidence of the shift from the fore-barrier environment to the back-barrier environment in
the form of tidal inlet successions that include ebb-tidal delta and fl ood-tidal deltas, there
are sedimentary successions that do not.
The lower member of the Wood Canyon Formation preserves lower shoreface and
upper shoreface sandstone deposits that are overlain by massively bedded dolostones of
a lagoonal environment. The Lower Member of the Deep Spring Formation preserves
hummocky cross-stratifi ed lower shoreface sandstones that are abruptly overlain by
trough cross-stratifi ed limestones and graded intraclastic limestones that represent
carbonate tidal channel and levee deposits, respectively. The Middle Member and Upper
Member of the Poleta Formation preserve hummocky cross-stratifi ed and quasi-planar
laminated sandstones that are abruptly overlain by fl aser bedded limestones or oligomict
pebble conglomerates, representing a carbonate tidal fl at environment and a solution
collapse breccia, respectively. The absence of barrier island facies between shoreface
sandstones and back-barrier carbonates might be explained if these barrier island systems
were self-destructive in nature and were eroded on the back-barrier side and they
prograded on the fore-barrier side, resulting in translation of the barrier island. While the
translation of barrier islands has been well documented for transgressive barrier islands,
such a process for prograding barrier islands has yet to be proposed.
The self-destructive prograding barrier island model, as proposed herein, predicts
that the response of the back-barrier environment to progradation of a barrier island is
erosion. As the barrier island experiences progradation, erosion from tidal channels and
73
storms results in the dissection of the landward side of the barrier. Progradation and
expansion of the barrier’s width might also promote back-barrier sediment starvation,
further enhancing erosion. If erosion keeps pace with progradation then the barrier
island can maintain a relatively constant width, rather than growing indefi nitely until
progradation ceases. This model also predicts the absence of barrier island dune deposits
in ancient prograding barrier island succession and suggests that the contact between
underlying sandy shoreface and foreshore deposits and overlying back-barrier tidal
channel, tidal fl at, and lagoonal deposits must represent an erosional unconformity.
Further scrutiny of the nature of the contact between fore-barrier and back-barrier
deposits within prograding barrier island succession will help to refi ne and test the
robustness of this model and may provide insight into the dynamics of barrier island
migration that are not apparent from studies of modern coastlines alone.
74
CHAPTER III
Lower Cambrian Grand Cycles of the Southern Great Basin: Implications for Mixed
Carbonate-Siliciclastic Systems
Introduction
‘Grand Cycles’ are depositional units comprised of a siliciclastic half-cycle
overlain by a carbonate half-cycle that form a conformable succession bounded by sharp
contacts (Aitken, 1966). These cycles were originally recognized by Aitken (1966)
within Middle Cambrian to Middle Ordovician strata of the southern Rocky Mountains
of Alberta, Canada. Sharp contacts that bound Grand Cycles are typically regarded as
representing disconformities due to the abruptness between underlying carbonate and
overlying siliciclastics (Aitken, 1966; 1978; Mount et al., 1991), although there are
exceptions (e.g., Mount and Bergk, 1998). Basal strata within the originally defi ned
Grand Cycles are typically comprised of silty shales with minor thin-bedded carbonate
layers (Aitken, 1966). Limestone beds thicken and become increasingly more abundant
upsection as shale decreases in abundance, and the transition between the siliciclastic and
carbonate half-cycles is gradational (Aitken, 1966).
Precambrian-Cambrian strata of the southern Great Basin, United States, have
been interpreted previously to contain similar Grand Cycles; however, they are confi ned
to the Ediacaran, lower Cambrian, and Middle Cambrian (Fritz, 1975; Palmer and Halley,
1979; Mount et al., 1991; Mount and Bergk, 1998). This study focuses on those of the
lower Cambrian in the White-Inyo and Death Valley regions of the southern Great Basin
(Fig. 3.1). Three Grand Cycles—Grand Cycles A, B, and C—were originally defi ned and
correlated by Fritz (1975) within the lower Cambrian of the North American Cordillera,
including the southern Great Basin, based upon trilobite biozones. Two more Grand
Cycles—DS 1 and DS 2—were added by Mount et al. (1991) for pre-trilobite strata,
75
W ood Canyon
Zabriskie
lower
middle
upper
Deep
Spring
Campito
Poleta
Harkless
ANDREWS
MOUNT AIN
GRAPEVINE
MOUNT AINS
NOP AH
RANGE
Lower
Middle
Upper
Andrews Mountain
Montenegro
Lower
Middle
Upper
WHITE-INY O SUCCESSION DEA TH V ALLEY SUC CESSION
km
0
1
DS 1
DS 2
A
B Bonnia-Olenellus
zone
Nevadella
zone
Fallotaspis zone
Conglomerate
Sandstone
Siltstone and sandstone
Limestone
Dolostone
Trilobite biozone boundary: dashed
where approximately located
Period boundary
Cambrian
Neoproterozoic
FIGURE 3.1. Generalized regional stratigraphy of the Neoproterozoic-Cambrian
succession in the southern Great Basin showing trilobite biozone correlations and the
placement of Grand Cycle boundaries (after Nelson, 1978; Stewart, 1982; Hunt, 1990;
Mount et al., 1991; Corsetti and Hagadorn, 2000; 2003; Hollingsworth, 2005). The
Nopah Range in the Death Valley region represents the most onshore locality examined,
while Andrews Mountain in the White-Inyo region represents the most offshore locality
examined.
76
some of which are now known to be Ediacaran in age (Corsetti and Hagadorn, 2003).
Grand Cycles of the southern Great Basin differ from those previously defi ned from
the southern Rocky Mountains of Alberta in that the siliciclastic half-cycles tend to be
coarser-grained and contain far less interbedded carbonates (Mount and Bergk, 1998).
They do, however, meet the typical criteria of Grand Cycles because they contain a sharp-
based siliciclastic unit, a gradational transition between siliciclastics and carbonates, and
a capping unit of peritidal carbonates (Mount et al., 1991).
Grand Cycles DS1 and DS2 in the White-Inyo Mountains correspond to the
Middle Member and Upper Member of the Deep Spring Formation, respectively, and
each consists of a siliciclastic-carbonate succession, with shelf and shoreface siliciclastics
transitioning upward into peritidal carbonates (Mount et al., 1991; Corsetti and Hagadorn,
2003). These Grand Cycles correlate lithologically to the middle and upper carbonate-
capped siliciclastic units of the lower member of the Wood Canyon Formation in the
Death Valley region (Stewart, 1966; 1970; Corsetti and Hagadorn, 2000; 2003). Grand
Cycle A consists of the Campito Formation and overlying Lower Member of the Poleta
Formation in the White-Inyo Mountains and is lithologically correlated to the uppermost
lower member, middle member, and lowermost carbonate-capped unit in the upper
member of the Wood Canyon Formation in the Death Valley region. Grand Cycle B
consists of the Middle and Upper Members of the Poleta Formation in the White-Inyo
Mountains that correlate to the upper carbonate-capped unit in the upper member of the
Wood Canyon Formation in the Death Valley region. Grand Cycle C is the uppermost
lower Cambrian Grand Cycle in the southern Great Basin and consists of the Harkless
Formation and overlying Mule Spring Limestone in the White-Inyo Mountains and the
Zabriskie Quartzite through the Thimble Limestone Member of the Carrara Formation in
the Death Valley region (Palmer and Halley, 1979).
77
The Precambrian-Cambrian boundary within these successions is found within
the siliciclastic portion of the Upper Member of the Deep Spring Formation in the White-
Inyo Mountains and the siliciclastic portion of the upper carbonate-capped unit of the
lower member of the Wood Canyon Formation in the Death Valley region (Corsetti and
Hagadorn, 2000; 2003). In both White-Inyo and Death Valley the Precambrian-Cambrian
boundary falls within the lower portion of Grand Cycle DS 2.
While these Ediacaran-Cambrian Grand Cycles of the southern Great Basin
have been examined as large-scale depositional packages—some over a kilometer in
thickness—they have yet to be examined at a higher resolution, particularly at the facies
and parasequence scale. The purpose of this study is to address the nature of the facies
changes and environmental shifts represented by the siliciclastic-carbonate transitions of
these Grand Cycles, and to examine the sequence stratigraphic position over which this
transition takes place. To do so, this study examines the siliciclastic-carbonate transition
of Grand Cycle A within the White-Inyo and Death Valley regions and Grand Cycle B in
the White-Inyo region.
Grand Cycle A
Grand Cycle A in the Death Valley region consists of the uppermost lower
member, middle member, and lowermost upper member of the Wood Canyon
Formation. The facies change associated with the siliciclastic-carbonate transition of
this Grand Cycle is exposed within the upper member of the Wood Canyon Formation
in the southern Nopah Range at Emigrant Pass (Fig. 1.7), California, and falls within
the Fallotaspis trilobite biozone (Hunt, 1990; Mount et al., 1991). The combined
sedimentary facies of the upper member have been interpreted as a mixed carbonate-
siliciclastic barrier island complex, with shallow shelf to shoreface quartz-rich
siliciclastics and fl ood-tidal delta and lagoonal carbonates (Mata et al., in review). A
78
Bidirectional cross-stratified
sandstone
Bidirectional cross-stratified
oolite
Sandy micritic dolostone
Hummocky cross-stratified sandstone Laminated siltstone
Siltstone and hummocky
cross-stratified sandstone
Archaeocyath-skeletal
limestone
Calcimicrobe-archaeocyath
limestone
1 m
Montenegro Member Campito Lower Member Poleta
Proximal
offshore
Archaeocyath
reef
Lower
shoreface
Offshore
transition
Lower
shoreface
Ebb-tidal
delta
Ebb-tidal delta
Ebb-tidal
delta
Lower
shoreface
Offshore
transition
Lower
shoreface
FS
FS
FS
Ebb-tidal
delta
FS
FS
Lagoon
Flood-tidal
delta
Flood-tidal
delta
Flood-tidal delta
Ebb-tidal
delta
Flood-tidal
delta
1 m
upper member Wood Canyon
Carbonate half-cycle Siliciclastic half-cycle
Erosional unconformity
WHITE-INYO SUCCESSION DEA TH V ALLEY SUCCESSION
FIGURE 3.2. Stratigraphy and depositional environments of the siliciclastic-carbonate
transition of Grand Cycle A within the White-Inyo and Death Valley regions. The
transition occurred fi rst in the Death Valley region and migrated offshore toward the
White-Inyo region. Siliciclastic facies in both regions consist of upward-coarsening
shallow shelf and shoreface deposits, however, carbonate facies differ. A barrier
island complex is developed across the transition in the Death Valley region, while an
archaeocyath reef complex is developed across the transition in the White-Inyo region.
79
single complete upward-shallowing succession consists of a basal unit of interbedded
siltstone and hummocky cross-stratifi ed very fi ne sandstone of the offshore transition that
grades upward into amalgamated hummocky-cross-stratifi ed fi ne sandstone of the lower
shoreface (Fig. 3.2). This is capped by bidirectional small-scale cross-stratifi ed medium
sandstone of an ebb-tidal delta. Ebb-tidal delta deposits are locally incised, representing
incision of a tidal inlet, and are overlain by medium-scale bidirectional cross-stratifi ed
oolite of a fl ood-tidal delta and heavily bioturbated sandy micritic dolostone of a lagoonal
environment (Fig. 3.3A-B). Parasequence stacking patterns are progradational leading
into the carbonate half-cycle of the upper member, shifting from siliciclastic-dominated
to carbonate-dominated, although the carbonate half-cycle still contains siliciclastic facies
(Fig. 3.2).
Grand Cycle A in the White-Inyo Mountains consists of the Campito Formation
and the Lower Member of the Poleta Formation. The siliciclastic-carbonate transition
occurs within the Nevadella trilobite biozone (Nelson, 1978), a biozone above this same
transition in the Death Valley region (Mount et al., 1991). The Campito Formation is a
predominantly siliciclastic succession with carbonate facies occurring within the upper
portion of its Montenegro Member. Siliciclastic facies are dominated by amalgamated
quartz-rich hummocky cross-stratifi ed sandstones and non-amalgamated hummocky
cross-stratifi ed sandstones interbedded with laminated siltstones and mudstones (Mount,
1982). These facies represent deposition within a storm-dominated shallow shelf and
shoreface, with amalgamated hummocky cross-stratifi ed sandstones preserved within the
lower shoreface and non-amalgamated beds preserved in the offshore transition between
fair-weather wave base and storm wave base. Carbonates of the Montenegro Member
consist of archaeocyathan limestones that form layers, lenses, and bioherms (Fuller, 1972;
Morgan, 1976) and occur as isolated carbonate bodies within hummocky cross-stratifi ed
fi ne to very fi ne sandstone of a lower shoreface environment (Fig. 3.3).
80
B
B C
A
1 cm
UP
UP
FIGURE 3.3. Outcrop photographs of the large carbonate body, interpreted as an
archaeocyath bioherm, found within the upper portion of the Montenegro Member of
the Campito Formation. A) The isolated carbonate body is enclosed within siliciclastic
strata consisting of hummocky cross-stratifi ed sandstone that is found below and occurs
laterally. Dashed red line represents marks the contact between underlying siliciclastics
and overlying carbonate. Box marks the location of fi g. 3.4B. B) Close-up view of the
contact between sandstone (lower left) and carbonate (upper right) marked by dashed red
line. C) Bedding surface view of in situ archaeocyaths within the carbonate body.
81
The overlying Lower Member of the Poleta Formation is comprised of
archaeocyathan-calcimicrobe and archaeocyathan-thrombolite boundstone interbedded
with lime mudstone, skeletal wackestone, and oolite (Rowland, 1984; Rowland and
Gangloff, 1988). The boundstone of the Lower Member has been interpreted as a reef
complex, with the calcimicrobes Renalcis and Epiphyton forming the reef fabric, which
has been shown to contain large meter-scale micrite-fi lled primary cavities with Renalcis
pendants suggestive of a rigid, wave-resistant framework (Rowland and Gangloff, 1988).
The siliciclastic-carbonate transition of Grand Cycle A is marked by the contact
between the Campito Formation and the Lower Member of the Poleta Formation. The
contact is irregular within the White-Inyo region because it represents diachronous
colonization of the seafl oor by different styles of archaeocyathan buildups (Rowland,
1978) rather than a fundamental lithologic change within the depositional system. The
siliciclastic-carbonate transition, examined at the northern end of Cedar Flat (Fig.
1.7), consists of an ~6 meter thick upward-coarsening siliciclastic succession that
transitions from siltstone with thin (1-2 cm) laminated very fi ne sandstone layers and
lenses of a proximal offshore environment, below storm wave base, to siltstone with
upward-thickening hummocky cross-stratifi ed and quasi-planar laminated very fi ne
sandstone beds of the offshore transition, between fair-weather and storm wave base
(Fig. 3.2). Associated with offshore transition deposits are rare hummocky cross-
stratifi ed archaeocyathan grainstones. The uppermost portion of the Campito Formation
is comprised of amalgamated hummocky cross-stratifi ed fi ne sandstones of the lower
shoreface with rare archaeocyathan grainstone lenses. The lowermost Poleta Formation
consists of several beds of hummocky cross-stratifi ed archaeocyathan grainstone with
coarse basal skeletal lags (Fig. 3.4C). This transitions upward into archaeocyathan-
thrombolite boundstone and skeletal wackestone with scattered fossil debris and in situ
archaeocyathans (Fig. 3.4D-E). These archaeocyathan-thrombolite boundstones represent
82
basal lag
reactivation
surface
ebb-tidal
delta
flood-tidal
delta
lagoon
thrombolite
archaeocyath
thrombolite
archaeocyath
A B
C
D E
FIGURE 3.4. Sedimentary facies of Grand Cycle A. A) Flood-tidal delta deposits
consisting of medium-scale bidirectional cross-stratifi ed oolite with reactivation surfaces
within the upper member of the Wood Canyon Formation at the Emigrant Pass locality
in the Death Valley region. B) Incision of channel within planar laminated ebb-tidal
delta deposits of the upper member of the Wood Canyon Formation. This channel is
fi lled with oolitic fl ood-tidal delta deposits that are overlain by sandy dolostone of a
lagoonal environment. C) Lowermost hummocky cross-stratifi ed grainstone of the Poleta
Formation overlying uppermost Campito sandstone at the Cedar Flat locality, White-Inyo
region. Notice archaeocyath lag at base. D) Bedding surface view of in situ archaeocyath
within lowermost Poleta Formation thrombolitic boundstone. E) Polished slab of sample
found in fi gure 3.3D showing embedded archaeocyathan within thrombolitic matrix.
83
the colonization stage of the earliest archaeocyathan reefs in the region (Rowland and
Gangloff, 1988).
Mount and Bergk (1998) have suggested that siliciclastic-carbonate transitions
within each Grand Cycle initiated onshore and migrated offshore over time because these
transitions are younger in the offshore facies of the White-Inyo Mountains than in the
onshore facies of the Death Valley region. The corollary to this is that the siliciclastic-
carbonate transition within Grand Cycle A represents one or more progradational events
that traversed the distance from the Death Valley region to the White-Inyo Mountains.
Palinspastic reconstructions place this distance during the Precambrian-Cambrian
transition at approximately 150 kilometers (Levy and Christie-Blick, 1989; 1991), which
is comparable to other ancient barrier island systems that are known to have prograded
long distances (i.e., over 100 km) (e.g., Leckie, 1985; Plint and Walker, 1987).
If this transition does indeed represent a single progradational event, then it
is noteworthy because it shows that archaeocyathans changed their ecology over the
time it took for the coastline to prograde from one side of the basin to the other, within
less than two trilobite biozones. Archaeocyathans within the upper member of the
Wood Canyon Formation in the southern Nopah Range are found only as transported
fossil accumulations within back-barrier fl ood-tidal delta and lower shoreface ebb-
surge deposits (Fig. 3.5A; Mata et al., in review). However, in younger strata of the
Montenegro Member of the Campito Formation, archaeocyathans can be found as in
situ components of bioherms (Fig. 3.5B; Fuller, 1972; Morgan, 1976; Mount and Signor,
1991). Capping the Campito Formation, archaeocyathans within the Lower Member of
the Poleta Formation occur in association with calcimicrobe and thrombolite frameworks
of a reef complex (Fig. 3.5C); Rowland, 1984; Rowland and Gangloff, 1988). This
pattern suggests dramatic changes in archaeocyathan ecology, from transported
fossils to bioherms to reefs, over a brief biostratigraphic interval, although the overall
84
Fair-weather wave base
Storm wave base
Fair-weather wave base
Storm wave base
Flood-
tidal
delta
Lagoon
Ebb-
tidal
delta
Lower
shoreface
Offshore
transition
Lower
shoreface
Bioherm
Fair-weather wave base
Storm wave base
Lagoon
Calcimicrobe-
archaeocyath
reef
Lower
shoreface
Offshore
transition
Offshore
upper member Wood Canyon
Campito-Poleta transition
Montenegro Member Campito
C
B
A
FIGURE 3.5. Diagrammatic illustration of the shifting environmental setting during the
progradation of the siliciclastic-carbonate transition of Grand Cycle A and associated
archaeocyath paleoecology. The siliciclastic-carbonate transition is fi rst estalished in the
Death Valley region with the formation of a barrier island complex in the upper member
of the Wood Canyon Formation. Archaeocyaths occur as tranported body fossils within
fl ood tidal delta and ebb-surge deposits within the lower shoreface. Archaeocyath then
occur as in situ components within isolated bioherms of the Montenegro Member of the
Campito Formation, just below the siliciclastic-carbonate transition in the White-Inyo
Mountains. The siliciclastic-carbonate transition in the White-Inyo Mountains is marked
by the development of a calcimicrobe-archaeocyath reef complex that initiated atop lower
shoreface deposits. Lithologic symbols are the same as in fi g. 3.2.
85
environmental settings are quite similar. All archaeocyaths found within the siliciclastic-
carbonate transition of Grand Cycle A occur in nearshore, high-energy carbonate settings
above fair-weather wave, either within the lower shoreface or landward of it.
Grand Cycle B
Grand Cycle B in the White-Inyo Mountains consists of the Middle Member and
Upper Member of the Poleta Formation. The Middle Member of the Poleta Formation
is comprised of shale, siltstone, and quartz-rich sandstones with limestone interbeds.
The Upper Member consists of cherty and silty limestones with micrite or quartz-fi lled
cavities and chicken wire structures. Near the northeast end of the Poleta Folds region of
the White-Inyo Mountains a canyon exposes a complete outcrop of the Middle Member
and the lower portion of the Upper Member. This has allowed for a complete sequence
stratigraphic analysis of the facies stacking patterns across this siliciclastic-carbonate
transition.
A single complete upward-shallowing succession within the Middle Member
of the Poleta Formation consists of a basal unit of laminated to bedded siltstone with
thin (< 2 cm) laminated sandstone layers and lenses, interpreted as a proximal offshore
environment. These sandstone beds thicken upward into hummocky cross-stratifi ed
and quasi-planar laminated layers interbedded with siltstone, interpreted as the offshore
transition. Upward, hummocky cross-stratifi ed and quasi-planar laminated sandstone
beds become amalgamated within deposits of the lower shoreface. These lower shoreface
deposits are capped by fl aser bedded limestone and siltstone, sometimes occurring across
an erosional unconformity. At the Middle-Upper Member contact, lower shoreface
deposits are capped by cherty limestones with micritic pebbles that form an oligomict
conglomerate. Associated with these strata are micrite or quartz-fi lled cavities over
which is found irregular, synsedimentary folding. The fl aser bedded limestones of the
86
Middle Member have been interpreted as a back-barrier tidal fl at of a barrier island
coastline, while the cherty limestones and oligomict conglomerates have been interpreted
as a solution collapse breccia in association with a supratidal fl at environment (Mata and
Bottjer, in review). The absence of barrier island deposits separating fore-barrier and
back-barrier facies is likely due to recycling of the barrier during progradation, which has
been suggested to erase the barrier deposits from the record and superimpose back-barrier
deposits over shoreface deposits across an erosional unconformity (Mata and Bottjer, in
review).
The facies stacking pattern within the Middle and Upper members of the Poleta
Formation suggests that the maximum fl ooding surface lies within the siliciclastic half-
cycle represented by the Middle Member (Fig. 3.6). A retrogradational stacking pattern
occurs within the lowermost ~9 meters of the Middle Member, while a progradational
stacking pattern is found within the upper ~12 meters of the Middle Member and
continues into the carbonate half-cycle of the Upper Member. It is therefore interpreted
that the lower ~9 meters of the Middle Member falls within the transgressive systems
tract and that the upper ~12 meters of the Middle Member and siliciclastic-carbonate
transition of the Middle-Upper Member contact fall within the highstand systems
tract. This would indicated that the maximum fl ooding surface occurs midway through
the siliciclastic half-cycle of Grand Cycle B, rather than at the siliciclastic-carbonate
transition, as has been suggested previously for other Grand Cycles (e.g., Mount and
Bergk, 1998).
Implications for Mixed Carbonate-Siliciclastic Systems
The mixed carbonate-siliciclastic sedimentary systems of the early Cambrian
documented in this study differ signifi cantly from those proposed for the region
previously (e.g., Palmer and Halley, 1979; Mount et al., 1991; Mount and Bergk, 1998),
87
FIGURE 3.6. Stratigraphic section measured through the Middle Member and lowermost
Upper Member of the Poleta Formation at the Poleta folds locality. Parasequence stacking
patterns are retrogradational in the lower third of the Middle Member and are interpreted
as the transgressive systems tract. This is capped by a maximum fl ooding surface and
subsequent progradational stacking patterns of the highstand systems tract leading
through the Middle-Upper Member contact. FS = fl ooding surface; MFS = maximum
fl ooding surface.
Lower
shoreface
Offshore
transition
Proximal
offshore
Offshore transition
Lower shoreface
Tidal flat
Lower shoreface
Tidal flat
Proximal
offshore
Offshore
transition
Lower shoreface
Offshore transition
Lower shoreface
Tidal flat
Proximal
offshore
Offshore
transition
Lower
shoreface
Supratidal flat
FS
FS
MFS
FS
Solution collapse
breccia
FS
Middle Member Poleta
Upper
Member
Highstand
Systems
Tract
Transgressive
Systems
Tract
2 m
FS
Siltstone and hummocky
cross-stratified sandstone
Bedded to laminated siltstone
with thin sandstone beds
Micritic limestone
Hummocky cross-stratified
sandstone
Erosional unconformity
Chert laminae
Chert layers with micritic pebbles
Silt flasers
Micrite or quartz-filled cavities
Chicken wire structures
88
and for other Cambrian mixed systems associated with Grand Cycles (Aitken, 1966;
1978; Chow and James, 1987). The main difference lies in the locus of carbonate
production in relation to siliciclastic environments. Middle Cambrian to Middle
Ordovician Grand Cycles from the Canadian Rocky Mountains of Alberta, originally
defi ned by Aitken (1966), have been interpreted as the result of shifts in facies belts,
including: i) an inner detrital belt of shale, siltstone, and subordinate carbonate
representing an inshore basin; ii) a middle carbonate belt with rare siltstone and sandstone
beds representing a shoal complex; and iii) an outer detrital belt with thin-bedded shale
and argillaceous limestones representing an open marine basin. Grand Cycles are
believed to result from onshore migration of these facies belts, as the more offshore
middle carbonate facies belt migrates landward over the inshore basin of the inner detrital
belt to form a siliciclastic-carbonate couplet.
Strata of the upper member of the Wood Canyon Formation, the Montenegro
Member of the Campito Formation, and the Poleta Formation reveal a different origin
for lower Cambrian Grand Cycles of the southern Great Basin that makes them markedly
different from other defi ned Grand Cycles. Specifi cally, they are notably different from
those defi ned in the Middle Cambrian to Middle Ordovician of the Canadian Rocky
Mountains (Aitken, 1966; 1978), the Middle-Upper Cambrian of western Newfoundland
(Chow and James, 1987), and the Middle Cambrian of the southern Great Basin (Palmer
and Halley, 1979), that rely on an inshore or intrashelf siliciclastic basin bounded seaward
by a carbonate shoal complex. Lower Cambrian strata of the White-Inyo and Death
Valley regions reveal that carbonate facies were deposited landward of siliciclastic facies
within these Grand Cycles. It is also of note that the siliciclastic-carbonate transition
of each Grand Cycle examined does not represent the same type of environmental shift,
even within a single Grand Cycle.
89
Lower Cambrian Grand Cycles in the southern Great Basin represent at least three
different types of siliciclastic-carbonate transitions. The transition of Grand Cycle A in
the Death Valley region represents progradation of a barrier island coastline with shallow
shelf and shoreface siliciclastics overlain by carbonate fl ood-tidal delta and lagoonal
deposits of a tidal inlet complex. This same Grand Cycle transition within the White-
Inyo Mountains represents the development of an archaeocyathan reef-lagoon complex
atop siliciclastic lower shoreface deposits. The transition of Grand Cycle B in the White-
Inyo Mountains represents a similar progradation of a barrier island coastline, as in Grand
Cycle A in the Death Valley region; however, there are no tidal inlet fi lls or tidal delta
deposits. Rather, siliciclastic shallow shelf and shoreface deposits are abruptly overlain
by back-barrier carbonate tidal fl at deposits.
The siliciclastic-carbonate transition within lower Cambrian Grand Cycles of the
southern Great Basin have been interpreted previously to occur at the maximum fl ooding
surface, which would represent the highest levels of sediment starvation (Mount and
Bergk, 1998). This interpretation was, however, based on large-scale stratigraphy, rather
than parasequence stacking patterns. At the facies and parasequence scale, it is apparent
that the maximum fl ooding surface occurs within the siliciclastic half-cycle and that
progradational stacking patterns lead upward into the siliciclastic-carbonate transition.
This suggests that the turnover from siliciclastics to carbonates was occurring within the
highstand systems tract and most likely during progradation of the coastline.
Conclusions
Understanding the depositional dynamics of Cambrian Grand Cycles has
important implications not only for mixed carbonate-siliciclastic sedimentation, but
also for evolutionary and paleoecological patterns. While siliciclastic facies within
lower Cambrian Grand Cycles are highly conservative and recur in many formations,
90
carbonate facies are highly variable, even within a single Grand Cycle. Siliciclastic
facies represent environments from shallow shelf to shoreface, while carbonate facies
represent a wide range of environments including tidal inlet complex, back-barrier tidal
fl ats, and archaeocyath reef-lagoon complex that are unique to specifi c Grand Cycles.
The transitions between siliciclastic and carbonate half-cycles are time-transgressive
and record the seaward progradation of carbonate environments over shallow shelf and
shoreface siliciclastic environments following maximum fl ooding midway through the
siliciclastic half-cycle. The time-transgressive nature of these transitions allows for a
record of paleoecological changes over the course of a progradational episode. Grand
Cycle A shows that archaeocyathans acquired new levels of ecology from level-bottom
forms to bioherm components to reef builders over a single siliciclastic-carbonate
progradational episode across the basin, spanning less than two trilobite biozones—less
than 6 million years in total duration (e.g., Peng and Babcock, 2008). While such a
transition might be accounted for by onshore-offshore patterns or distributions, a more
parsimonious conclusion is that the environment remained conservative, while the
ecology changed within it. This seems a valid conclusion considering all archaeocyath
occurrences are within shallow carbonate settings above fair-weather wave base,
found either within the lower shoreface or landward of it. Grand Cycles are highly
unique depositional episodes that can provide a wealth of both sedimentological and
paleoecological information, and continued refi ning of Grand Cycle depositional
models may provide further insight into the depositional dynamics of mixed carbonate-
siliciclastic systems and could shed new light on the environmental patterns of the
Cambrian radiation.
91
CHAPTER IV
Facies Control on Lower Cambrian Wrinkle Structure Development and
Paleoenvironmental Distribution, Southern Great Basin, United States
Introduction
Modern microbial mats can be found in a wide range of environments including
tidal fl ats, lagoons, outer shelf to slope environments, and deep ocean basins in
siliciclastic settings (e.g., Horodyski et al., 1977; Soutar and Crill, 1977; Javor and
Castenholz, 1984; Gerdes et al., 1993; Noffke, 1998; Jørgensen and Gallardo, 1999;
Bertics and Ziebis, 2009). Microbial mats can also be found in lagoons, tidal channels
and fl ats, and sabkhas in carbonate settings (Logan, 1961; Logan et al., 1974; Hardie and
Ginsburg, 1977; Dravis, 1983; Dill et al., 1986). In mineralizing carbonate environments
the distribution of microbial mats is typically recorded by cemented microbialites, which
have a high preservation potential due to early lithifi cation at or near the seafl oor (e.g.,
Dravis, 1983; Dupraz and Visscher, 2005), even if the microbial mat itself leaves no
organic record. In siliciclastic settings, however, microbially mediated sedimentary
structures require exceptional preservational conditions because physical and biological
processes can destroy the soft microbial mat before burial and subsequent lithifi cation
(e.g., Javor and Castenholz, 1984; Fenchel, 1998).
Of paramount importance to the destruction of microbial mats is bioturbation,
which is the disruption and mixing of sediment due to organisms. In most modern
oxygenated subtidal settings the upper sediment column (~10 cm depth) is constantly
and thoroughly mixed due to biological activity, developing a surface mixed layer (e.g.,
Berger et al., 1979). This perpetual reworking can inhibit microbial mat development
at the sediment surface and prevent preservation of a surface microbial mat upon burial.
This destruction can occur through grazing and consumption of the mat at the sediment
92
surface (Javor and Castenholz, 1984; Fenchel, 1998) or physical disruption of the layered
mat in the shallow or deep subsurface, which can destroy essential chemical gradients
important for biogeochemical cycling (Paerl and Pinckney, 1996). While these biological
aspects have received much study, the physical and chemical processes that inhibit
microbial mat growth and preservation are less well understood.
One process by which microbial mats may be excluded from an environment is
through physical erosion of the mat due to high-energy depositional conditions (e.g.,
Neumann et al., 1970; Hagadorn and McDowell, 2012). Excessive wave stress or tidal
energy may lead to the erosion of a pre-existing mat on a stable substrate or inhibit
development of a mat on a frequently reworked substrate (Ginsburg and Planavsky, 2008;
Mata and Bottjer, 2009b). An additional physical parameter is irradiance reaching the
sediment surface, which is an important parameter for phototrophic microbial mats in
subtidal settings, but is less signifi cant for non-phototrophic mats or for intertidal mats
that are periodically exposed and receive suffi cient light for photosynthesis (Mata and
Bottjer, 2009b).
Addressing the roles that these physical parameters play on microbial mat
distribution in modern environments is hindered by the presence of extensively developed
bioturbation in most oxygenated subtidal marine settings. This is the case throughout
much of the Phanerozoic in which metazoan activity effectively precluded extensive
microbial mat development (Bottjer et al., 2000; Mata and Bottjer, 2009b; 2011). The
physical factors affecting microbial mat distribution must therefore be examined during
times in Earth’s history in which bioturbation was not the primary control on microbial
mat distribution (e.g., Noffke et al., 2002). This restricts the window for observing
the role of these physical processes to the Cambrian and Precambrian, prior to the
development of extensively developed and deep-penetrating bioturbation (e.g., Bottjer et
al., 2000; Marenco and Bottjer, 2008).
93
The Cambrian explosion and Ordovician radiation mark a signifi cant turning
point for the paleoenvironmental distribution of microbial mats throughout Earth’s
history. Microbially mediated sedimentary structures in siliciclastic settings are abundant
and diverse in Precambrian and Cambrian shallow shelf and marginal marine deposits,
but are rare thereafter (Mata and Bottjer, 2009b). Extent of vertical bioturbation shows
initial increases in the early Cambrian, although it isn’t until the Late Ordovician that
bioturbation reaches levels more comparable to those found throughout the remainder
of the Phanerozoic (Droser and Bottjer, 1989). The post-Cambrian Phanerozoic
paleoenvironmental distribution of siliciclastic microbial mat structures is much more
restricted and most occurrences are found primarily within tidal fl at and marginal marine
settings, except during the aftermath of mass extinction events (Mata and Bottjer, 2009b;
2011). This distribution refl ects increases in ecospace utilization that likely began in the
Ordovician and were halted only temporarily during the aftermath of mass extinctions in
which ecological constraints were relaxed (Schubert and Bottjer, 1992; Mata and Bottjer,
2011).
The Precambrian-Cambrian transition in the southern Great Basin consists of a
mixed siliciclastic-carbonate succession that has been studied extensively with regard to
sedimentology and stratigraphy (e.g., Nelson, 1962; Stewart, 1970; Mount et al., 1991;
Corsetti and Hagadorn, 2003) and trends in bioturbation (Droser and Bottjer, 1988;
1989; Marenco and Bottjer, 2008). While the succession is known to yield microbial
mat structures within its siliciclastic facies (e.g., Hagadorn and Bottjer, 1997; 1999;
Bailey et al., 2006), there has yet to be a systematic study of the paleoenvironmental
distribution of these features. The most common microbial mat features throughout
the southern Great Basin succession are wrinkle structures, which consist of mm-scale
sinuous fl at-topped ridges and troughs and represent the former presence of a surface
microbial mat (Hagadorn and Bottjer, 1997). Wrinkles structures are most extensively
94
developed within the lower Cambrian portion of the succession, which predates the
development of extensive and vertically-oriented bioturbation within siliciclastic settings
for the region (Marenco and Bottjer, 2008). The purpose of this study is to document the
paleoenvironmental distribution of wrinkle structures within lower Cambrian strata of the
southern Great Basin to examine the physical and taphonomic controls that may affect the
formation and preservation of microbial mats in siliciclastic shallow shelf and nearshore
settings and explore the nature of the early Cambrian substrate.
Geological Setting and Methods
The Precambrian-Cambrian succession of the southern Great Basin are divided
into four interfi ngering successions that span the ancient craton to shelf, and included
from onshore to offshore are the Craton, Mojave (Craton Margin), Death Valley, and
White-Inyo successions (Nelson, 1978; Corsetti and Hagadorn, 2000; Fedo and Cooper,
2001). This study focuses on the Death Valley and White-Inyo successions that are found
within a widely distributed Precambrian-Cambrian outcrop belt in the southern Great
Basin (Fig. 1.7).
Siliciclastic facies predominate in the lower Cambrian portion of the Death Valley
and White-Inyo successions and represent a wide range of depositional environments,
including deep shelf, shallow shelf, shoreface, tidal fl at, and fl uvial environments
(e.g., Klein, 1971; Mount, 1982; Mount and Signor, 1985; Fedo and Cooper, 1990;
Prave, 1992). Carbonate facies represent subtidal bioherm, reef, lagoonal, and tidal fl at
environments (e.g., Moore, 1976; Rowland, 1984; Mount and Signor, 1985; Mata et al.,
in review). This study focuses on the middle member of the Wood Canyon Formation
in the Death Valley region, and the Montenegro Member of the Campito Formation, the
Middle Member of the Poleta Formation, and the Harkless Formation in the White-Inyo
Mountains (Fig. 4.1).
95
FIGURE 4.1. Generalized regional stratigraphy of the Neoproterozoic-Cambrian
succession in the southern Great Basin showing the stratigraphic intervals examined
in this study (after Nelson 1978; Stewart 1982; Corsetti and Hagadorn 2000; 2003). 1
= Campito-Poleta contact; 2 = Middle Members of the Poleta Formation; 3 = Harkless
Formation; 4 = middle member of the Wood Canyon Formation.
2
WHITE-INYO REGION
DEATH VALLEY REGION
3
1
4
Neoproterozoic
Cambrian
W ood Canyon
Zabriskie
Carrara
lower
middle
upper
Deep
Spring
Campito
Poleta
Harkless
km
0
1
Lower
Middle
Upper
Andrews Mountain
Montenegro
Lower
Middle
Upper
Conglomerate
Sandstone
Siltstone + sandstone
Limestone
Dolostone
96
Sedimentary facies were defi ned for each of these formations and depositional
environments were interpreted. Wrinkle structures were then placed into this facies and
depositional context and were not used to aid or infl uence environmental interpretation,
so as to avoid circular reasoning in addressing their environmental distribution. Partial
stratigraphic sections were measured to show facies stacking patterns, as well as to
precisely place wrinkle structures at a bed-scale resolution within facies.
Death Valley Succession - Middle Member Wood Canyon Formation
The middle member of the Wood Canyon Formation is Terreneuvian in age and
occurs above the fi rst occurrence of Treptichnus pedum, but below the fi rst occurrence
of trilobites in the region (e.g., Corsetti and Hagadorn, 2000; 2003). The lower portion
of the middle member consists of conglomerate and conglomeratic quartzite with
interbedded siltstone, interpreted primarily as braided fl uvial environments with rare
tidally infl uenced fl uvial environments (Stewart, 1970; Diehl 1974; 1979; Fedo and
Cooper, 1990). A general fi ning-upward trend exists within the middle member and its
upper portion consists primarily of interbedded quartzite and siltstone, interpreted as a
mixed tidal fl at environment (Klein, 1971). The middle member was examined at the
Southern Salt Spring Hills, located at the southeast end of the Death Valley region.
The measured partial stratigraphic section consists of upward-fi ning and upward-
thinning sandstones interbedded with siltstone and sandy siltstone (Fig. 4.2). A complete
upward-fi ning succession is comprised of a basal unit of amalgamated planar laminated
and low-angle cross-stratifi ed fi ne sandstone with rare herringbone cross-stratifi cation,
ripple cross-lamination, and fl aser bedding (Fig. 4.3A); a middle unit of sandy siltstone
interbedded with planar laminated and low-angle cross-stratifi ed fi ne to very fi ne
sandstone with rare wavy bedding (Fig. 4.3B); and an upper unit of sandy siltstone with
97
Sandy siltstone
Low-angle cross-stratified sandstone
Cross-laminated sandstone with silt flasers
Planar laminated sandstone
FS
VC
C
M
F
VF
Silt
Shale
Lithology
FS
Lower flat
Lower flat
Lower flat
Mid flat
Mid flat
Mid flat
FS
Mid flat
Upper flat
Upper flat
0
1
2
3
4
5
6
7
8
9
10
11
12
13
14
meters
middle member Wood Canyon
Wrinkle Structures
Planolites
FIGURE 4.2. Partial stratigraphic section measured through the middle member of the
Wood Canyon Formation at the southern Salt Spring Hills and interpreted depositional
settings. Base of section begins at N 35° 36.709 ′, W 116° 16.061 ′. FS = fl ooding surface.
98
thin (< 5 cm) cross-laminated very fi ne sandstone layers and lenses (Fig. 3C). This facies
succession represents a prograding mixed tidal fl at setting (e.g., Klein, 1975).
Mixed tidal fl ats are characterized by an onshore fi ning of sediment from
amalgamated planar laminated and ripple cross-laminated sands with fl aser bedding of
a lower tidal fl at environment that transition landward into interbedded sand and mud
with wavy bedding of a mid tidal fl at environment, and lastly mud with lenticular sands
and pinstripe bedding of an upper tidal fl at environment (Reineck, 1967; Klein, 1971).
The rhythmic alternations of sand and mud result from transportation and deposition of
sand from traction during tidal fl ow, while mud settles out of suspension during slack-
water conditions (Reineck and Wunderlich, 1968). The succession produced by the
progradation of a mixed tidal fl at setting consists of the superposition of progressively
onshore facies and results in an upward fi ning succession from sand-dominated to mud-
dominated (Klein, 1975). Upward-fi ning successions within the middle member of the
Wood Canyon Formation are therefore interpreted as the progradation of a mixed tidal
fl at with planar laminated and low-angle cross-stratifi ed sandstones interpreted as a lower
tidal fl at, interbedded sandy siltstone and planar laminated sandstone beds interpreted as
a mid tidal fl at, and sandy siltstone with thin layers and lenses as an upper tidal fl at (e.g.,
Klein, 1971; 1975).
Wrinkle structures within the middle member of the Wood Canyon Formation
occur primarily on thick (~10 cm-thick) planar laminated and low-angle cross-laminated
sandstone beds deposited within lower and mid tidal fl at environments (Fig. 4.4A-B).
Wrinkle structures exhibit a patchy distribution and can be found on the same bedding
planes as the simple horizontal trace fossil Planolites.
99
flaser
bedding
A
B
C
5 cm
5 cm
FIGURE 4.3. Sedimentary facies from middle member of the Wood Canyon Formation,
southern Salt Spring Hills locality. A) Outcrop photograph of planar laminated and fl aser
bedded fi ne sandstone interbedded with siltstone partings and layers; lower tidal fl at. B)
Outcrop photograph of a trough cross-stratifi ed fi ne sandstone bed with an erosional base
within planar laminated siltstone; mid tidal fl at. Hammer is 24 cm long. C) Interbedded
planar laminated siltstone with thin (< 5 cm) cross-laminated very fi ne sandstone layers
and lenses; upper tidal fl at.
100
A
B
1 cm
FIGURE 4.4. Outcrop photographs of wrinkle structures from the middle member of
the Wood Canyon Formation, southern Salt Spring Hills locality. A) Bedding surface
view of wrinkle structures found on a 35 cm thick planar laminated and low-angle cross-
laminated fi ne sandstone bed; approximately 8.2 meters above base of partial section. B)
Bedding surface view of wrinkle structures found on a 30 cm thick planar laminated fi ne
to very fi ne sandstone bed; approximately 13.8 meters above base of partial section. Scale
bar = 4 cm.
101
White-Inyo Succession - Montenegro Member Campito Formation
The Montenegro Member of the Campito Formation belongs to Cambrian Stage 3
of Series 2 and lies above the fi rst occurrence of trilobites, but below the fi rst occurrence
of the trilobite genus Olenellus. The member consists of interbedded shale, quartz-rich
siltstone, and fi ne-grained sandstone with lenses and layers of archaeocyathan limestone
near its top (Nelson, 1962). The Montenegro Member was examined at the northwest
edge of Cedar Flat, within the White-Inyo Mountains, where it contacts the Lower
Member of the Poleta Formation. The measured partial section spans the Campito-Poleta
contact, although wrinkle structures are only found in the Campito Formation (Fig. 4.5).
The uppermost ~6 meters of the Montenegro Member consists of three units
that stack to form an overall upward-coarsening and thickening siliciclastic succession,
which transitions upward into the lowermost Poleta Formation limestones. The basal
unit of the Campito-Poleta transition consists of layered and laminated siltstone with
thin (1-2 cm thick) planar laminated very fi ne sandstone beds with sharp bases (Fig.
4.6A). The overlying unit consists of layered siltstone with interbedded quasi-planar
laminated (sensu Arnott, 1993) and hummocky cross-stratifi ed very fi ne sandstone beds
that thicken upsection (Fig. 4.6B). Interbedded with the siltstones of this unit are also
hummocky cross-stratifi ed archaeocyathan grainstone beds. The uppermost unit of the
Campito Formation consists of amalgamated hummocky cross-stratifi ed fi ne and very
fi ne sandstone (Fig. 4.6C). Sandstones are commonly erosive-based, while bed tops are
locally truncated. Within this unit are also found rare archaeocyathan grainstone lenses
and layers that are massively bedded. The lowermost Poleta Formation consists of a
thin zone (~30 cm thick) that is comprised of hummocky cross-stratifi ed archaeocyathan
grainstone with coarse fossil lags at the base of each bed. This zone is overlain by clotted
thrombolites with in situ archaeocyathans that typify the basal Lower Member of the
Poleta Formation (Rowland and Gangloff, 1988).
102
VC
C
M
F
VF
Silt
Shale
Lithology
Proximal
offshore
Offshore
transition
Lower
shoreface
Archaeocyath
reef
0
1
2
3
4
5
6
7
8
9
10
meters
Archaeocyath thrombolitic limestone
Non-amalgamated quasi-planar laminated
and hummocky cross-stratified sandstone
Amalgamated hummocky cross-stratified sandstone
Montenegro Member Campito Lower Member Poleta
11
12
Bioclastic archaeocyath limestone
FIGURE 4.5. Partial stratigraphic section measured through the Campito-Poleta
Formation transition at the northern end of Cedar Flat and interpreted depositional
settings. Base of section begins at N 37° 18.070 ′, W 118° 08.834 ′.
103
5 cm
A
C
B
D
FIGURE 4.6. Outcrop photographs of facies and wrinkle structures from the Montenegro
Member of the Campito Formation at the northern end of Cedar Flat. A) Planar laminated
siltstone interbedded with thin (1-2 cm thick) planar laminated or massively bedded very
fi ne sandstone layers; proximal offshore. Staff scale is in decimeters. B) Interbedded
planar laminated siltstone and quasi-planar laminated very fi ne sandstone; offshore
transition. C) Amalgamated hummocky cross-stratifi ed fi ne sandstone; lower shoreface.
D) Bedding surface view of wrinkle structures found on a hummocky cross-stratifi ed very
fi ne sandstone bed that is capped by siltstone; offshore transition.
104
The Campito-Poleta transition is interpreted to represent an upward-shoaling
siliciclastic shallow shelf to shoreface succession overlain by the initial stages of an
archaeocyathan reef complex. The lowermost unit of the transition is interpreted as
a proximal offshore environment, below storm wave base, but within the reach of
storm generated currents. Thin planar laminated or normally graded sandstone layers
within fi ne-grained shelf strata are typically the result of deep storm processes in which
suspended sandy sediment entrained by storms is deposited by gravity fl ows as graded
rhythmites (e.g., Reineck and Singh, 1972). The overlying unit of interbedded siltstone
and hummocky cross-stratifi ed sandstone is interpreted as the offshore transition,
between fair-weather wave base and storm wave base. Siltstones were deposited out
of suspension, while hummocky cross-stratifi ed and quasi-planar laminated sandstones
are the result of offshore-directed storm transport of sand due to the interaction of
large storm waves and storm-generated currents (e.g., Aigner and Reineck, 1982; Dott
and Bourgeois, 1982; Dumas and Arnott, 2006). Wrinkle structures are found atop a
hummocky cross-stratifi ed very fi ne sandstone bed, beneath less resistant siltstone, of
this offshore transition environment (Fig. 4.6D). The uppermost unit of the Campito
Formation consisting of amalgamated hummocky cross-stratifi ed sandstone is interpreted
as a lower shoreface environment, above fair-weather wave base within the zone of wave
shoaling where tempestites are commonly amalgamated (Harms et al., 1975; MacEachern
and Pemberton, 1992). Basal hummocky cross-stratifi ed archaeocyathan grainstones of
the Poleta Formation refl ect a similar lower shoreface environment, while the overlying
thrombolites with in situ archaeocyathans represent the initiation of an archaeocyathan
reef complex that is extensively developed within the Lower Member of the Poleta
Formation (Rowland, 1984; Rowland and Gangloff, 1988).
105
White-Inyo Succession - Middle Member Poleta Formation
The Middle Member of the Poleta Formation contains the boundary between
Cambrian Stage 3 and 4, marked by the fi rst occurrence of the trilobite genus
Olenellus within its uppermost portion (Nelson, 1978). The Middle Member consists
of interbedded shale, siltstone, and quartz-rich sandstones with minor interbedded
limestones (Nelson, 1962). The Middle Member was examined at the northeast end of
the Poleta folds region of the White-Inyo Mountains, and the partial section measured
records its uppermost portion (Fig. 4.7).
The Middle Member consists of similar facies to the Montenegro Member of the
Campito Formation and contains three primary siliciclastic facies: i) laminated siltstone
with thin (< 2 cm) sharp-based planar laminated very fi ne sandstone beds (Fig. 4.8A);
ii) sharp-based thick (3-10 cm) quasi-planar laminated and hummocky cross-stratifi ed
fi ne to very fi ne sandstone beds with thin (1-3 cm) siltstone layers and partings (Fig.
4.8B); and iii) amalgamated quasi-planar laminated fi ne sandstone with rare hummocky
cross-stratifi cation (Fig. 4.8C). These three facies stack to form upward-coarsening and
upward-thickening successions that represent deposition within the proximal offshore,
offshore transition, and lower shoreface, respectively. This is similar to the interpreted
environments of the siliciclastic facies of the uppermost Campito Formation. Thin (< 2
cm) sandstone beds of facies i were likely deposited below storm wave base from deep-
penetrating storm-generated currents as distal tempestites or graded rhythmites (Aigner
and Reineck, 1982). Quasi-planar laminated sandstone beds interbedded with siltstone
are interpreted as the result of combined-fl ow during storm events (Arnott, 1993), with
siltstone being deposited during fair-weather conditions. Amalgamated quasi-planar
laminated sandstones were likely preserved within a lower shoreface environment in
which storm-generated combined-fl ows predominate (Arnott et al., 1995).
106
VC
C
M
F
VF
Silt
Shale
Lithology
Taphrhelminthopsis
Conichnus
Skolithos
Wrinkle structures
Flooding surface
Proximal
offshore
Offshore
transition
Lower
shoreface
Offshore
transition
Lower
shoreface
Tidal flat
Proximal
offshore
Offshore
transition
Lower
shoreface
FS
FS
0
1
2
3
4
5
6
7
8
9
10
11
12
13
14
FS
15
16
17
18
19
20
21
22
meters
Limestone with siltstone flasers
Siltstone and quasi-planar laminated sandstone
Quasi-planar laminated and hummocky
cross-stratified sandstone
Planar laminated siltstone
Middle Member Poleta
FS
Erosional unconformity
FIGURE 4.7. Partial stratigraphic section measured through the upper portion of the
Middle Member of the Poleta Formation within the Poleta folds region and interpreted
depositional settings. Base of section begins at N 37° 18.877 ′, W 118° 05.444 ′. FS =
fl ooding surface.
107
5 cm
1 cm
A B
C
D
FIGURE 4.8. Outcrop photographs of facies and wrinkle structures from the Middle
Member of the Poleta Formation at the Poleta folds locality. A) Massively bedded
siltstone with thin (1-2 cm thick) planar laminated very fi ne sandstone layers. B)
Interbedded planar laminated siltstone and quasi-planar laminated very fi ne sandstone;
offshore transition. C) Amalgamated quasi-planar laminated fi ne sandstone with low-
angle truncations; lower shoreface. D) Bedding surface view of wrinkle structures found
on a quasi-planar laminated very fi ne sandstone bed that is capped by siltstone; offshore
transition.
108
Wrinkle structures are found atop a quasi-planar laminated very fi ne sandstone
bed, beneath less resistant siltstone, within facies ii of the Middle Member, representing
deposition within the offshore transition (Fig. 4.8D). The upward-coarsening siliciclastic
succession of the Middle Member can be overlain by bedded limestone with siltstone
fl asers and wavy bedding, interpreted as a carbonate tidal fl at (e.g., Demicco, 1983).
These limestones rest unconformably atop lower shoreface sandstones, and occur across
an erosional unconformity. Rare trace fossils are present throughout the section and
are found exclusively within hummocky cross-stratifi ed and quasi-planar laminated
sandstones of the lower shoreface. Ichnogenera include Conichnus, Skolithos, and
Taphrhelminthopsis (e.g., Hagadorn et al., 2000).
White-Inyo Succession - Harkless Formation
The Harkless Formation belongs to Cambrian Stage 4 of Series 2 and occurs
above the fi rst occurrence of the trilobite genus Olenellus, but beneath the fi rst occurrence
of the trilobite genus Oryctocephalus (Nelson, 1978). The Harkless Formation consists
of shale, siltstone, and quartz sandstone with interbedded archaeocyathan limestones
near its base and top (Nelson, 1962). A partial section through the lower portion of
the Harkless Formation was measured just north of Cedar Flat, within the White-Inyo
Mountains (Fig. 4.9).
The Harkless Formation north of Cedar fl at consists of primarily two facies
that stack to form upward-coarsening and upward-thickening successions. The lower
portion of these successions consists of hummocky cross-stratifi ed and low-angle cross-
stratifi ed fi ne sandstones interbedded with sandy siltstone (Fig. 4.10A). Sandstone beds
are sharp-based and thicken upward within the facies from 3 to 20 cm thick. Overlying
the interbedded siltstone and sandstone facies is found amalgamated hummocky cross-
109
VC
C
M
F
VF
Silt
Shale
Lithology
Wrinkle structures
Archaeocyath limestone
Non-amalgamated hummocky
cross-stratified sandstone
Amalgamated hummocky
cross-stratified sandstone
Planar laminated siltstone
0
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
meters
Harkless Formation
Offshore
transition
Lower
shoreface
FIGURE 4.9. Partial stratigraphic section measured through the Harkless Formation at
Cedar Flat and interpreted depositional settings. Base of section begins at N 37° 17.111 ′,
W 118° 08.035 ′. FS = fl ooding surface.
110
AB
C D
5 cm
FIGURE 4.10. Outcrop photographs of facies and wrinkle structures from the Harkless
Formation, Cedar Flat locality. A) Non-amalgamated hummocky cross-stratifi ed very
fi ne sandstone interbedded with less resistant planar laminated siltstone; offshore
transition. Staff scale is in decimeters. B) Amalgamated hummocky cross-stratifi ed fi ne
sandstone; lower shoreface. Hammer is 24 cm long. C) Isolated lens (dashed outline) of
archaeocyath packstone within amalgamated hummocky cross-stratifi ed fi ne sandstone;
lower shoreface. D) Bedding surface view of wrinkle structures found on a hummocky
cross-stratifi ed very fi ne sandstone bed that is capped by less resistant siltstone.
111
stratifi ed fi ne to very fi ne sandstone with 10-40 cm thick archaeocyathan packstone and
grainstone layers and lenses that are low-angle cross-stratifi ed to massively bedded (Fig.
4.10B-C).
The Harkless Formation facies represent deposition within the offshore transition
and lower shoreface. Upward-thickening hummocky cross-stratifi ed sandstones
interbedded with siltstone represent distal to proximal sandy tempestites of the offshore
transition (e.g., Aigner and Reineck, 1982), while the amalgamated hummocky cross-
stratifi ed sandstones represent storm wave-dominated deposition within the lower
shoreface (Harms et al., 1975). Wrinkle structures are found exclusively within offshore
transition deposits and occur atop sandstone bedding surfaces that are exposed beneath
less resistant siltstone (Fig. 10D). Layers and lenses of archaeocyath packstone and
grainstone may represent local skeletal banks or build-ups, similar to the bioherms
documented in the lower portion of the Harkless Formation by McKee and Gangloff
(1969).
Paleoenvironmental Distribution of Wrinkle Structures
Throughout the Death Valley and White-Inyo regions, lower Cambrian wrinkle
structures are found preserved atop quartz sandstone beds within heterolithic strata of
interbedded sandstone and siltstone. In the Death Valley region, the middle member
of the Wood Canyon Formation preserves wrinkle structures on planar laminated and
low-angle cross-laminated fi ne to very fi ne sandstone beds of lower to mid tidal fl at
environments. In the White-Inyo region, the Montenegro Member of the Campito
Formation preserves wrinkle structures on a hummocky cross-stratifi ed very fi ne-grained
proximal sandy storm deposit of the offshore transition, lying seaward of a siliciclastic
shoreface and emerging archaeocyathan reef complex. The Middle Member of the Poleta
Formation preserves wrinkle structures on a very fi ne-grained quasi-planar laminated
112
sandy tempestite of the offshore transition, seaward of a siliciclastic shoreface and
back-barrier carbonate tidal fl at system. The Harkless Formation preserves wrinkle
structures on numerous hummocky cross-stratifi ed very fi ne-grained proximal sandy
tempestites of the offshore transition, lying seaward of a siliciclastic shoreface with local
archaeocyathan skeletal banks or low-relief build-ups. In all instances, wrinkle structures
are preserved at the interface between underlying sandstone and overlying siltstone.
This pattern of preservation within heterolithic strata is the most common type found for
wrinkle structures in both the Precambrian and Phanerozoic (e.g., Noffke et al., 2002;
Mata and Bottjer, 2009b).
The preference for the preservation of wrinkle structures at sand-silt interfaces is
likely due to multiple factors. Noffke et al. (2002) suggested that phototrophic microbial
mats might show a preference for the colonization of quartz sand due to its high
translucence, as compared to other sediment types. This type of sediment is, however,
most typically associated with high-energy environments and mobile substrates that are
not conducive to microbial mat formation (Mata and Bottjer, 2009b). For this reason, it
has been suggested by Mata and Bottjer (2009b) that most wrinkle structures throughout
Earth’s history actually formed on sandy tempestites deposited below fair-weather
wave base. These tempestites would allow for a requisite translucent substrate, yet still
offer low-energy conditions conducive to microbial mat formation and preservation.
Fine-grained deposition of silt and mud during fair-weather conditions would allow
for preservation without erosion (e.g., Noffke et al., 2002). A similar process might be
expected within tidally infl uenced environments in which high-energy tidal currents
promote clean sandy substrates for colonization, while preservation of the microbial mat
would be favored during slack-water conditions where silt and mud could blanket the
substrate. This qualitative assessment can be further expanded into a semi-quantitative
assessment utilizing threshold entrainment velocities for modern microbial mats as
113
analogue fl ow parameters for the controls on microbial mat development and preservation
in ancient environments.
Hydrodynamic Limits on Microbial Mat Formation and Preservation
Layered microbial mats are capable of enmeshing underlying sediment and can
produce a biostabilized cohesive substrate that is more erosion-resistant than unbound
sediment alone (e.g., Ginsburg and Lowenstam, 1958; Bathurst, 1967; Neumann et al.,
1970; Noffke et al., 2001; Cady and Noffke, 2009; Hagadorn and McDowell, 2012). This
is because increased binding by microbes and extracellular polymeric substances tends
to lead to increased threshold entrainment velocities (Dade et al., 1990). Stemming from
these observations grew a suite of experiments seeking to redefi ne and quantify threshold
velocities for the entrainment of microbial mat-bound sediment (e.g., Neumann et al.,
1970; Cady and Noffke, 2009; Hagadorn and McDowell, 2012), similar to those defi ned
previously for sterile unbound sediment (e.g., Miller et al., 1977; Hammond and Collins,
1979).
Early pioneering work was conducted by Neumann et al. (1970), which examined
the entrainment threshold velocity of tidally-infl uenced mat-bound fi ne to medium
carbonate sand of the Little Bahama Bank. Utilizing an in situ subaqueous fl ume, it
was shown that mat-bound sediment required at least twice the unidirectional current
velocity of unbound sediment in order to initiate entrainment. Unbound sediment within
the studied areas had a threshold entrainment velocity of approximately 20 cm/s, while
microbially-bound sediment required velocities of approximately 30-40 cm/s to initiate
entrainment and sometimes up to 110-120 cm/s dependent upon the nature of the initial
mat conditions. Also, originally unbroken mats could withstand much higher velocities
than those that experienced previous breakage.
114
This pioneering work by Neumann et al. (1970) has been followed up and
reproduced by subsequent experiments. Cady and Noffke (2009) through a similar in
situ fl ume experiment on the tidal fl ats of Portsmouth Island, North Carolina, showed
that endobenthic mats are capable of resisting up to 90 cm/s fl ow velocities before
entrainment, while thicker epibenthic mats could resist up to 160 cm/s fl ow velocities.
Hagadorn and McDowell (2012) examined the threshold entrainment velocity
and bedform stability of cultured microbial mat layers of varying growth stages (10-
48 days) within a laboratory fl ume containing medium-grained quartz sand. In young
mats, growing for 10 days, threshold entrainment velocities ranged from 25-28 cm/s. In
thicker moderate-aged mats, growing for 12-35 days, entrainment began around 25-32
cm/s. In both cases, the fi rst type of mat failure with increasing velocity is mat fl ip-overs
in which the mat is undercut before fl ipping over itself. In the oldest mats (42-49 days),
entrainment of the mat occurred at values similar to those for the younger 10-day mats;
however, this may have been facilitated by the development of gas blisters that fostered
mat buoyancy.
While these studies have focused primarily on the erosion of suffi ciently
developed microbial mats that grew at lower hydrodynamic conditions, it is unclear what
fl ow conditions are necessary to inhibit the initiation of microbial mat development.
Presumably, velocities equal to those required to erode a microbial mat should be
suffi cient to inhibit initial mat formation, and may even be less to inhibit incipient forms
similar to the younger 10-day mats of Hagadorn and McDowell (2012). Also, while
these studies utilizing unidirectional fl ow are widely applicable for tidal fl at environments
in which this type of fl ow predominates, they are not directly applicable for examining
the notable exclusion of microbial mat features from modern and ancient shoreface
environments, which are dominated by waves and oscillatory fl ow.
115
Oscillatory and unidirectional currents do not necessarily exhibit identical
threshold velocities for a given grain size (Hammond and Collins, 1979). In spite of this,
there are some notable qualitative relationships between oscillatory and unidirectional
fl ow conditions that can aid in translating unidirectional fl ow observations on microbial
mats into interpretations about the effects of oscillatory fl ow. Equal orbital and
unidirectional velocities can have comparable entrainment threshold values for a wide
range of grain sizes from fi ne to very coarse sand, although unidirectional fl ows tend to
have slightly lower thresholds (5-10 cm/s less) for a given grain size (e.g., Hammond
and Collins, 1979; Paphitis et al., 2001). This has been determined experimentally for
simulated oscillatory conditions with wave periods between 5 and 12 seconds, with
decreasing entrainment threshold velocities with decreasing wave period (Paphitis et al.,
2001). Subsequent discussion will treat unidirectional and oscillatory fl ows as roughly
comparable, as the difference between their threshold entrainment velocities for a given
grain size are likely well within the natural variability of a given environment (i.e., +/-
5-10 cm/s)
Shoreface and Nearshore Dynamics
The shoreface consists of the high-energy, constantly wave-reworked zone that
lies between mean low tide and fair-weather wave base—the depth to which fair-weather
waves can move sedimentary particles. The morphology of the shoreface is generally
concave-upward and steepest at the top, gently sloping seaward before fl attening out
where it merges with the innermost portion of the shelf. The shoreface is subject to both
fair-weather and storm processes; however, they operate over different timescales and
exert different magnitudes of infl uence.
When examining ancient depositional conditions it is apparent that the product
of storms (i.e., storm deposits) can predominate within a sedimentary succession,
116
yet the amount of time represented by these storms is very minimal compared to the
time represented by the fair-weather deposits between them (Dott, 1983). It is likely,
therefore, that fair-weather conditions are those that should be considered when
examining the potential for formation of a microbial mat within a particular environment
because the growth of the mat is on the same timescale of the fair-weather conditions
between storms. A microbial mat, however, is more likely to be eroded and destroyed
during higher-energy storm conditions because a growing microbial mat should be in
equilibrium with fair-weather processes and unlikely to be eroded by them.
Oscillatory fl ow velocities within shoreface settings (< 15 meters depth)—
which are contingent upon wave height and water depth for solitary waves—average
10-30 cm/s under fair-weather conditions (e.g., Cook and Gorsline, 1972; Hill et al.,
2003). Commonly, a weak cross-shore or along-shore unidirectional current (< 10
cm/s) can be superimposed over this oscillatory current (e.g., Hill et al., 2003). Under
storm conditions, oscillatory velocities can far exceed those present under fair-weather
conditions, and mean velocities for a given storm can often approach or exceed 100 cm/s
(e.g., Hequette and Hill, 1993; Amos et al., 1996; Hill et al., 2003) and can reach values
of almost 200 cm/s (Amos et al., 1996). These oscillatory currents are often coupled with
offshore-directed unidirectional storm-generated currents that have velocities that closely
co-vary with nearshore storm surge levels (Hequette and Hill, 1993). Unidirectional
near-bottom currents associated with downwelling during these storm surges can range
from 10-160 cm/s (e.g., Murray, 1970; Leckie and Krystinik, 1989; Hequette and Hill,
1993; Hill et al., 2003). A wide range of oscillatory, unidirectional, or combined fl ows
can be present during both fair-weather or storm conditions, and no particular type of
fl ow is specifi c to a given shoreface.
While oscillatory fl ow velocities within the shoreface under fair-weather
conditions are close to the velocities required to initiate erosion on a young microbial mat
117
(e.g., Hagadorn and McDowell, 2012)—assuming that oscillatory fl ow velocity roughly
tracks the threshold velocities of unidirectional currents—there is indeed potential for
microbial mat development if the shoreface is situated on a low-energy coastline. This
would require the presence of waves on the low end of the spectrum for fair-weather
fl ow velocities. Most moderate to high-energy storms, however, could certainly provide
both oscillatory and unidirectional fl ow conditions with velocities high enough to erode
an incipient mat, but might not be strong enough to completely erode a well-developed
mature mat similar to the epibenthic mats documented on Portsmouth Island by Cady and
Noffke (2009) that could withstand unidirectional velocities of up to 160 cm/s.
Storm wave-dominated shoreface successions typically consist of trough cross-
stratifi ed sandstones with abundant scours in the upper shoreface and amalgamated
hummocky cross-stratifi ed sandy tempestites in the lower shoreface (e.g., Harms et al.,
1975; MacEachern and Pemberton, 1992). This distribution exists because the upper
shoreface is typically erosional and sources sandy sediment for the lower shoreface
and shallow shelf during storms. The amalgamated nature of storm-dominated lower
shoreface successions consisting exclusively of hummocky cross-stratifi ed storm deposits
refl ects the total absence of preserved fair-weather deposits. Each successive tempestite
incises into the one below it, eroding away the intervening fair-weather deposit. If a
microbial mat managed to form within the shoreface under fair-weather conditions, its
potential for preservation within a storm-dominated setting would be very low because
of tempestite amalgamation. It is only within environments below fair-weather wave
base in which suffi cient fi ne-grained sediment can drape a tempestite during fair-weather
conditions to inhibit amalgamation and promote microbial mat preservation.
Heterolithic strata similar to those of the offshore transition can also be found in
tidal fl at environments. In these settings high-energy tidal currents alternate with periods
of slack-water, resulting in interbedded sand, silt, and mud (e.g., Reineck, 1967). Tidal
118
currents are highly variable and differ between environments. Tidal currents fl owing
across tidal fl ats can average 5-40 cm/s (Reineck, 1967; Wells, 1990; Widdows et al.,
2004), while channelized fl ow within inlets and channels can average up to 120-130
cm/s (Kumar and Sanders, 1974; Reddering, 1983; Fenster and FitzGerald, 1996). It
appears, however, that most reported occurrences of wrinkle structures from ancient
tidally infl uenced environments are found in tidal fl at deposits, rather than channels (e.g.,
Wunderlich, 1970; Fedo and Cooper, 1990; Mángano and Buatois, 2004; Noffke et al.,
2006). These wrinkle structures likely formed on the low end of the spectrum for tidal
currents (i.e., 5-40 cm/s). Much like in this study, wrinkle structures in these settings
are preserved on sandstone beds that are interbedded with siltstone or mudstone (e.g.,
Wunderlich, 1970; Fedo and Cooper, 1990). A similar preservation process to shelf strata
occurs within these environments. Microbial mats likely formed on quartz-rich substrates
that were eventually draped and buried by slack-water fi ner-grained sediment, resulting in
microbial mat preservation.
Facies Model for Wrinkle Structure Development
Wrinkle structures show a preference for heterolithic strata of the offshore
transition and mixed tidal fl ats in this study and in previous studies (e.g., Wunderlich,
1970; Fedo and Cooper, 1990; Banerjee and Jeevankumar, 2005; Noffke et al., 2002;
Mata and Bottjer, 2009a; 2009b). While the absence of preserved wrinkle structures in
the high-energy shoreface is not surprising, the absence of preserved wrinkle structures in
environments below storm wave base is more problematic. Many types of microbial mat-
forming organisms can exist under low light conditions or no light conditions and utilize
either phototrophy or chemotrophy, respectively. The absence of wrinkle structures from
ancient deep-water environments should therefore likely not be due to an absence of mat-
forming organisms or an absence of mats. Rather, the absence may be taphonomic.
119
The preference for wrinkles structures to be found on quartz sandstone bedding
planes may be due to the benefi ts of a translucent substrate for phototrophic mats (e.g.,
Noffke et al., 2002), or it may simply be that fi ner-grained sediments, consisting of
silt and mud, tend not to develop extensive bedding planes (Mata and Bottjer, 2009b).
Therefore, the lack of wrinkle structures within mud-rich shales may not necessarily be
due to a lack of former microbial mats, but rather a lack of adequate preservation to allow
for their later observation. In fact, many microbial mat features have been documented
in fi ner-grained sediments, including microfault sets and millimeter ripples in siltstones
(Pfl üger, 1999) and wavy-crinkly carbonaceous laminae in shales (Schieber, 1999), yet
not wrinkle structures. Wrinkle structures appear to be restricted primarily to sandstones.
An additional possibility is that the microbes responsible for making wrinkle structures
are exclusively phototrophic and must remain in the euphotic zone. If this were the case,
it might be that the base of the euphotic zone and storm wave base occur at comparable
depths (Mata and Bottjer, 2009b), with both commonly ranging from approximately 40-
60 meters in depth (e.g., Saito, 1989; Nelson, 1999; Kuwahara et al., 2000). It is notable,
however, that some wrinkle structures have been reported from deep sea environments,
which may have been well below the euphotic depth (e.g., Buatois and Mángano, 2003).
If the requirements for wrinkle structure preservation are a quartz-rich sandy
substrate for formation and casting, and a fi ner-grained drape of silt or mud for
preservation, then this restricts the types of environments in which wrinkle structures
would be predicted to occur (Fig. 2.11). Environments below storm wave base in which
the deposition of sand is rare should be expected to be generally lacking in wrinkle
structures in spite of a prevalence of contemporaneous microbial mat features in adjacent
environments. It may also be that microbial mats were present in these environments,
but not preserved. Exceptions to this include the presence of wrinkle structures on lower
Cambrian deep sea sandy turbidite deposits (Buatois and Mángano, 2003). Heterolithic
120
Siltstone and hummmocky
cross-stratified sandstone
Planar laminated or massive
mudstone and siltstone
Hummocky cross-stratified
sandstone
Fair-weather wave base
Storm wave base
Planar laminated sandstone
Percent Sand
Percent sea surface PAR at seafloor
Average hydrodynamic energy conditions
Low
High
0%
100%
0%
100%
Lagoon
Upper
shoreface
Lower
shoreface
Offshore
transition
Offshore
Tidal
flat
Foreshore/
washover
fan
Trough cross-stratified
sandstone
Massive or mottled
mudstone
Flaser, wavy, and lenticular
bedded sandstone and siltstone
Lower Cambrian wrinkle structure distribution
FIGURE 4.11. Facies model showing the preferred environments for wrinkle structure
formation and preservation (marked by dashed lines) based upon the parameters
of average hydrodynamic energy conditions, percent sand, and percent sea surface
photosynthetically active radiation (PAR) at the seafl oor (modifi ed from Mata and Bottjer
2009b). If most wrinkle structures are formed by phototrophic microbial mats, this would
limit most wrinkle structures to above the euphotic depth, which is approximated by 1%
PAR. Wrinkle structures show a preference for being preserved atop quartz sandstones,
and this would limit wrinkle structure preservation to environments where sand can be
delivered by storm waves, currents, and tidal activity. This would limit wrinkle structures
to environments primarily above storm wave base. Also, wrinkle structures should be
prohibited from high-energy environments like the shoreface, because fl ow velocities
are strong enough to erode an incipient mat. With these restrictions, this leaves the
offshore transition, lagoons, and tidal fl ats as the primary environments in which wrinkle
structures should be preferentially preserved. Lower Cambrian wrinkle structures are
found primarily within the proximal portion of the offshore transition and within mid
tidal fl at settings.
121
facies of the offshore transition, between fair-weather wave base and storm wave base,
would be expected to preserve wrinkle structures if microbial mats were present, but only
if colonization occurred on sandstone beds. The high-energy shoreface may allow for
periods of microbial mat development, but storm deposition might preclude preservation
due to erosion and frequent amalgamation of tempestites. Mixed tidal fl ats are also
expected to preserve wrinkle structures if microbial mats were present, but the highest
potential would likely be in mid tidal fl ats. The lowermost portion of a tidal fl at typically
contains facies very similar to the shoreface and consists of amalgamated bidirectional
trough and tabular cross-stratifi ed sandstones, with only minimal deposition of fi ner-
grained sediment. The uppermost portion of a tidal fl at tends to be muddy and lacks
suffi cient sand to allow for the casting and preservation of wrinkle structures, even if
microbial mats were present. Modern lagoons are very hospitable environments for the
development of microbial mats, yet their record of siliciclastic microbial mat deposits
is all but lacking, except for rare occurrences (e.g., Noffke, 2000). In spite of this,
heterolithic lagoonal deposits, in which sand is delivered to the back-barrier environment
through storm washover, might be expected to record a more faithful signal of microbial
mat development and wrinkle structure preservation than mud-rich lagoons in which
storm washover is less signifi cant. More data on ancient lagoonal systems is required,
however, to explore this in more detail.
Implications for Early Cambrian Shallow Marine Siliciclastic Substrates
The Precambrian-Cambrian transition saw a marked increase in trace fossil
diversity and complexity that began in the Ediacaran and increased gradually through the
early Cambrian (Crimes, 1992; MacNaughton and Narbonne, 1999). This increase in
trace fossil diversity was accompanied by an increase in extent and depth of bioturbation
in carbonate environments at the end of the Terreneuvian series (Droser and Bottjer,
122
1989). The nature of the early Cambrian trace fossil record is notably different depending
upon which environments are considered, so it is important to place the discussion of the
early Cambrian substrate within a paleoenvironmental context.
Nearshore environments, primarily above fair-weather wave base, were colonized
by deep-penetrating vertical burrowers by the end of the early Cambrian at the latest
(e.g., Hiscott et al., 1984; Droser, 1991; Prave, 1992; McIlory and Garton, 2004). These
burrowers belong primarily to the Skolithos pipe-rock biotope (Droser, 1991; McIlroy
and Garton, 2004), consisting of the trace fossil Skolithos with lesser Monocraterion.
Skolithos are known to penetrate to an average depth of 10-15 cm (Droser, 1991; Miller
and Byers, 1984). In addition to the simple vertical tubes of Skolithos, there are many
additional deep-penetrating trace fossils found within early Cambrian nearshore settings.
The anemone trace fossils Conichnus and Dolopichnus are known to have penetrated
to depths of 19 cm and 27 cm, respectively, within tidal inlet complex deposits of
the lower Cambrian L’Anse-au-Clair Member of the Bradore Formation in Labrador,
Canada (Hiscott et al., 1984). In these same deposits, the spade-shaped inarticulate
brachiopod trace fossil Lingulichnus is known to have penetrated to nearly 10 cm in
depth (Pemberton and Kobluk, 1978). Anemone trace fossils similar to those from the
Brador Formation are found within the upper member of the Wood Canyon Formation in
the Death Valley region, USA, and Conichnus and Dolopichnus penetrate to depths of 29
cm and 19 cm, respectively (Mata et al., in review).
Early Cambrian bioturbation within muddy shelf environments, below fair-
weather wave base, shows a much more restricted depth of burrowing. Droser et al.
(2002) show that lowermost Cambrian fi ne-grained shelf sediments exhibit evidence for
fi rmground conditions at depths close to the sediment-water interface. These fi rmgrounds
are marked by exceptional preservation of surface markings and scratch marks that
would not be preserved if there was a surface mixed layer and a softground seafl oor.
123
While fi rmgrounds can form due to exhumation of buried and compacted sediment along
specifi c stratigraphic surfaces, these early Cambrian fi rmground conditions appear to
be ubiquitous in many shelf settings and are a refl ection of low levels of bioturbation
(Droser et al., 2002; 2004). While depth of burrowing within nearshore settings can
approach 30 cm or more, bioturbation on the shelf appears to be restricted to less than 10
cm in depth, and more commonly less than 5 cm (McIlroy and Logan, 1999; Droser et al.
2004; Marenco and Bottjer, 2008). These early Cambrian shelf settings appear to be the
primary depositional environment in which wrinkle structures are preserved.
In addition to low levels of bioturbation, Cambrian shelf settings are notable
for containing abundant evidence for seafl oor microbial mats in the form of wrinkle
structures (e.g., Hagadorn and Bottjer, 1997; 1999; Bailey et al., 2006; Mata and Bottjer,
2009b; Buatois and Mángano, 2012). In some instances these wrinkle structures
can be associated with horizontal trace fossils (Hagadorn and Bottjer, 1999; Buatois
and Mángano, 2012) or with body fossils, including inarticulate brachiopods and the
enigmatic agglutinated fossil Volborthella (Bailey et al., 2006). The presence of these
microbial mat features indicates the lack of a surface mixed-layer and an absence of
extensively developed bioturbation in early Cambrian siliciclastic shelf settings. Even in
deeper shelf environments where microbial mat features do not appear to be preserved,
there is strong evidence for fi rmground conditions. The early Cambrian helicoplacoid
echinoderms lived as suspension-feeding sediment stickers, an ecology that had its
roots in the matground ecosystems of the Ediacaran, but became scarce following
the Cambrian (Bottjer et al., 2000; Dornbos and Bottjer, 2000; 2001). The sediment
sticker ecology is contingent upon the presence of a fi rm substrate and the absence
of a mixed layer, and most helicoplacoids are associated with laminated shales with
limited to no bioturbation (Dornbos and Bottjer, 2000; 2001). These helicoplacoids
are found primarily within the Middle Member of the Poleta Formation, and inhabited
124
an environment further out on the shelf than the heterolithic wrinkle structure-bearing
offshore transition environments documented in this study from the Middle Member.
The persistence of wrinkle structures in siliciclastic shelf settings throughout the
duration of the early Cambrian in the southern Great Basin suggests that a signifi cant
mixed-layer had yet to develop within these settings. This is further reinforced by
previous studies examining the extent and depth of bioturbation within lower Cambrian
siliciclastic shelf strata. Dornbos and Bottjer (2000) showed that outer shelf deposits of
the Middle Member of the Poleta Formation consist primarily of bedding-parallel trace
fossils or are not bioturbated at all. Marenco and Bottjer (2008) quantifi ed the extent
and depth of bioturbation within siliciclastic shelf strata of the Campito, Poleta, and
Harkless formations to show that the primary style of bioturbation present is that of the
simple horizontal burrow Planolites. These burrows did not extend much below the
sediment-water interface and their predominant horizontal orientation suggests that they
may have served as undermat miners during this earliest Cambrian interval (Marenco and
Bottjer, 2008). The nature of the shelf environment stands in stark contrast to nearshore
environments that record deep-penetrating bioturbation in the early Cambrian. While it
might be expected that this deep-penetrating bioturbation is responsible for the absence
of wrinkle structures from these nearshore settings, a more likely interpretation is that
high hydrodynamic energy conditions were detrimental enough to inhibit microbial mat
formation or possibly preservation.
Conclusions
The lower Cambrian record of the southern Great Basin preserves ancient
microbial mats in the form of wrinkle structures. A detailed facies analysis reveals that
wrinkle structures show a preference for being cast on underlying sandstone beds and
draped with overlying siltstone. In terms of depositional environment, wrinkle structures
125
show a preference for heterolithic strata of mixed tidal fl ats and the offshore transition—
between fair-weather wave base and storm wave base. Mixed tidal fl at deposits consist of
fl aser bedded, wavy laminated, and planar laminated sandstone and siltstone with wrinkle
structures being preserved exclusively on planar laminated sandstone beds. Offshore
transition deposits consist of interbedded siltstone and hummocky cross-stratifi ed
or quasi-planar laminated sandy tempestites with wrinkle structures being preserved
exclusively on sandstone beds.
The preference for wrinkle structures within these environments may be due to
several factors. If wrinkle structures were made exclusively by phototrophic microbial
mats, then this might explain their absence from most deep-water environments. The
ubiquitous preservation of wrinkle structures at the interface between underlying
sandstone and overlying siltstone or shale suggests that these two contrasting lithologies
are possibly an essential element to taphonomic preservation. If this is the case, then it
would be expected that wrinkle structures should be most prevalent in environments that
consist of heterolithic deposits, and this is the pattern revealed by this study. In all strata
examined, wrinkle structures are only found in heterolithic strata and are absent from
deposits consisting of only siltstone or only sandstone.
The preservation of wrinkle structures within heterolithic strata can be explained
by the observation that the hydrodynamic energy requirements for microbial mat
formation and growth are different from those needed for preservation. Ideally a
phototrophic mat would be best suited to a clean sandy substrate with minimal mud or
clay content to inhibit sunlight penetration. Previous studies also suggest that sandstone
may also be required for the casting of wrinkle structures (e.g., Mata and Bottjer, 2009b).
Achieving such substrate conditions requires moderate-energy, but not necessarily
high-energy conditions. The preservation of a microbial mat, in contrast, appears to
be restricted to instances of low-energy conditions. All wrinkle structures examined in
126
this study were all draped by the deposition of silt below fair-weather wave base or the
deposition of silt during slack-water conditions on a mixed tidal fl at. The persistence of
either high-energy conditions or low-energy conditions within an environment would be
detrimental to the ultimate preservation of a microbial mat.
Modern studies on the threshold entrainment velocity of microbial mats
suggest that only low-energy shorefaces may be capable of allowing for microbial mat
formation; however, storm deposition across any shoreface would most likely inhibit
preservation. This is because most proximal storm deposits within the shoreface tend to
be amalgamated and result in erosion prior to deposition. Therefore it would be unlikely
to preserve wrinkle structures on the soles of amalgamated proximal sandy tempestites
or on underlying deposits. Tempestites distal to the shoreface would have a much greater
preservation potential for wrinkle structures because they are non-amalgamated and are
separated by periods of fair-weather silt deposition. A similar phenomenon is expected
for the lower reaches of mixed tidal fl ats where tidal currents can amalgamate subtidal
and intertidal sand bodies, inhibiting wrinkle structure preservation. Environments
higher on the tidal fl at, however, experience less amalgamation and more slack-
water deposition of mud and silt. These environments would therefore have a higher
preservation potential for wrinkle structures.
This study shows that wrinkle structures do indeed show a facies preference
for heterolithic strata, a conclusion reached by previous authors (e.g., Noffke et al.,
2002; Mata and Bottjer, 2009b). The persistence of these microbial features in shelf
settings throughout the duration of the early Cambrian corroborates previous studies that
have documented a widespread development of fi rmground conditions in fi ne-grained
siliciclastic shelf strata and the predominance of horizontal styles of bioturbation.
Taking a facies and paleoenvironmental approach is essential when examining the
paleoenvironmental distribution of microbial mat features, and especially important
127
when documenting the earliest Cambrian trace fossil record, which shows a wide range
of environmental disparity. By understanding the physical controls on microbial mat
distribution, it is possible to better reconcile the effects that the emergence of vertical
bioturbation had on their paleoenvironmental distribution throughout the remainder of the
Phanerozoic.
128
CHAPTER V
Early Cambrian Microbial Mat-Associated Sediment Mounds and Funnels from the
Middle Member of the Wood Canyon Formation, Southern Great Basin, United States
Introduction
The Precambrian-Cambrian transition is marked by increases in extent and
depth of bioturbation that began in the early Cambrian, but reached typical early-
mid Paleozoic levels by the Late Ordovician (Bottjer and Ausich, 1986; Droser and
Bottjer, 1989). During this transition there is a notable diversifi cation of trace fossil
morphologies (Crimes, 1992), even though overall levels of bioturbation remained
low (Crimes and Droser, 1992). These low levels of bioturbation during the earliest
Cambrian allowed for very unique environmental conditions in which radiating metazoan
life coexisted with extensively developed microbial mat ecosystems, and may have
even been morphologically adapted to them (Bottjer et al., 2000). The same appears
true for trace fossil behaviors, and there are indications that earliest Cambrian deep sea
organisms behaved as microbial mat grazers and undermat miners, yet did not destroy the
underlying mat in the process (Buatois and Mángano, 2003; Seilacher et al., 2005).
Microbial mat-associated trace fossils require unique conditions for preservation,
and are limited by the environmental and temporal distribution of microbial mats. In
modern environments, the presence of surface microbial mats is evident and conspicuous,
while in the rock record microbial mats must be largely inferred based upon the presence
of unique sedimentary structures that refl ect the response of microbially-bound sediment
to physical and biological processes (e.g., Hagadorn and Bottjer, 1997). One of the most
common forms is termed ‘wrinkle structures’, which consists of mm-scale meandering
ridges interspersed with pits and sinuous troughs (Hagadorn and Bottjer, 1997). Utilizing
these microbially mediated sedimentary structures, it is possible to infer past microbial
129
mat-associated behaviors based upon their close association with body fossils and trace
fossils (e.g., Seilacher, 1999; Bottjer et al., 2000; Buatois and Mángano, 2003; Seilacher
et al., 2005).
The lower Cambrian middle member of the Wood Canyon Formation in the
southern Salt Spring Hills (Fig. 1.7) preserves vertical tubes associated with wrinkle
structures across an extensively exposed bedding plane surface. Surrounding these
vertical tubes are relatively fl at sediment mounds and funnels that are cm-scale to dm-
scale in diameter and typically contain less than a centimeter of vertical relief. Similar
types of features are known to form from physical processes (e.g., water escape),
microbial processes (e.g., gas escape), or biological processes (e.g., fecal mounds).
The purpose of this study is to interpret the depositional setting of these microbial-mat
associated funnels and sediment mounds, explore what processes led to their formation
and preservation, and what bearing this might have on the Cambrian radiation.

Stratigraphy and Depositional Setting
The middle Member of the Wood Canyon Formation is lower Cambrian in
age and falls within the Terreneuvian series of the Cambrian (Fig. 5.1) (e.g., Babcock
and Peng, 2007). It occurs above the fi rst occurrence of the Precambrian-Cambrian
boundary-marking trace fossil Treptichnus pedum, but below the fi rst occurrence of
trilobites in the region (e.g., Corsetti and Hagadorn, 2000). The middle member can
be subdivided into a lower ‘conglomerate arkose’ unit and an upper ‘arkose-feldspathic
quartzite’ unit (Diehl, 1979). The lower unit consists of quartz sandstone and pebble
conglomerate, while the upper unit consists of interbedded quartz sandstone and siltstone.
The burrows of this study fall within the upper unit of the middle member.
A partial section was measured through the interval containing the burrows and
consists of upward-fi ning and upward-thinning beds of sandstone interbedded with
130
Siltstone + sandstone
DEATH VALLEY
SUCCESSION
Studied
Interval
W ood Canyon
Zabriskie
Carrara
lower
middle
upper
km
0
1
Shale
Limestone
Dolostone
Johnnie
Stirling
Neoproterozoic Cambrian
Conglomerate
Sandstone
FIGURE 5.1. Generalized regional stratigraphy of the Neoproterozoic-Cambrian
succession in the Death Valley region of the southern Great Basin showing the
stratigraphic interval examined in this study (after Stewart 1982; Corsetti and Hagadorn
2000).
131
siltstone (Fig. 5.2). These upward-fi ning successions have been interpreted previously
as prograding mixed tidal fl at deposits (Klein, 1971; 1975). The lowermost portion of
an upward-fi ning succession consists of amalgamated planar laminated and trough cross-
stratifi ed fi ne sandstone with lesser fl aser bedding, herringbone cross-stratifi cation, and
thin siltstone layers and partings. The middle portion consists of interbedded siltstone
and planar laminated and low-angle cross-stratifi ed fi ne to very fi ne sandstone beds
in roughly equal proportions. The uppermost portion of an upward-fi ning succession
consists of laminated to bedded siltstone with thin (< 5 cm) very fi ne sandstone layers
and lenses. This facies succession represents the progradation from the basal lower tidal
fl at sand-dominated facies, through mid tidal fl at heterolithic facies, and into upper tidal
fl at silt-dominated facies (e.g., Klein, 1971).
Methods
In addition to outcrop observations, hand samples from the middle member of the
Wood Canyon Formation were slabbed and examined as polished slabs, in x-radiographs,
and within thin section. For x-radiographs, slabs were cut to less than 1 cm in thickness
and were x-rayed using a Penetrex industrial x-ray machine. Samples were exposed at
8 mA and 96 kV for 12-36 minutes depending upon thickness. X-radiograph negatives
were digitized using a UMAX PowerLook 2100XL scanner. Since negatives are being
utilized in all examinations, brighter colors represent higher densities and darker colors
represent lower densities in all x-radiographs shown. Also, due to the thickness of the
slabs, x-radiographs represent an averaging of all material penetrated throughout the
width of the slab and can therefore superimpose physical sedimentary structures and
biogenic structures from different slices through the samples, even though they do not
come into contact.
132
Sandy siltstone
Low-angle cross-stratified sandstone
Cross-laminated sandstone with silt flasers
Planar laminated sandstone
FS
VC
C
M
F
VF
Silt
Shale
Lithology
FS
Lower flat
Lower flat
Lower flat
Mid flat
Mid flat
Mid flat
FS
Mid flat
Upper flat
Upper flat
0
1
2
3
4
5
6
7
8
9
10
11
12
13
14
meters
middle member Wood Canyon
Wrinkle structures
Mound or funnel
FIGURE 5.2. Partial stratigraphic section through the middle member of the Wood
Canyon Formation at the Salt Spring Hills showing the distribution of lithologies,
environments, and the location of the funnels and sediment mounds examined.
133
Sediment Mound and Funnel Descriptions
Sediment mounds and funnels within the middle member of the Wood Canyon
Formation are found preserved across a single laterally exposed bedding plane that caps
a bed that is approximately 1 meter thick (Fig. 5.3). The lithology of the bed consists of
planar laminated fi ne sandstone, and sediment mounds and funnels are found only on the
upper surface of the bed. Interspersed between sediment mounds are wrinkle structures,
which preserve the record of ancient microbial mats (Hagadorn and Bottjer, 1997).
Amongst all features found, there are three primary morphologies of mounds and funnels
preserved: 1) sediment mounds found within shallow depressions; 2) isolated funnels;
and 3) isolated sediment mounds.
Sediment mounds found within shallow depressions are the most common and
distinctive feature found on the bedding surface (Fig. 5.4A-C). In bedding plane view, at
the center of each sediment mound is a single vertical tube that ranges in size from 1-1.5
cm in diameter. Surrounding this central tube is an apron or mound of sediment that is
relatively fl at lying, but shows defi nite relief above the surrounding bedding plane. In
some instances, this apron may be absent. The apron can extend from the central tube for
a radius of 1-4 cm and can be highly irregular and not necessarily circular in shape. This
apron occurs as a discrete sediment layer above the host sediment. Outward from this
zone can be found a trough that encircles the apron and extends 1-2 cm from the rim of
the sediment apron. This trough, in many instances, has a negative relief and falls below
the level of the surrounding bedding plane. Outside of the trough is found an outer bulge
of sediment that is roughly circular, relatively fl at-topped, and shows relief above the
surrounding bedding plane. This bulge can have widths of 1-2 cm and its outer portion
appears more regular and circular than the interior portion of the bulge, which can have a
very irregular contact with the trough. Unlike the inner sediment apron, this bulge does
not appear to be a separate layer from the surrounding bedding plane or host sediment.
134
30 cm
ws
ws
ws
ws
ws
ws
FIGURE 5.3. Outcrop photographs of the bedding plane exposure examined. Wrinkle
structures (WS) are distributed in patches across the bedding plane. Sediment mounds
and funnels are found distributed sporadically with no discernible pattern.
135
Isolated funnels also occur on the bedding surface and are the second most
common morphology, following the sediment mounds found within shallow depressions
(Fig. 5.4D-F). Funnels are approximately 2-3 cm in diameter and penetrate to 0.5-1 cm
in depth below the surface of the bedding plane. The surface of the funnels, which can
be very irregular, is pierced by a vertical sediment tube. These funnels occur in close
association with the sediment mounds found within shallow depressions; however, as will
be shown subsequently, there is no apparent subsurface connection between these two
features.
Isolated sediment mounds not found in association with depressions are relatively
rare on the bedding plane, but do occur (Fig. 5.4G). These mounds also have a different
morphology than those associated with shallow depressions. Vertical tubes are 1-1.5 cm
in diameter and the sediment mounds that surround them tend to be steeper and more
irregular. The radius of the sediment mound from the central tube is smaller and typically
only extends 1-2 cm from the edge of the central vertical shaft. Also, the mounds can
have up to 1 cm of vertical relief above the bedding surface. In some instances the
sediment mounds can be very asymmetric.
Due to the thickness and nature of the bedding plane, not all surface morphologies
could be sampled for examination in cross section, and only representative samples of
isolated funnels and isolated sediment mounds could be obtained. In cross section, the
tubes found associated with the isolated funnels tend to penetrate to a typical depth of
2-4 cm, but can reach lengths of up to 10 cm below the bedding surface (Fig. 5.5A-
F). Downward defl ected laminae are the most common feature within the adjacent
host rock and are found primarily near the uppermost portion of the funnel (Fig. 5.5A).
Disturbed zones, in the form of irregular margins and diffuse boundaries resembling
mottling, can be found along the outer periphery of the tube walls, but do not extend
more than a centimeter from the tube (Fig. 5.5B-C). In rare instances, funnels connect
136
A B C
D
E F G
outer bulge
trough
inner bulge
vertical
tube
outer bulge
inner mound
vertical
tube
mound
mound
funnel
funnel
funnel
funnel
wrinkle
structures
wrinkle
structures
wrinkle
structures
2 cm 2 cm
5 cm
2 cm 2 cm 1 cm
vertical
tube
funnel
FIGURE 5.4. Outcrop photographs of the bedding plane expression of sediment mounds
and funnels. A) Sediment mound pierced by a vertical tube found within a shallow
depression rimmed by a low relief bulge. The mound and bulge are separated by a trough.
B) Sediment mound and vertical tube abutting the lower relief bulge. C) Sediment
mounds and funnels found in close association on a bedding plane. D) Funnels found
in associated with patchy wrinkle structures. E) Close-up of funnel in association with
wrinkle structures. F) Two small funnels found in close association. G) Isolated sediment
mound with a vertical tube piercing it.
137
in the subsurface, forming u-shaped tubes (Fig. 5.5D). The fi ll of most tubes associated
with surface funnels is massive; however, some funnels contain laminated sediment or
are fi lled with fossil debris (Fig. 5.5C-F). Fossils found in some funnels consist of the
agglutinated cones of the lower Cambrian taxon Volborthella, which has been reported
previously from lower Cambrian strata of the southern Great Basin (e.g., Signor and
Ryan, 1993; Hagadorn and Waggoner, 2002; Bailey et al., 2006). Because Volborthella
consists of agglutinated heavy mineral grains, including magnetite, pyrite, and zircon
(e.g., Lipps and Sylvester, 1968; Bailey et al., 2008), it shows up as bright spots on x-ray
negatives. Volborthella cones are loosely packed within the funnels, and are usually
found within massively bedded infi ll of funnels.
In cross section, the tubes found at the center of the isolated sediment mounds are
vertical to sub-vertical in orientation and can extend to a depth of up to 10 cm, although
most fall within 2-4 cm in length (Fig. 5.5G-I). The walls of the tubes are irregular and
the length of the tube does not necessarily form a straight shaft (Fig. 5.5I). The base of
the tube can be blunt and rounded or can taper to a point. At the top of these vertical
tubes is found a funnel that expands toward the upper surface of the bed and merges
with the capping sediment mound. At their widest, these subsurface funnels reach a size
of 1.2-2 cm in diameter. The junction of the funnel with the underlying vertical tube
occurs at a depth of 1-2 cm. Laminated infi llings of funnels consist of convex-downward
projecting alternating sand and silt drapes that mirror the topography of the funnel. The
sediment mound can have a massive-looking internal structure or can consist of convex-
upward laminae that merge laterally with the adjacent host sediment.
In thin section, the isolated surface funnels can be seen to consist of convex-
downward drapings of sand and mud that parallel the geometry of the underlying
funnel morphology (Fig. 5.6A). Also, the infi ll of the funnel tends to contain a higher
mud or silt content than the underlying shaft of the burrow, which consists primarily
138
deflected
laminae
irregular
margin
funnel
irregular
margin
diffuse
boundary
A
B C
D
E F
G H
I
1 cm 1 cm
sediment
mound
funnel
vertical
tube
funnel
tube continues
1 cm 1 cm
1 cm 1 cm
1 cm 1 cm
funnel
Volborthella
cones
1 cm
sediment mound
FIGURE 5.5. X-radiographs and polished slabs of cross sections of isolated funnels and
sediment mounds. A) X-radiogrpah of an isolated funnel with a downward tapering tube.
The tube has an irregular margin and defl ects adjacent laminae. B) X-radiograph of an
isolated funnel with an irregular margin that tapers downward. C) X-radiograph of an
isolated funnel with a diffuse boundary and loosely packed Volborthella cones within a
massive infi ll. D) X-radiograph of two surface funnels that connect in the subsurface to
form a u-shaped tube. E) X-radiograph of a surface funnel with a laminated infi ll leading
downward into a vertical tube with an irregular margin. F) X-radiograph of a shallow-
penetrating funnel without a signifi cantly developed subsurface tube. G) Polished slab
of small sediment mound capping a vertical tube. H) Polished slab of isolated surface
sediment mound surrouding a funnel. The vertical tube continues downward past
what can be seen, as it leaves the plane of view (sediment mound from fi g. 5.4G). I)
X-radiograph of sediment mound and funnel from fi g. 5.5H, showing an irregular bend in
the vertical tube.
139
of clean sand. Funnels associated with surface sediment mounds are also fi lled with
laminated sediment that alternates between sand-rich and mud-rich laminae (Fig. 5.6B).
These laminae have a convex-downward geometry and contour the morphology of the
underlying funnel shape. The sediment that comprises the mound itself can consist
of clean quartz sand, or can consist of alternating sand and mud layers (Fig. 5.6C). In
the case of alternating layers, mud-rich layers are well-sorted and generally deprived
of coarse grains (Fig. 5.6D). Sand-rich layers, however, are poorly sorted and consist
of a mud matrix with fl oating quartz grains. Also found within the sand-rich layer are
elongate mica grains that are interspersed with quartz grains and are oriented parallel to
the surface slope of the sediment mound (Fig. 5.6E).
Sediment Mound Formational Processes in the Modern
Sediment mounds at the surface of a substrate can form due to both physical
and biological processes. In modern marine environments, biogenic sediment mounds
typically form from the expulsion of material from a burrow opening due to excavation
or deposit feeding (e.g., Swinbanks, 1981). In the case of deposit feeding behavior, the
sediment of the mound will consist of aggregates of fecal castings. For simple excavation
behavior the sediment is typically just ejected from the top of the burrow, where it
accumulates as a growing sediment mound (Fig. 5.7A). Water currents generated by
the burrowing organisms are typically responsible for the sediment transport associated
with this excavation behavior and can lead to the size sorting of sedimentary particles.
Due to the limited transport capacity of these types of biogenic currents, fi ner particles
are preferentially ejected from the burrow opening. Following ejection, these particles
accumulate within the surface sediment mound, while coarser particles too heavy to be
transported are preferentially left behind (e.g., Swinbanks, 1981; Ziebis et al., 1996).
140
AB C
DE
mud drapes in funnel
D
E
sand-rich layer
mud-rich layer
mud-rich layer
sand-rich layer
5 mm 5 mm 5 mm
1 mm 500 μm
FIGURE 5.6. Photomicrographs of isolated funnels and sediment mounds. A) Isolated
funnel consisting of an sand-rich infi ll with mud laminae drapes that follow the funnel
morphology. B) Convex-downward alternating sand-rich and mud-rich laminae within
a funnel associated with a surface sediment mound (polished slab from fi g. 5.5H). C)
Small sediment mound consisting of a lower mud-rich layer and an upper sand-rich layer
(polished slab from fi g. 5.5G). D) Close-up of fi g. 5.6C showing the diffuse boundary
between the mud-rich layer and the sand-rich layer. E) Close-up of fi g. 5.6D showing that
the sand-rich layers consist of a mud matrix with fl oating quartz grains. Also, abundant
mica grains are present and are oriented concordant with the slope of the surface
sediment mound.
141
Sandy sediment mounds that form through physical processes mostly fall under
the category of injection features. Sediment can be ejected from the substrate due to
many causal factors including fl uidization of sediment during sedimentary compaction,
the expulsion of gas, and upwelling of subsurface water to the sediment surface. The
fl uidization of sediment typically occurs from the rapid compaction of water-rich
sediment, which can be prompted by slumping, rapid sedimentation, or seismicity. The
resultant features of this fl uidization are sand volcanoes (Fig. 5.7B). Sand volcanoes
are roughly circular in shape, consist of a sediment mound that exhibits relief above
the surrounding substrate, and exhibit a central crater or depression (e.g., Gill and
Kuenen, 1957; Neumann-Mahlkau, 1976). In cross section, a vertical shaft commonly
extends down from the central crater and passes through the sediment mound into the
underlying substrate to a source bed from which the sediment originated. This source
bed is typically of a lower density than the capping bed near the sediment surface, and the
capping bed typically has an irregular or convoluted base (Burne, 1970). A funnel-shaped
morphology can be developed where the shaft passes through the sediment mound (e.g.,
Burne, 1970) and this funnel morphology refl ects the downward defl ection of laminae
from the sediment mound as they approach the central shaft. This downward defl ection
is very similar to that produced by simple sediment collapse (e.g., Buck and Goldring,
2003).
Sediment mounds can also be formed from the expulsion of gas, which is most
often associated with the development of surface microbial mats (e.g., Gerdes et al.,
1993). The most common features that result from the accumulation of trapped gas
beneath a surface microbial mat are gas domes and gas pits (e.g., Gerdes et al., 1993;
Draganits and Noffke, 2004; Dornbos et al., 2007) (Fig. 5.7C). Gas domes consist of
a mound-shaped morphology with an interior central cavity that is buoyed-up by gas
pressure. This gas forms from the subsurface degradation of organic matter (Gerdes et
142
5 cm 5 cm 5 cm
funnel
mound
respiration
current
respiration
current
sand volcano
gas-filled
hollow space
microbial
mat
clay
lamina
convolute
base
source
pipe
reduced
slushy
sediments
zone of decay
and gas production
degassing
channel
source bed
AB C
gas
dome
FIGURE 5.7. Methods by which sediment mounds can form. A) Burrow of Arenicola
consisting of a u-shaped tube with a surface funnel and a surface mound (after
Swinbanks, 1981). Respiration currents expel fi ne-grained particles from the surface
funnel, where they accumulate as a clay lamina. Fecal casting are ejected from the
other burrow opening to form a surface mound. B) Sand volcanoes are the result of
sediment injection from a subsurface source bed during dewatering or compaction (after
Burne 1970; Gill and Kuenen 1957) . The sediment is carried upward through a pipe
and accumulates at the sediment surface, while the contact between the source bed and
overly material becomes convoluted during compaction. C) Gas domes result from the
subsurface decay of organic matter, which releases gas that can accumulate beneath a
sealing microbial mat (after Gerdes et al., 1993).
143
al., 1993). The preservation of this central cavity is only possible with early cementation
or mineral infi lling. If the gas manages to escape and the dome collapses, the resulting
feature is a gas pit. Gas pits consist of a doughnut-shaped circular dome with a shallow
central depression or pit. This pit refl ects the collapse of a prior gas dome. The top of the
dome near the central pit is typically fl at-topped (Draganits and Noffke, 2004; Dornbos et
al., 2007).

Early Cambrian Sediment Mound and Funnel Origins
Given the data, the most parsimonious interpretation is that the mounds and
funnels on the examined bedding plane have multiple origins. The isolated funnels and
isolated sediment mounds have features most similar to biogenic sediment mounds and
funnels that form through the expulsion of sediment during deposit feeding or excavation.
Despite differences in surface morphology, most surface isolated sediment mounds
and funnels have a subsurface cross section that is funnel-shaped at the top, leading
downward into a vertical tube. The funnel morphology found in the subsurface cross
section is typical for vertical burrows in unconsolidated sediment. The surface funnel
morphology results from the collapse of loose sediment around the burrow aperture
as the organism shifts up and down within its constituent tube (Hallam and Swett,
1966; Stamhuis et al., 1997). Surface funnels may also be formed from the drainage
of water into the burrow aperture, resulting in surface scour and cave-in (Boyd, 1966).
Subsequent burrowing allows for the re-establishment of the vertical tube if it becomes
fi lled with sediment. The irregular, winding, and diffuse nature of the margins of the
funnels and associated tubes, as well as the presence of u-shaped morphologies, is most
consistent with bioturbation. Sediment injection and dewatering tends to produce sharp
tube walls with straight, vertical geometries (Gill and Kuenen, 1957; Burne, 1970), and
are not known to produce pipes with bends or u-shaped morphologies. Also, sediment
144
injection often leads to the convolution of bedding and laminae, as well as radiating
rill marks at the sediment surface (Boyd and Ore, 1963). No such features have been
observed in any samples. In addition, the concentration of fossil debris within some
funnels likely indicates that the funnels were open features that were fi lled from above,
rather than from below. This concentration makes sense as burrows are more likely
to capture passively transported biogenic particles than surrounding fl at sediments
(e.g., DePatra and Levin, 1989). No source bed for a sediment injection origin can be
recognized, and the fossils concentrated within the funnels were likely colonizing the
sediment surface, rather than coming up from the shallow subsurface. The laminated
infi ll of some funnels also indicates a top-down fi lling. This type of laminated fi lling
is common in tidal fl at environments in which alternating tidal current and slack-water
conditions produce laminated sand and mud couplet laminae within the burrow (Gingras
et al., 2000).
Sediment mounds associated with these interpreted funnel-shaped burrows are
thought to be biogenic, rather than perhaps sediment injection through an existing burrow.
Biogenic sediment mounds tend to be irregular and often consist of sediment grains that
are fi ner than those that comprise the burrow and can consist of fecal pellets (e.g., Frey,
1970). These mounds are typically low-relief, small, and are restricted by the amount of
sediment excavated from the burrow (e.g., Frey, 1970; Swinbanks, 1981; Duncan, 1987),
although deeper and more complex burrows can have larger mounds (e.g., Curran and
Martin, 2003; Ziebis et al., 1996). Mounds associated with sediment injection, in the
form of sediment volcanoes, can range in diameter and height from the centimeter-scale
to the meter-scale and are commonly associated with convolute bedding and slumping, as
well as radiating rill marks or channels (e.g., Dionne, 1976; Johnson, 1977). Given what
can be ascertained, the relatively uniform size of tubes and associated subsurface funnels,
the laminated nature of the associated funnel fi ll, and the size sorting within the mound
145
itself suggest a biogenic origin for these mounds. The lack of any of the typical physical
structures indicating sediment injection and the laminated infi ll of the associated funnels
suggest that these were indeed open burrows. The sorting of sediment within the mound
is atypical for sediment extruded during dewatering, which is typically fi nely laminated
(e.g., Burne, 1970). Rather, the sorting is likely biogenic in origin. The concentration of
mica and quartz grains within a fi ne-grained matrix is typically considered a signature of
trapping and binding activity of a microbial mat, whereby the elongate grains adhere fl at
to the mat surface and become trapped (Gerdes et al., 2000; Noffke et al., 2003). This
conclusion seems valid considering that biogenic sediment mounds within microbial mat-
dominated environments can be stabilized by microbial mats (e.g., Curran and Martin,
2003; Baucon, 2008). While these isolated mounds are rare and cannot be scrutinized in
detail, it is apparent that the signatures of physical processes for their origin are lacking,
and a biogenic origin fi ts better with the observations.
Sediment mounds found within shallow depressions are the most enigmatic of
the features exposed on the bedding plane and appear to have an origin not associated
with bioturbation. These sediment mounds were most likely formed due to the erosion of
microbial gas domes. Gas domes are typically preserved in full relief as actual domes or
are expressed in their collapsed form as gas pits, with a small central depression within
a fl at-topped mound (Draganits and Noffke, 2004; Dornbos et al., 2007). For these
early Cambrian examples, the most telling feature of the shallow depressions associated
with these sediment mounds is the irregular nature of the margin separating the outer
bulge from the trough. The irregularity of this margin is highly diagnostic for the torn
edge of a surface microbial mat (Noffke, 1999; Noffke, 2009), and is typically not found
in unbound sediment. This is a reasonable conclusion considering that this bedding
surface is known to have been coated by microbial mats in the form of wrinkle structures.
The sediment mound found at the center of the depression may be closely tied to the
146
depression itself, and may represent sediment injection following a phase of compaction
or depressurization.
The following model is proposed for the formation of these features: 1) the
formation of a surface microbial mat across a tidal fl at surface results in the sealing of
the underlying sediment; 2) accumulated subsurface gas leads to the development of
a surface gas dome that is fed by a subsurface degassing channel or pipe (Fig. 5.8A);
3) strong tidal currents lead to the erosion of the surface gas dome, leaving behind a
circular-shaped torn microbial mat margin where the dome connected to the underlying
substrate (Fig. 5.8B); 4) rapid degassing and depressurization of the underlying sediment
leads to compaction and subsidence, resulting in sediment expulsion through the former
degassing channel; this results in the generation of a surface sediment apron and the
formation of a shallow depression beneath it (Fig. 5.8C); 5) the torn microbial mat
margins collapse or fl op down onto the substrate forming the outer bulge, the inner
irregular margin refl ects the torn margin of the microbial mat, and the smooth outer
portion of the bulge marks the location where the former gas dome intersected the
bedding surface (Fig. 5.9). While this model is plausible given the known depositional
conditions and does explain the morphological features observed on the bedding
surface, it is diffi cult to confi rm the predicted subsidence beneath the sediment mounds,
which would be observed as down-warped laminae. Also, it might be expected that
the subsurface lamination beneath the sediment mounds should exhibit some degree on
convolution, as the process being invoked is sediment injection. It is diffi cult, at present,
to test these assertions until adequate samples can be acquired.
Trace Fossil Affi nities
The isolated funnels and sediment mounds found in the middle member of
the Wood Canyon Formation can be classifi ed into three distinct morphologies, based
147
gas-filled
hollow space
microbial
mat
gas
dome
trough
sediment
apron
outer bulge
(collapsed
margin)
torn
margin
erosive
tidal
currents
torn margin
of gas dome
torn margin
of gas dome
zone of decay
and gas production
zone of decay
and gas production
degassing
channel
rapid
pressure
release
compaction
and
subsidence
expulsion
of sediment
AB C
FIGURE 5.8. Model for the formation of sediment mounds within shallow depressions.
A) Accumulated subsurface gas beneath a microbial mat leads to the development of
a surface gas dome that is fed by a subsurface degassing channel or pipe. B) Strong
tidal currents lead to the erosion of the surface gas dome. C) Rapid degassing and
depressurization of the underlying sediment leads to compaction and subsidence,
resulting in sediment expulsion through the former degassing channel; this leads to the
generation of a surface sediment apron and the formation of a shallow depression beneath
it. Also, the torn margins of the microbial mat collapse to form a fl at-topped bulge around
the depression.
148
contact between
former gas dome
and substrate
mound formed from
sediment injection
during compaction
degassing pipe
utilized during
sediment
injection
irregular
torn microbial
mat margin 5 cm
FIGURE 5.9. Interpretations of the observed features for the sediment mounds found
within shallow depressions. The contact between the outer bulge and the outlying
sediment surface marks the position where the former gas dome connected to the
substrate. The irregular margin where the outer bulge contacts the trough represents the
torn margin of the microbial mat that once constituted the former gas dome. The fl at-
topped nature of the bulge itself is the result of the inward collapse of this torn margin.
The vertical pipe represents the former degassing chamber, while the surrounding
sediment mound likely formed from the expuslsion of sediment during compaction.
149
primarily on cross sections: 1) u-shaped tubes with surface funnels; 2) vertical tubes with
surface funnels; 3) vertical tubes with surface funnels surrounded by a sediment mound.
The u-shaped tubes with surface funnels are most closely allied with the ichnogenus
Arenicolites (Fig. 5.10A). Arenicolites consists of a simple u-shaped tube with the
apertures, sometimes enlarged or funnel-shaped, connecting to the sediment surface
(Howard and Frey, 1984). In the modern ocean, these types of u-shaped burrows can
be constructed by the polychaete Arenicola (e.g., Swinbanks, 1981) and the amphipod
Corophium (e.g., Gingras and Bann, 2006).
The vertical tubes with surface funnels fall within the ichnogenus Monocraterion
(Fig. 5.10B). Monocraterion consists of a funnel-shaped burrow (up to 4 cm in diameter)
that is penetrated by a vertical shaft than extends below the funnel (e.g., Hiscott et al.,
1984). Many groups of organisms are capable of generating vertical burrow shafts, and
in many instances the surface funnel can be tied to sediment collapse or drainage into the
burrow during feeding (Boyd, 1966; Stamhuis et al., 1997). For these reasons, assessing
the possible type of trace maker for these burrows is diffi cult. Modern organisms
capable of making this type of funnel burrow with a vertical tube include the acorn worm
Balonoglossus (e.g., Duncan, 1987), the polychaete Diopatra (e.g., Boyd, 1966), and
some modern ghost shrimp (e.g., Rowden and Jones, 1995; Ziebis et al., 1996; Stamhuis
et al., 1997). While these organisms are capable of producing the funnel and vertical
shaft, their burrows tend to be far less simple and include u-shaped geometries and
complex branching patterns.
The vertical tubes with sediment mounds are similar to those described for
the ichnogenus Chomatichnus (Fig. 5.10C). Chomatichnus is defi ned as a conical
mound of fecal castings surrounding a vertical shaft, and was recognized and defi ned
originally within a well-preserved Carboniferous limestone deposit (Donaldson and
Simpson, 1962). The recognition of castings in the rock record is based highly on the
150
quality of preservation, and subsequent documentation of Chomatichnus has focused
primarily on the morphology of the trace fossil, rather than the recognition of castings
(e.g., Chamberlain, 1971; Dawson and Reaser, 1984). In the absence of fecal castings,
Chomatichnus is defi ned simply as a conical mound surrounding a vertical shaft. The
vertical tubes and sediment mounds from the middle member of the Wood Canyon meet
these criteria and should likely be placed within this ichnogenus.
Burrows similar to Chomatichnus can be made through several processes. For
numerous modern organisms the main process by which a biogenic surface sediment
mound forms is through the deposition of fecal castings around the burrow aperture.
While some organisms such as fi ddler crabs simply deposit a scattering of castings
around the openings of their burrows, only a few produce what could actually be termed
a mound. Balanoglossus is one such organism, and can generate large fecal castings that
attain relief above the substrate (Frey, 1970). Modern ghost shrimp are also well-known
for generating large sediment mounds and in some groups produce simple vertical shafts
with only minimal branching or winding (Griffi s and Suchanek, 1991). While some
degree of deposit feeding and expulsion of fecal casting occurs in association with the
excavation of sediment and the construction of sediment mounds (e.g., Pryor, 1975), the
expelled sediment does not necessarily need to be in the form of fecal castings (Griffi s
and Suchanek, 1991), but can simply be higher in organic content that the surrounding
sediment (Stamhuis et al., 1997).
The polychaete Lanice conchilega also inhabits a vertical burrow and forms a
sediment mound around its aperture, but does so in response to sedimentation (Carey,
1987). When these polychaete burrows are fi lled by sedimentation, the organism will
re-excavate the burrow and deposit the sediment as a mound around the aperture (Carey,
1987). In this instance, the biogenic mound is not comprised of fecal castings. In the
absence of fecal casting or pellets within the Wood Canyon sediment mounds, which may
151
Isolated Funnel
[Monocraterion]
Isolated Sediment Mound
[Chomatichnus]
U-shaped Tube
[Arenicolites]
AB C
FIGURE 5.10. Burrows from the middle member of the Wood Canyon Formation and
their assigned ichnotaxonomy. A) The u-shaped burrows with surface funnels are most
closely associated with the ichnogenus Arenicolites. B) The isolated funnel burrows
belong to ichnogenus Monocraterion. C) The isolated sediment mound burrows fall
within the ichnogenus Chomatichnus.
152
be due to original absence or poor preservation, it is inconclusive whether these mounds
represent deposit feeding behavior or simple sediment excavation.
Conclusions
The lower Cambrian middle member of the Wood Canyon Formation in the
southern Salt Spring Hills preserves funnels and sediment mounds in association with
wrinkle structures, which represent the former presence of a surface microbial mat.
These features occur within heterolithic strata of a mixed tidal fl at environment. Three
morphologies are recognized in bedding plane view: 1) sediment mounds found within
shallow depressions; 2) isolated funnels; and 3) isolated sediment mounds. Sediment
mounds found within shallow depressions are interpreted as the result of the erosion
of a microbially-stabilized gas dome. The shallow depression formed as a result
of compaction during depressurization of the gas dome, while the sediment mound
resulted from sediment injection during this compaction. Isolated funnels are connected
to subsurface vertical tubes that have irregular margins. These isolated funnels are
interpreted as funnel-shaped burrows assignable to the ichnogenus Monocraterion.
In rare instances, surface funnels connect in the subsurface to form u-shaped tubes
assignable to the ichnogenus Arenicolites. Isolated sediment mounds can be connected
to subsurface vertical tubes or to a subsurface funnel. These sediment mounds and tubes
are assignable to the ichnogenus Chomatichnus. Funnels associated with these latter two
morphotypes contain a laminated infi ll consisting of alternating layers of sand and mud,
indicating a passive fi ll from above, rather than sediment injection from below. Also,
sediment mounds can exhibit a similar layering, which results possibly from trapping and
binding by a microbial mat.
The presence of wrinkle structures across the studied bedding plane indicates
that it was extensively colonized by microbial mats. The interpretation that the sediment
153
mounds within shallow depressions represent former gas domes indicates that these
microbial mats were capable of sealing-off underlying sediment, resulting in the
accumulation of subsurface gas following the decay of organic matter. The types of
burrows, including funnels and sediment mounds, developed across this ancient tidal fl at
surface are very similar to those documented for modern microbial mat-covered tidal fl at
environments (e.g., Curran and Martin, 2003; Baucon, 2008; Bertics and Ziebis, 2009).
Fecal castings and their resultant sediment mounds typically have a low preservation
potential and can be easily reworked and redistributed by tidal currents (Oertel, 1973;
Pryor, 1975). The fact that these sediment mounds are indeed preserved may be due to
microbial stabilization; however, a microbial signature within the burrow mounds—aside
from potential trapping and binding fabrics—is not entirely clear. While the presence of
associated wrinkle structures and gas domes indicates that the underlying substrate was
highly stabilized, it cannot be confi rmed whether this stabilization was extended to the
burrow mounds as well.
Tidal fl at environments of the middle member of the middle member of the Wood
Canyon Formation provide a record of deep-penetrating vertical burrows (up to 10 cm in
depth) in the early Cambrian. In spite of the depth of bioturbation, burrows are widely
distributed across the bedding plane and generate an ichnofabric index of no more than
2. This low levels of sediment mixing may be responsible for the coexistence of these
burrows and associated microbial mat features, a phenomenon documented on many
modern coastlines (e.g., Curran and Martin, 2003; Baucon, 2008; Bertics and Ziebis,
2009). This record differs from deep-penetrating vertical burrows from the overlying
upper member of the Wood Canyon Formation, which preserves deep-penetrating vertical
burrows (up to 30 cm in depth) within a lagoonal environment that generate ichnofabric
indices up to 4, destroying most of the original physical sedimentary structures (Mata et
al., in review). The tidal fl at burrow mounds and funnels from the middle member of the
154
Wood Canyon Formation record the earliest occurrence of complex interactions between
vertical burrows and microbial mats in the southern Great Basin and represent behaviors
more complex than other microbial mat-associated horizontal burrows found throughout
early Cambrian shallow shelf environments of the region (e.g., Hagadorn and Bottjer,
1999; Marenco and Bottjer, 2008). Further documentation of these complex animal-
microbial mat interactions and their paleoenvironmental context can provide additional
insight into the dynamics of the Cambrian radiation and may shine new light on the
antiquity of these types of behaviors.
155
CHAPTER VI
Early Cambrian Anemone Burrows from the Upper Member of the Wood Canyon
Formation, Death Valley Region, United States: Paleoecological and Paleoenvironmental
Signifi cance
Introduction
Large and vertically-oriented trace fossils (deeper than 6 cm) are rare in lower
Cambrian deposits representing shallow shelf and deeper-water environments (Ausich
and Bottjer, 1982; Bottjer and Ausich, 1986; MacNaughton and Narbonne, 1999;
McIlroy and Logan, 1999; Marenco and Bottjer, 2008), which are characterized by
shallow-penetrating and horizontally-oriented trace fossils (Droser and Bottjer, 1989;
MacNaughton and Narbonne, 1999; McIlroy and Logan, 1999; Marenco and Bottjer,
2008). Deep-penetrating vertical burrows do occur, however, in surprising abundance
and in dense populations in lower Cambrian strata deposited in nearshore settings ranging
from shoreface to intertidal (Cornish, 1986; Hiscott et al., 1984; Droser, 1991; McIlroy
and Garton, 2004). The most ubiquitous Cambrian deep-penetrating, nearshore burrows
are those assignable to the ichnogenus Skolithos, which commonly occur as dense
concentrations within nearshore sandstones, colloquially termed ‘pipe rock’ (Peach and
Horne, 1884; Hallam and Swett, 1966; Droser, 1991; McIlroy and Garton, 2004). Pipe
rock emerged in the early Cambrian, reached peak abundance during the Cambrian, and
experienced a long-term decline through the Paleozoic (Droser, 1991). Skolithos burrows
within these pipe rock deposits are commonly up to 15 cm in length and are generally
less than a centimeter in diameter (Miller and Byers, 1984; Droser, 1991).
The lower Cambrian upper member of the Wood Canyon Formation in the Death
Valley region, USA, preserves large vertical burrows up to 30 cm in length and 7 cm in
diameter within oolitic and sandy dolostones and dolomitic sandstones (Corsetti, 2007).
156
These burrows predate the appearance of extensively developed Skolithos ‘pipe rock’ in
the region, found in the overlying lower Cambrian Zabriskie Quartzite (Stewart, 1970;
Barnes and Klein, 1975; Droser, 1991; Prave, 1992). The large upper Wood Canyon
burrows exceed typical Skolithos burrows in terms of length and diameter, and differ
in shape. The burrows are assignable to the ichnogenera Bergaueria, Conichnus, and
Dolopichnus, all of which have been interpreted previously to represent the behavior of
anemone-like organisms (Prantl, 1945; Alpert, 1973; Alpert and Moore, 1975; Howard
and Frey, 1984; Pemberton et al., 1988). The purpose of this study is to characterize
the depositional setting, taphonomy, and paleoecology of these trace fossils to explore
ecospace utilization, including extent and depth of bioturbation, recorded by ichnofabric
indices (sensu Droser and Bottjer, 1986) and depth of tiering (sensu Ausich and Bottjer,
1982), respectively, during the earliest Phanerozoic.
Geologic Background
The Wood Canyon Formation consists of three informal members—the lower,
middle, and upper members—established by Stewart (1966) and encompasses the
Precambrian-Cambrian boundary, which is found in its lower member (Fig. 6.1) (Corsetti
and Hagadorn, 2000; 2003). The upper member of the Wood Canyon Formation is
lower Cambrian in age and crops out over a wide region of eastern California and
western Nevada near Death Valley. According to updated Cambrian chronology (e.g.,
Babcock and Peng, 2007) nearly all of the upper member of the Wood Canyon Formation
falls within Stage 3 of Cambrian Series 2 because it is stratigraphically above the fi rst
occurrence of trilobites in the region, found near the base of the upper member (e.g.,
Diehl, 1979; Hunt, 1990), but lies stratigraphically below the fi rst occurrence of the
trilobite genera Olenellus or Redlichia (e.g., Hunt, 1990; Mount et al., 1991). The Wood
Canyon Formation is part of a thick succession of strata that was deposited on a thermally
157
Planolites
first trilobites
archaeocyaths
Ernietta
T. pedum
Swartpuntia
Cloudina
Johnnie
Stirling
W ood
Canyon
Cambrian Ediacaran
Zabriskie
Carrara
limestone
shale
dolostone
conglomerate
quartzite
int. sand/siltstone
Death Valley Succession
A B
arthropod traces
Emigrant
Pass
116°W
36°N
Sho-
shone
Death Valley
Primary Locality
Other Localities
40 km
NEV ADA
CALIFORNIA
Daylight
Pass
Beatty
lower middle upper
Bare Mt.
FIGURE 6.1. A) Generalized stratigraphy of the Neoproterozoic-Cambrian succession
in the Death Valley region showing the position of the Lower Cambrian upper
member of the Wood Canyon Formation. Fossils and trace fossils depicted mark only
fi rst occurrences. B) Generalized locality map showing the regional distribution of
Neoproterozoic-Cambrian units (shaded) in east-central California and southwest Nevada
(after Stewart, 1970).
158
subsiding passive continental margin that was developed following late Neoproterozoic
rifting of Laurentia (Stewart, 1970; Stewart and Poole, 1974; Stewart and Suczek, 1977;
Fedo and Cooper, 2001).
The upper Wood Canyon Formation contains early Cambrian body fossils
including trilobites, archaeocyathans, echinoderms, inarticulate brachiopods, and hyoliths
(Stewart, 1970; Diehl, 1974; 1979) and records middle to late early Cambrian time
(Stewart, 1970; Diehl, 1974; Fedo and Cooper, 1990). According to early Cambrian
trilobite zonations (e.g., Fritz, 1972), the upper member includes the Fallotaspis trilobite
biozone, the Nevadella biozone, and terminates within the lowermost portion of the
Bonnia-Olenellus biozone (Hunt, 1990; Mount et al., 1991; Mount and Bergk, 1998;
Hollingsworth, 2005). Carbonate units within the upper Wood Canyon Formation in
the southern Nopah Range that contain the trace fossils of this study fall within the
Fallotaspis zone and are lithostratigraphically correlated to the Lower Member of
the Poleta Formation in the White-Inyo Mountains of east-central California (Nelson,
1978; Mount and Bergk, 1998; Hollingsworth, 2005). This study focuses on outcrops
of the upper Wood Canyon Formation located approximately 1.5 km west northwest of
Emigrant Pass, Inyo County, California (N 35° 53.577 ′, W 116° 04.800 ′) (Fig. 6.1B).
Depositional Environments
To address the paleoenvironmental implications of the large upper Wood Canyon
Formation trace fossils, a detailed sedimentological analysis was conducted through
a partial section of the member at Emigrant Pass, California. This section documents
a siliciclastic-carbonate transition that occurs between lowermost upper member
siliciclastics and an overlying dolostone subunit that is found toward the top of the
member. The large upper Wood Canyon Formation trace fossils are found exclusively
159
in the dolostone subunit. At the Emigrant Pass locality, the upper member of the Wood
Canyon Formation can be subdivided into 5 sedimentary facies (Fig. 6.2).
Facies I consists of interbedded sandy siltstone and fi ne to very fi ne-grained
quartz sandstone beds. Siltstones are laminated or massive and sandstone beds are 10-15
cm in thickness and contain hummocky cross-stratifi cation or quasi-planar lamination
(sensu Arnott, 1993) (Fig. 6.3A). Upper and lower contacts of sandstone beds are
typically planar and sharp. This facies is interpreted to have been deposited between fair-
weather wave base and storm wave base in the offshore transition (sensu Howard, 1972).
The siltstone was deposited during fair-weather conditions, while the sandstone beds are
the result of offshore-directed sand transport during storms by strong oscillatory fl ow,
forming hummocky cross-stratifi cation (Harms et al., 1975; Walker, 1985; Duke, 1990;
Duke et al., 1991), or combined-fl ow forming quasi-planar lamination (Arnott, 1993;
Arnott et al., 1995).
Facies II is comprised of amalgamated beds of hummocky cross-stratifi ed fi ne
to very fi ne-grained quartz sandstone with rare quasi-planar lamination (Fig. 6.3B).
Individual beds reach thicknesses of up to 40 cm and have sharp upper and lower
contacts. Facies II is interpreted to represent deposition within the lower shoreface,
above fair-weather wave base within the zone of shoaling swells. While hummocky
cross-stratifi cation forms under a wide variety of strong oscillatory fl ow conditions
(Greenwood and Sherman, 1986; Arnott and Southard, 1990), it is commonly preserved
as amalgamated sandstone beds within the lower shoreface where fi ner sediment is less
likely to accumulate between depositional events (Harms et al., 1975; Reinson, 1984;
Leckie and Krystinik, 1989; MacEachern and Pemberton, 1992). Quasi-planar laminated
sandstones are also amalgamated in the lower shoreface and refl ect a similar process,
except with combined-fl ow dominating over oscillatory fl ow (Arnott, 1993; Arnott et al.,
1995)
160
Lower
shoreface
Ebb-tidal
delta
Ebb-tidal
delta
Ebb-tidal
delta
Lower
shoreface
Offshore
transition
VC
C
M
F
VF
Silt
Shale
Lithology
Lower
shoreface
FS
FS
FS
Ebb-tidal
delta
FS
FS
Lagoon
Flood-tidal
delta
Flood-tidal
delta
Flood-tidal delta
Ebb-tidal
delta
Flood-tidal
delta
Siltstone and hummocky
cross-stratified sandstone
Hummocky cross-stratified
sandstone
Trough cross-stratified sandstone
+/- herringbone cross-stratification
Planar laminated sandstone
Small-scale trough cross-
stratified sandy dolostone
Medium-scale trough cross-
stratified oolitic dolostone
Dolostone with sandy
layers and vertical fabric
Conichnus
LITHOLOGY AND SEDIMENTARY
STRUCTURES
BIOGENIC STRUCTURES
Dolopichnus
Bergaueria
0
2
4
6
8
10
12
14
16
18
20
22
24
26
28
30
1
3
5
7
9
11
13
15
17
19
21
23
25
27
29
31
meters
FIGURE PHOTO
Erosional discontinuity
FIGURE 6.2. Partial stratigraphic section of the upper member of the Wood Canyon
Formation measured at Emigrant Pass, California, and interpreted depositional settings.
Base of section begins at N 35° 53.577 ′, W 116° 04.800 ′.
161
FIGURE 6.3. Siliciclastic facies from the upper member of the Wood Canyon Formation,
Emigrant Pass locality. A) Outcrop photograph of interbedded sandy siltstone and quasi-
planar laminated fi ne to very fi ne sandstone beds of facies I; offshore transition. Chisel
is 18.5 cm long. B) Outcrop photograph of hummocky cross-stratifi ed fi ne sandstone of
facies II; lower shoreface. Hammer is 24 cm long. C) Outcrop photograph of bidirectional
small-scale trough cross-stratifi ed medium sandstone of facies III. Pen is 15 cm long.
162
Facies III consists of fi ne to medium-grained quartz sandstone with lesser
dolomitic sandstone and sandy dolomitic packstone. Sedimentary structures include
small-scale (< 5 cm thick) unidirectional and bidirectional trough cross-stratifi cation
with rare herringbone cross-stratifi cation and planar laminae (Fig. 6.3C). Dolomitic
packstones, when present, are comprised of echinoderm ossicles, with rare archaeocyaths
and hyoliths. Facies III is interpreted to represent deposition within the surf zone or
seaward side of a tidal inlet. Trough cross-stratifi cation in the marine realm forms due
to the migration of subaqueous dunes in the surf zone (Clifton et al., 1971), while the
bidirectional nature implies a signifi cant tidal component (e.g., Hayes, 1980; Dalrymple
et al., 1990; Shanley et al., 1992). Facies III was likely deposited within the upper
shoreface, corresponding to the build-up and surf zone, and may have been infl uenced
locally by signifi cant tidal activity within an ebb-tidal delta seaward of a tidal inlet.
Planar laminae likely formed due to swash activity (Clifton, 1969), either within the
foreshore (Clifton, 1969; Clifton et al., 1971; Harms et al., 1975) or localized within
swash bars of an ebb-tidal delta (Hayes, 1980; Imperato et al., 1988).
Facies IV is comprised of oolitic dolostone with medium-scale (5-30 cm thick)
tabular and trough cross-stratifi cation with reactivation surfaces (Fig. 6.4A). At one
horizon this facies occurs within a broad channel incised into facies III (Fig. 6.4B).
Dolostones are primarily packstones and grainstones, and fossil content consists of rare
echinoderm ossicles and archaeocyaths. Trace fossils within this facies include rare
Bergaueria, rare Conichnus, and rare Dolopichnus. This facies is interpreted to represent
deposition within a fl ood-tidal delta on the landward side of a tidal inlet. The medium-
scale cross-stratifi cation with reactivation surfaces is typical for channelized fl ow within
tidal inlets and tidal deltas (Hayes, 1980; Hennessy and Zarillo, 1987; Dalrymple et al.,
1990; Shanley et al., 1992). Facies IV typically overlies facies III stratigraphically, and
this vertical succession from small-scale bidirectional cross-stratifi cation to medium-scale
163
FIGURE 6.4. Carbonate facies from the upper member of the Wood Canyon Formation,
Emigrant Pass locality. A) Outcrop photograph of medium-scale, trough and tabular
cross-stratifi ed oolitic dolostone with reactivation surfaces of facies IV . B) Outcrop
photograph of facies IV deposited within channel incised into facies III; facies IV is
overlain by facies V . This outcrop is found lateral to the main locality in which the section
of this study was measured and is correlative to a horizon approximately 25.5 meters
above the base of the partial section in fi gure 3. Hammer is 24 cm long. C) Interbedded
sandy dolostone and dolomitic sandstone with low-angle laminae that are obscured by
extensive bioturbation (ichnofabric index = 4) of facies V .
164
bidirectional cross-stratifi cation is commonly associated with the shift from the small
ebb-tidal delta to the larger fl ood-tidal delta of a wave-dominated tidal inlet complex
(Hubbard et al., 1979; Hayes, 1980). The shift in lithology from facies III sandstone to
facies IV dolostone may therefore mark a shift from the clastic-dominated seaward side
of a barrier to a back-barrier environment characterized by clastic sediment starvation and
carbonate precipitation. Ooids that comprise the dominant allochems within facies IV
were likely formed due to frequent agitation associated with ebb and fl ood-tidal currents
within the inlet and tidal delta (e.g., Gonzalez and Eberli, 1997; Rankey et al., 2006).
Facies V consists of sandy dolostone with dolomitic sandstone layers and laminae
(Fig. 6.4C). Sedimentary structures include rare planar and low-angle cross-laminae,
with most sedimentary structures obscured by extensive bioturbation with an ichnofabric
index of 4-5 (sensu Droser and Bottjer, 1986). The bioturbation is marked by a strong
vertical fabric comprised of the trace fossils Bergaueria, Conichnus, and Dolopichnus
(as discussed in further detail below). Rare allochems include ooids and echinoderm
fragments. Sandstone and dolostone beds are fi nely interlaminated and interbedded
in parts and show a strong lithologic segregation despite extensive vertical mixing by
bioturbation. Facies V is interpreted to have been deposited further landward than the
fl ood tidal delta deposits, represented by facies IV , within a back-barrier setting. The high
degree of bioturbation, coupled with low preservation of sedimentary structures, suggests
a low-energy environment, most likely an open lagoon or distal portion of a fl ood tidal
delta. Carbonate was likely formed in situ, while siliciclastic material was potentially
sourced from the seaward side of the barrier through tidal activity or storms. In storm-
infl uenced coastlines, barrier island sediments are frequently eroded down to the water
table during storm surge and can be redistributed seaward to the shoreface or landward
as washover fans into a back-barrier lagoon (Leatherman, 1979; Houser et al., 2008), and
may account for abundant sandy sediments in quiet-water lagoonal environments.
165
These facies together defi ne an ancient wave-dominated tidal inlet complex
(Fig. 6.5A-B). Characteristics of wave-dominated tidal inlet settings include small to
poorly defi ned ebb-tidal deltas and well-developed fl ood-tidal deltas that extend into
large, open back-barrier lagoons (e.g., Hubbard et al., 1979; Hayes, 1980; Berelson and
Heron, 1985). Siliciclastic material on the seaward side of the barrier was likely sourced
laterally from longshore transport—creating the barrier in the process—while carbonate
was formed in situ in the back-barrier lagoon, removed from signifi cant siliciclastic
input (Fig. 6.5A). Mixing of carbonate and siliciclastic material in some facies may be
due to exchange across the tidal inlet, and discrete carbonate beds can be found within
hummocky cross-stratifi ed sandstone units of the lower shoreface, and a signifi cant
siliciclastic component can be found in some of the back-barrier dolostones.
The presence of an incised channel into facies III sandstones may represent
the location of the actual inlet throat, which would then be overlain by fl ood-tidal
delta deposits of facies IV and lagoonal deposits of facies V in a shallowing-upward
succession (Fig. 6.4B). The absence of preserved facies deposited atop the barrier island,
including eolian dunes, may be due to their poor preservation within subaerially exposed
environments and the observation that subaqueous lagoon, tidal inlet, and shoreface
environments are much more likely to be preserved (Hoyt and Henry, 1967). Also,
barrier island deposits are volumetrically less signifi cant than tidal inlet deposits (Kumar
and Sanders, 1974) and tidal inlet migration is commonly responsible for the reworking
and destruction of barrier island facies (Moslow and Tye, 1985), in addition to extensive
erosion of barrier island sediments during storm surges (Leatherman, 1979; Houser et al.,
2008).
166
Facies III
Facies II
Facies I
Facies IV
Flood-
tidal
delta
Lagoon
Ebb-
tidal
delta
Lower
shoreface
Offshore
transition
Offshore
Fair-weather wave base
Storm wave base Facies V
Barrier island
Flood-
tidal
delta
Tidal flat
Ebb-
tidal
delta
Lower shoreface
Upper shoreface
Lagoon
Offshore transition
X
X’
X’ X
Lateral import of
siliciclastics through
longshore transport
In situ carbonate
production
Vertical burrows
A
B
FIGURE 6.5. Paleoenvironmental reconstruction of the distribution of facies and
depositional environments of the upper member of the Wood Canyon Formation in A)
plan view and B) cross-section.
167
Trace Fossils
Three distinct ichnogenera are recognized within the extensively burrowed
interval corresponding to facies V . Due to the high ichnofabric indices within most of the
interval, select discrete trace fossils near the upper and lower boundaries of the unit were
used for trace fossil identifi cation. The lack of 3-dimensionally preserved specimens
and diffi culties in observing burrow margins, possibly due to extensive dolomitization,
preclude assignment of these ichnotaxa to the ichnospecies level. Below are presented
the descriptions of the recognized ichnogenera, their taphonomy, the lines of evidence
that suggest that they are indeed biogenic structures, and the behaviors that they
represent.
Bergaueria Prantl 1945
The ichnogenus Bergaueria was originally described by Prantl (1945) and is a
common trace fossil in strata from the early Cambrian through the Recent. It is defi ned
as a hemispherical to shallow cylindrical, vertical structure with a rounded base (Alpert,
1973; Pemberton et al., 1988). Its diameter is generally greater than or equal to its length.
Burrow walls are smooth and unornamented, although a lining may be present, and the
base may contain a shallow central depression and radial ridges (Alpert, 1973; Pemberton
et al., 1988). Burrow fi ll is generally structureless and most commonly attached and
genetically related to sediment from the overlying bed forming convex hyporelief
preservation (Pemberton et al., 1988).
Trace fossils within the Wood Canyon Formation assignable to the ichnogenus
Bergaueria are hemispherical to cylindrical with a slightly rounded to blunt termination
(Fig. 6.6A-C). Burrow walls are smooth and show no evident ornamentation. Most
noted specimens have a diameter greater than length. Primary physical sedimentary
structures adjacent to the trace fossil occur as defl ected laminae or are cut cleanly by
168
FIGURE 6.6. A-C) Outcrop photographs of Bergaueria in cross section from the upper
member of the Wood Canyon Formation. All photographs are from specimens found in
facies V . Emigrant Pass locality. A-B) Bergaueria with blunt, rounded terminations with
no evident ornamentation. C) Bergaueria consisting of a shallow depression with an
upward-projecting central depression. Truncated laminae can be seen along the left fl ank
of the burrow.
169
the burrow. In rare instances, a shallow upward-projecting central depression can be
observed at the basal termination.
Conichnus Männil 1966
The ichnogenus Conichnus was originally described by Männil (1966) and has
been subsequently revised and redefi ned by Frey and Howard (1981) and Pemberton
et al. (1988). Its stratigraphic range is from the early Cambrian through the Recent.
It is generally described as a conical, downward-tapering vertical structure oriented
perpendicular to bedding that terminates at a rounded but distinct basal apex (Frey and
Howard, 1981; Pemberton et al., 1988). Structure fi ll tends toward coarse sediment at the
core, with sediment fi ning toward the periphery of the structure, and may exhibit nested
funnel or chevron-like laminae oriented convex downward (Frey and Howard, 1981;
Pemberton et al., 1988). Structures typically exhibit a lining, albeit thin, that marks a
discontinuity between the structure and the adjacent surrounding sediments (Frey and
Howard, 1981).
Burrows assignable to Conichnus within the Wood Canyon Formation are
unbranched cylindrical structures (Fig. 6.7A-C). They are straight to slightly sinuous
with a maximum diameter from 19 to 70 mm that can vary throughout the length of
the burrow, and length of the burrows extends from 55 mm to 290 mm. Burrows are
oriented perpendicular to slightly oblique to the bedding plane and are open at one end
and taper to a close at the other. Walls are generally smooth, although some specimens
show annulations along the burrow margin (Fig. 6.7A). Adjacent laminae within the host
rock can be defl ected downward (Fig. 6.7B) or can be truncated (Fig. 6.7C). Burrows
tend to occur in dense, unevenly spaced populations, and are sometimes nested within
one another. The burrows are fi lled with carbonate sediment that contains abundant
echinoderm debris that is sometimes found in aggregated piles at the base of the burrow.
170
FIGURE 6.7. A-C) Outcrop photographs of Conichnus in cross section from the upper
member of the Wood Canyon Formation. Photographs are from facies IV and facies V .
Emigrant Pass locality. A) Conichnus with irregular, annulated burrow wall (marked by
arrows). B) Conichnus burrow that defl ects adjacent laminae within the host rock. C)
Conichnus consisting of an inverted cone that truncates adjacent laminae within the host
rock.
171
Dolopichnus Alpert and Moore 1975
The ichnogenus Dolopichnus was described by Alpert and Moore (1975) and
has been revised by Pemberton et al. (1988). It is a rare trace fossil, but is known from
the lower Cambrian Poleta Formation in the White-Inyo Mountains, California, USA,
where it was originally described (Alpert and Moore, 1975), and is also known from
the lower Cambrian Bradore Formation in Labrador, Canada (Hiscott et al., 1984), the
Silurian Thorold Sandstone of Ontario, Canada (Pemberton and Risk, 1982), the lower
Mississippian Price Formation of West Virginia, United States (Bjerstedt, 1988), and the
lower Cretaceous McMurray Formation of Alberta, Canada (Pemberton et al., 1982).
It is generally described as a vertical, cylindrical to sub-conical structure containing a
central cylindrical core comprised of coarser grained material and exhibits an indistinct
termination or bulbous expansion at its base (Alpert and Moore, 1975; Pemberton et al.,
1988). Outer walls of the structure are smooth and unornamented but may show irregular
constrictions (Alpert and Moore, 1975; Pemberton et al., 1988).
Trace fossils assignable to Dolopichnus within the Wood Canyon Formation
are cylindrical to downward tapering structures that terminate in a circular to ovoid
expansion at their base (Fig. 6.8A-E). Burrows are 110-190 mm in length and 20-55
mm in diameter, while the diameter of the basal expansion ranges from 28-68 mm.
Burrow fi ll appears homogenous in cross section and is comprised of echinoderm
ossicles. On bedding plane surfaces, some burrows do show a central core of coarser
material surrounded by the outer periphery of the burrow (Fig. 6.8F). In some specimens
a central depression is found on the basal termination, similar to those found within
Bergaueria (Pemberton et al., 1988; Seilacher, 2007), and is seen in cross section as an
upward projecting point at the bottom of the basal expansion (Fig. 6.8A). Rarely, the
basal expansion can exhibit a convex-upward termination (Fig. 6.8B). Laminae adjacent
172
FIGURE 6.8. A-E) Outcrop photographs of Dolopichnus in cross section from the upper
member of the Wood Canyon Formation. Photographs are from facies IV and facies V .
A) Dolopichnus is comprised of a vertical cylindrical tube that terminates in a basal
expansion. This specimen contains a central depression on its basal termination, which is
not found in all specimens. B) Dolopichnus showing a convex-up basal termination. C)
Downward tapering Dolopichnus with a bulbous basal expansion. D-E) Some specimens
show the presence of a lateral branch; however, this is not common. F) In bedding plane
surface view Dolopichnus consists of a central core surrounded by the periphery of the
burrow. Emigrant Pass locality.
173
to the structures are rarely defl ected, unlike those associated with Conichnus within the
Wood Canyon Formation. In some specimens a small lateral branch is observed that
is positioned near the central portion of the burrow and projects upward at a slant (Fig.
6.8D-E). This branch is shorter than the main shaft and has a smaller diameter.
Burrow Taphonomy
All three types of trace fossils exhibit a similar style of preservation in which the
burrows are emplaced within dolomitic sandstone or sandy dolostone and are infi lled
with sediment dominated by echinoderm ossicles. Conichnus is the most abundant
trace fossil and therefore most readily sampled, so taphonomic analysis was conducted
primarily on this trace fossil. Analysis of thin sections indicates that the vertical burrows
were emplaced in fi rm quartz sand, which was later fi lled by a carbonate matrix (95%
carbonate grains and cement plus 5% quartz grains) and cemented with calcite (Fig.
6.9A-B). Burrow walls are generally unlined, although a micrite-rich zone is found
in some specimens; however, this may be due to stylolitization. Examination of the
contact between the burrow fi ll and the host rock reveals that the quartz sand grains at
the burrow-host rock interface are tightly packed, much more so than in the host rock
away from the burrows. This is highlighted by a thin zone (~300 μm thick) dominated
by quartz sand that separates the carbonate-rich burrow fi ll and the micritic matrix of the
host rock. Therefore, as the burrow was created within the sand, adjacent sediment was
actively compacted near the burrow wall, possibly enabling the burrows to stay open for
an extended period of time, only to be later fi lled in with a contrasting sediment type.
Evidence for Biogenicity
The burrows from the upper member of the Wood Canyon Formation fall within
the class of sedimentary structures termed ‘conical sedimentary structures’ (sensu Buck
174
FIGURE 6.9. A-B) Scanned thin section showing the boundary between burrow fi ll and
host rock in A) plan view and B) cross sectional view. The host rock consists of dolomitic
sandstone and sandy dolostone, while the burrow fi ll is dominantly echinoderm ossicles.
A zone of compacted quartz-rich sand separates the host rock from the burrow fi ll. C)
Photomicrograph of the contact between burrow fi ll and host rock. A dense concentration
of quartz sand marks a zone of compaction between the burrow fi ll and the host rock.
The micritic zone near the burrow wall may represent a lining, but may also be due to
stylolitization. Emigrant Pass locality.
175
and Goldring, 2003). Conical sedimentary structures are common throughout rocks of
many ages and can be formed due to physical or biological processes, and sometimes
both. Purely physical processes by which conical sedimentary structures can form
include gas or fl uid escape and sediment collapse into a cavity (Buck and Goldring,
2003). Biological processes by which these conical structures can form are commonly
tied to physical processes, usually collapse, and can be generated through upward and
downward locomotion, including adjustment and escape, in addition to excavation. Buck
and Goldring (2003) outlined criteria for discriminating between several types of conical
sedimentary structures that are pertinent for this study: collapse structures, locomotion
structures, and dwelling traces. These criteria are outlined below.
Collapse structures form abiogenically due to the collapse of loosely packed
sediment into a cavity. Collapse structures are defi ned by U- or V-shaped downwarping
of sedimentary laminae. The vertical displacement of laminae decreases upward to
a shallow surface depression, yet the zone of deformation widens upward. In these
collapse structures the sediment within the conical structure matches that of the host
sediment, and the boundary of the conical structure is defi ned by highly deformed shear
planes.
Locomotion structures can be formed through continuous movement or through
adjustment and equilibration. A locomotion structure formed though continuous vertical
movement is comprised of a long trail of consistent deformation with a central core that is
homogenized. This core is surrounded by an inner deformation zone where laminae are
defl ected in the direction of locomotion and an outer deformation zone where laminae are
defl ected opposite to locomotion. Alternatively, adjustment and equilibration structures
consist of laminated, nested cone-in-cone structures in which the sediment within the
cone is identical to the host sediment around it. Laminae are generally warped downward
and can be traced across the structure.
176
Dwelling traces represent the cast of the body of the organism that created and
occupied the cavity of the conical sedimentary structure. The shape of dwelling traces
can be variable, but are commonly conical and may expand downward. The lower
termination is usually rounded and a central upward-projecting depression may be present
at the base. Fills are structureless and can be variable as they are contingent upon the
sediment sources available. Burrow wall ornamentation may be present, yet weak, and
can be smooth. Also, the sharpness of the contact between the wall of the structure and
the host sediment refl ects the amount of consolidation present within the host sediment at
the time of the structure formation.
The conical structures from the upper member of the Wood Canyon Formation
are best interpreted as dwelling traces, and therefore represent a cast of the body of the
organism that made them, rather than physical processes such as collapse. The shapes
of the traces are highly variable, especially in Conichnus and Dolopichnus in which
the shape and the orientation of the boundaries change dramatically along the length
of the burrow. All three recognized ichnogenera are generally conical, with this shape
being best expressed in all Conichnus and some Dolopichnus. Downward expansion
is recognized in all Dolopichnus specimens examined, with one exhibiting a shallow
central depression on its basal termination. The fi lls of all ichnogenera examined are
structureless and are of a lithology that contrasts with the host rock, suggesting that these
are not collapse structures or locomotion structures. The sharp contact separating the
burrow fi lls from the host rock is due to a strong lithologic contrast, but may also be due
to burrow wall compaction and consolidation during the construction of the cavity. The
zone of quartz-rich sediment separating the burrow fi ll from the host rock is likely due
to active compaction during burrow formation, which is known to reduce porosity and
permeability in the adjacent host sediment (Meadows and Tait, 1989; Jégou et al., 2001).
Also, it has been shown that compaction during burrowing can lead to the expulsion
177
of fi ner particles from the interstices between coarser particles through the process of
liquefaction, leading to a segregation of grain sizes at the burrow wall (Genise and Poire,
2000). The quartz-rich zone may therefore represent a region of compaction in which
fi ner-grained particles were expelled from between coarser quartz grains, resulting in a
tighter packing of quartz sediment.
Burrowing Behavior
Past work has assumed an anemone affi nity for the ichnogenera Bergaueria,
Conichnus, and Dolopichnus (e.g., Alpert, 1973; Alpert and Moore, 1975; Frey and
Howard, 1981; Pemberton et al., 1988). Here we more rigorously document that
the morphology of these ichnogenera from the upper member of the Wood Canyon
Formation is consistent with known anemone burrowing depths and behaviors.
Bergaueria is interpreted as the resting trace or dwelling burrow of an anemone-like
organism (Prantl, 1945; Alpert, 1973; Pemberton et al., 1988) that penetrated very
shallowly into the underlying substrate, as most specimens of Bergaueria are generally
no more than a few centimeters in depth (Fig. 6.10A). The shallow penetration may also
be due to formation as undertraces across lithologic contrasts (Seilacher, 2007). The
shallow central depression found on the base of some Bergaueria refl ects the observation
that anemones scrape sediment away from their bodies radially during burrowing, leaving
a relict upward-projecting point at the center of the burrow termination (Seilacher, 2007).
Conichnus is also interpreted as a dwelling burrow or resting trace of an anemone-
like animal, with a possible escape equilibration component (Fig. 6.10B) (Frey and
Howard, 1981). Studies of modern anemones confi rm that Conichnus-like structures
can be generated through their burrowing activity, including dwelling, equilibration, and
escape traces (Shinn, 1968; Bromley et al., 1975; Curran and Frey, 1977). Conichnus
within the Wood Canyon Formation exhibit rare instances of nested funnel laminae,
178
Escape
Dolopichnus
Bergaueria
A B D
C
Conichnus
FIGURE 6.10. Burrowing behavior represented by the actinian trace fossils found in
the upper member of the Wood Canyon Formation. A) Bergaueria represents the resting
or dwelling trace of an anemone, and is often shallow penetrating. B) Conichnus also
represents the resting or dwelling trace of an anemone, but is often associated with
features indicative of equilibrium response or escape. C) The formation of Dolopichnus
likely occurs during active burrowing of an anemone as its physa penetrates into the
sediment during elongation and D) expands to serve as an anchor during retraction into
the sediment. Modifi ed from Shinn (1968); Ansell and Trueman (1968); Seilacher (2007).
179
which may represent escape or an equilibration response; however, the most ubiquitous
preservation style consists of structureless burrow fi ll surrounded by downward defl ected
laminae in the adjacent host rock and is more characteristic of dwelling traces than escape
or equilibration (Buck and Goldring, 2003). Also, the presence of annulations in some
specimens may record the peristaltic movements of the anemone during burrowing, either
due to migration within its burrow or during its construction (Fig. 6.7A).
Dolopichnus was originally interpreted by Alpert and Moore (1975) as an
anemone dwelling burrow with the central cylindrical core of coarse fossil debris
representing a cast of the anemone’s coelenteron—the hollow cavity inside the anemone.
This was based on the presence of trilobite debris within the central core of the burrow,
and these fossil fragments were interpreted as the stomach contents of the anemone.
Pemberton et al. (1988) reinterpreted the central core as the actual living chamber of the
anemone and the outer burrow as a zone of displaced sediment that results from upward
and downward peristaltic contractions. The presence of a central coarser core surrounded
by a fi ner periphery is also a common style of burrow fi ll in the anemone trace fossil
Conichnus (Frey and Howard, 1981) and may just represent the way sediment is sorted
as it fi lls the burrow. In anemones, the basal portion of the animal is often shaped into
a thin-walled, bulb-like termination termed a physa that can be expanded or contracted
and is used for penetrating the sediment during burrowing or anchoring within it during
retraction into a burrow (Fig. 6.10C-D) (Ansell and Trueman, 1968). It has been
interpreted that the basal expansion found in Dolopichnus could be due to the expansion
of the physa of the anemone during anchoring and retraction into the burrow (Pemberton
et al., 1988) and has been illustrated by Ansell and Trueman (1968) for the modern
anemone Peachia hastata. The shallow central depression found at the base of the
basal expansion may also be due to the expansion of the physa, which scrapes sediment
out radially from beneath the anemone during swelling (Seilacher, 2007). This style
180
of depression is most commonly associated with the trace fossil Bergaueria; however,
it should be equally likely in Dolopichnus if each trace fossil represents a similar
behavior. Lateral branches similar to those observed in some Dolopichnus specimens
have been observed associated with the burrows of the modern anemone Ceriantheopsis
americanus, but their function is undocumented, although it is speculated that they may
represent the burrow of a juvenile anemone generated asexually by the main burrow
dweller (Frey, 1970).
The burrows from the upper member of the Wood Canyon Formation, which are
attributed to the burrowing activity of anemones, are of a size and morphology that is
comparable to those produced by modern anemones. Modern actinian (sea anemone)
and cerianthid (tube anemone) anthozoans have been well documented with regard to
their burrowing behavior, including the burrowing mechanisms and depth of penetration.
Although actinian anemones are a much more abundant and diverse order of anthozoans,
it has been noted by Bromley (1996) that cerianthid anemones may in fact have more
ichnological importance because they form deep, semi-permanent, walled burrows that
are commonly mucus-lined.
While it might be expected that burrowing depth scales with body size, this is not
necessarily the case, and burrow depth can be much greater than the actual length of the
anemone. The cerianthid anemone Ceriantheopsis americanus is capable of producing
burrows of 30-35 cm in depth, while its actual body size is no more than 5-15 centimeters
in length (Frey, 1970; Mariscal et al., 1977; Kristensen et al., 1991; Holohan et al., 1998).
Cerianthus lloydii, also a cerianthid anemone, can reach body lengths of up to 30 cm
and is capable of submerging the entirety of its body completely within the sediment
(Schäfer, 1962), thereby making its burrow approximately the length of its body. The
actinian anemone Peachia hastata, which is known to produce structures with a similar
morphology to the ichnogenus Dolopichnus, has documented burrowing depths of up
181
to 10 cm (Ansell and Trueman, 1968), while their body ranges from approximately 5-8
cm in length (Stephenson, 1935). Deeper burrows are made by the actinian anemone
Harenactis attenuata, which can reach lengths of 40-50 cm at full extension (Child,
1908) and has documented burrowing depths of up to 30 cm (Ricketts et al., 1985). H.
attenuata burrowing behavior is similar to that of P . hastata whereby anchoring by an
expanded physa allows for retraction into the burrow (Ricketts et al., 1985).
Although modern anemones may not be perfect analogs for ancient organisms of
the same class, it is clear that they are capable of producing the depth and architecture
represented by the burrows of the upper member of the Wood Canyon Formation and
that large and deep burrows do not necessitate large trace makers. The lack of signifi cant
evidence for equilibration within these burrows (e.g., nested cone-in-cone structures
matching the host sediment), except in rare instances, and the nature of the burrow fi ll
sediments that extend the length of the burrow without interruption suggests that these
structures were indeed open burrows that were fi lled by one or more depositional events
following the vacation of the structure by the burrower.
Discussion - Regional Extent
The upper member of the Wood Canyon Formation also occurs at the California-
Nevada border near Daylight Pass, Nevada (Fig. 6.1B). Here, the upper member is at
least twice as thick as the section at Emigrant Pass and contains a greater proportion of
fi ne-grained sediment. An extensive search reveals vertical burrows reminiscent of those
found at Emigrant Pass, but the contrast between the burrows and the surrounding rock
is so low that discrete burrows are diffi cult to observe and the burrow walls are diffi cult
to separate from stylolites that are pervasive in this particular locality. Nevertheless, the
presence of the burrows at the Daylight Pass locality indicates that the burrow makers
were not limited to the Emigrant Pass area; rather, the Emigrant Pass locality simply
182
exhibits the best preservation. A third locality near Bare Mountain contains Emigrant
Pass-style bioturbation, but as with Daylight Pass, the preservation is inadequate for
further study (Fig. 6.1B).
Discussion - Lower Cambrian Anemone Burrows
Lower Cambrian assemblages containing abundant anemone-style trace fossils are
also known from Labrador, Canada and southern Poland. The L’Anse-au-Clair Member
of the Bradore Formation of Labrador is comprised of small to large scale (30-150 cm
thick) sets of tabular and trough cross-stratifi ed quartz sandstone with bidirectional
paleocurrents, in addition to thoroughly bioturbated sandstone with Skolithos (Hiscott et
al., 1984). These facies are interpreted as a progradational succession with tabular cross-
stratifi ed sandstones of an ebb-tidal delta, trough cross-stratifi ed sandstones of a tidal
inlet, and heavily bioturbated sandstones of an intertidal sand fl at (Hiscott et al., 1984).
The trace fossils Conichnus (up to 19 cm in length), Dolopichnus (up to 27 cm in length),
and ?Monocraterion are common in tabular cross-stratifi ed and trough cross-stratifi ed
sandstones, representing ebb-tidal delta and tidal inlet environments respectively (Hiscott
et al., 1984). This distribution is quite similar to that observed in the upper member of
the Wood Canyon Formation.
The Gocza łkowice Formation of southern Poland contains anemone-style trace
fossils assignable to Bergaueria (32-56 mm in length), Conichnus (15-76 mm in length),
and Conostichus (35-40 mm in length) within lower shoreface and upper shoreface
deposits associated with other vertically-oriented trace fossils including Diplocraterion,
Monocraterion, and Skolithos (Pacze śna, 2010).
Little is known about the depositional setting in which Dolopichnus of Alpert
and Moore (1975) from the Poleta Formation of the White-Inyo Mountains formed,
aside from that it occurs in fi ne to very fi ne-grained quartzite with trilobite fragments
183
infi lling the burrows. This may indicate a nearshore environment, but without further
documentation an interpretation remains lacking.
While the anemone-style trace fossils Conichnus and Dolopichnus appear
to be uncommon in lower Cambrian strata, Bergaueria is present in a wide range of
environments globally, but occurs typically as an accessory trace fossil and not as a
prominent assemblage component, although concentrations do occur (Pemberton and
Magwood, 1990). The lower Cambrian succession of the White-Inyo Mountains where
Dolopichnus was defi ned is known to also contain Bergaueria, and in addition to being
found within the Poleta Formation, it also occurs in the underlying Campito Formation
and overlying Harkless Formation, both of which are lower Cambrian in age (Nelson,
1978). In the Campito Formation Bergaueria is found in several facies, including
interbedded fi ne quartzite with hummocky cross-stratifi cation and siltstone, amalgamated
hummocky cross-stratifi ed fi ne quartzite, and planar-laminated fi ne to medium-grained
quartzite with localized channelization (Mount, 1982). These facies likely correspond
to the offshore transition, lower shoreface, and upper shoreface-foreshore—possibly
dissected by tidal channels—respectively.
Bergaueria is also known from the Chapel Island Formation of Newfoundland
where it occurs in interbedded siltstone and graded sandstone beds with current ripples
and gutter casts (Crimes and Anderson, 1985; Narbonne et al., 1987). The depositional
environment is interpreted as a storm-dominated subtidal shelf, quite similar to the types
of environments represented by the Campito, Poleta, and Harkless Formations of the
White-Inyo Mountains in which Bergaueria is found (Mount, 1982; Mount and Signor,
1985; Bailey et al., 2006).

184
Discussion - Environmental Signifi cance
It is well documented that during the Cambrian and Ordovician there is a
strong preference for nearshore origination of communities and evolutionary novelties
followed by offshore expansion (Jablonski et al., 1983). A similar trend is observed
in ichnofabrics, whereby intense bioturbation is fi rst recorded strongly in Cambrian
carbonate inner shelf settings prior to expanding into the middle and outer shelf by the
Ordovician (Droser and Bottjer, 1993). Data from this study, combined with previous
work (e.g., Hiscott et al., 1984; Pacze śna, 2010), show that the lower Cambrian deep-
penetrating burrows Conichnus and Dolopichnus show a strong preference for nearshore
environments, primarily above fair-weather wave base in the shoreface and in tidally
infl uenced settings. This parallels the record of the deep-penetrating trace fossil Skolithos
during this interval. It is notable, however, that these anemone burrows penetrate far
deeper than typical Skolithos. The shallower-penetrating Bergaueria shows a less
restricted distribution in the early Cambrian and occurs in environments from nearshore
to shelf, being found in tidally infl uenced environments in this study.
Jablonski et al. (1983) suggested that the general ‘onshore-innovation, offshore-
archaic’ pattern resulted from differential extinction rates among onshore and offshore
clades, as well as differential origination rates of new ecological adaptations or
evolutionary novelties in nearshore environments. Many nearshore clades have been
shown to be more extinction resistant, increasing the probability that innovations that
develop in nearshore communities will persist long enough for the community to
diversify and spread into offshore environments (Jablonski et al., 1983; Bottjer et al.,
1996). Additionally, although speciation rates are lower in onshore environments, the
more erratic temporal and spatial environment found in the nearshore may be more
conducive to the production of evolutionary novelties or new ecological associations or
adaptations (Jablonski et al., 1983; Bottjer et al., 1996). Deep burrowing by anemones
185
during this interval may be an evolutionary novelty that originated in nearshore
settings; however, throughout the remainder of the Phanerozoic these types of burrows,
specifi cally the more common burrow Conichnus, do not occur far from the high-energy
nearshore zone—including the shoreface, foreshore, and intertidal environments (e.g.,
Howard and Frey, 1984; MacEachern and Pemberton, 1992; Pemberton et al., 1992;
Zonneveld et al., 2001; Savrda, 2002)—except for the opportunistic colonization of sandy
storm deposits in shelf environments (Pemberton and MacEachern, 1997). Dolopichnus
is also found in deeper-water environments in the Mississippian where it occurs in
interbedded siltstone and hummocky cross-stratifi ed sandstone inner shelf deposits of the
Price Formation (Bjerstedt, 1988). So, although these deep-penetrating vertical burrows
do show an onshore origination, they do not necessarily experience a dramatic offshore
expansion throughout the remainder of the Phanerozoic. This is similar to the pattern
of the trace fossil Skolithos, which also appears to emerge in the high-energy nearshore
zone, yet is not prominent in shelf sediments throughout the Phanerozoic, except for
the opportunistic colonization of sandy storm deposits (V ossler and Pemberton, 1988;
Pemberton and MacEachern, 1997).
Conclusions
The upper member of the Wood Canyon Formation of the Death Valley region
records extensively developed and deep-penetrating bioturbation that predates the
emergence of Skolithos pipe rock in the region. Trace fossils found include Bergaueria,
Conichnus, and Dolopichnus, all of which are interpreted as the burrowing activity of
anemones. Although the trace makers of most moderate to deep-penetrating vertical
burrows in the Cambrian—including Arenicolites, Diplocraterion, Monocraterion,
and Skolithos—remain elusive, the trace makers of the early Cambrian burrows of the
upper member of the Wood Canyon Formation are most likely due to anemones. This
186
also shows that anemones were at least one of possibly many types of deep-penetrating
vertical burrowers during this interval, occupying deeper tiers than most other ichnotaxa
(up to 30 cm), and that their activity was extensive enough to destroy most or all primary
physical sedimentary structures present during deposition (ichnofabric index 4-5).
With limited preservation of soft-bodied cnidarian fossils in the Phanerozoic,
except for rare exceptional preservation (e.g., Hagadorn et al., 2002; Cartwright et
al., 2007; Hagadorn and Belt, 2008), the presence of these trace fossils provides key
information for understanding the ecological roles that cnidarians played during the
Cambrian radiation, the environments they inhabited, and the infl uences that they had
upon the marine substrate. This study shows that the trace makers responsible for
Bergaueria, Conichnus, and Dolopichnus inhabited a tidally-infl uenced nearshore zone
including fl ood-tidal delta and lagoonal environments and that they were responsible for
extensive modifi cation of the marine substrate through deep burrowing activity. Further
reconciliation of these types of vertical trace fossils and their depositional context in
Cambrian strata may further emphasize the role that anemone-like organisms played in
the environmental and ecological changes that paralleled the Cambrian radiation.
187
CHAPTER VII
Conclusions
A revision of the sedimentology of the Precambrian-Cambrian transition of the
southern Great Basin and a new set of paleoenvironmental interpretations has allowed
for the development of new ideas surrounding the Cambrian radiation and earliest
Phanerozoic ecologic patterns. This paleoenvironmental revision suggests that mixed
carbonate-siliciclastic sedimentary rocks of the White-Into and Death Valley successions
fi t the depositional architectures of barrier island coastlines. This barrier island model
interprets that siliciclastic sedimentary rocks were deposited within shallow shelf and
shoreface environments, while carbonate rocks formed within back-barrier environments
including fl ood-tidal delta, lagoon, tidal channel, intertidal fl at, and supratidal fl at. This
contrasts with previous assessments that placed the locus of carbonate production within
offshore settings seaward of nearshore siliciclastic deposition. The possible inhibitor to a
barrier island interpretation in previous studies may have been the absence of conclusive
barrier island deposits. This absence, however, is now explained by the ‘self-destructive
prograding barrier island model’ that shows that the nature of barrier island progradation
implies that barrier island deposits—including backshore dune deposits—are not likely
to be preserved in the rock record. This is supported by evidence from Precambrian-
Cambrian successions of the southern Great Basin that reveal that siliciclastic facies can
be independently modeled as a shallow shelf and shoreface succession, while carbonate
facies can be modeled independently as back-barrier tidally-infl uenced environments.
The transition between an underlying siliciclastic succession and an overlying
carbonate succession is often marked by an erosional unconformity, implying that a
phase of erosion separates these conformable shallow shelf/shoreface and back-barrier
associations.
188
These siliciclastic-carbonate transitions are apparent within lower Cambrian
Grand Cycles that consist of a lower siliciclastic half-cycle and an upper carbonate half-
cycle. Siliciclastic-carbonate transitions of Grand Cycles of the southern Great Basin
refl ect a shift from shallow shelf and shoreface siliciclastic environments to back-barrier
carbonate environments. This transition occurs within the highstand systems tracts
and the maximum fl ooding surface occurs within the siliciclastic half-cycle. This is in
contrast to previous interpretations that placed the transition at the maximum fl ooding
surface.
The reinterpretation of Precambrian-Cambrian southern Great Basin strata also
has implications for the paleoenvironmental distribution and paleoecological patterns of
early Cambrian fossil organisms, most notably the archaeocyathans. Archaeocyathans
exhibit dramatic shifts in ecology within carbonate environments of Grand Cycle A,
which prograded across the ancient shelf from the Death Valley region to the White-Inyo
region. Archaeocyaths fi rst occur as transported fossil assemblages within carbonate
fl ood-tidal delta deposits of the upper member of the Wood Canyon Formation in the
Death Valley region that were deposited within a tidal inlet complex landward of a
siliciclastic shoreface. Temporally, archaeocyaths occur next as components of bioherms
within lower shoreface deposits of the Montenegro Member of the Campito Formation
that predate carbonate facies of Grand Cycle A in the White-Inyo region. The arrival of
carbonate facies of Grand Cycle A within the Lower Member of the Poleta Formation
marks the fi rst development of archaeocyathan reefs within the southern Great Basin.
This reef development is initiated atop uppermost lower shoreface deposits of the
Montenegro Member of the Campito Formation. While these different archaeocyath
ecologies might be attributed to onshore-offshore trends, the depositional environments
in which the archaeocyaths occur throughout these different formations are relatively
189
conservative and all represent nearshore carbonate environments above fair-weather wave
base found landward of a siliciclastic shoreface.
While the lower Cambrian rock record of the southern Great Basin is notable for
its fossil succession, it is also well known for its extensive development of microbial
mats in the form of wrinkle structures. The reevaluation of the paleoenvironmental
distribution of siliciclastic and carbonate sedimentary rocks reveals that most fossil-
rich rocks are found within the back-barrier and nearshore carbonate successions,
while wrinkle structures are found primarily within shallow shelf environments. This
distribution is converse to what is found throughout much of the Phanerozoic where
microbial mats were restricted primarily to marginal marine settings and fossil-rich
deposits extended across the shelves. In the southern Great Basin, wrinkle structures
occur exclusively within heterolithic deposits of siltstone and hummocky cross-stratifi ed
sandstone of the offshore transition, as well as within deposits of interbedded siltstone
and planar laminated sandstone of a mixed tidal fl at system. Wrinkle structures show a
preference for being cast on hummocky cross-stratifi ed or planar laminated sandstone
beds that are capped by less-resistant siltstone. The presence of these microbial mat
structures in shallow shelf settings indicates the absence of a surface mixed layer, which
is in accord with previous observations that bioturbation was very shallow-penetrating
(typically less than 5 cm deep) and not extensively developed during the early Cambrian
in siliciclastic shelf environments.
Early Cambrian tidal fl at environments of the southern Great Basin reveal a
different story of bioturbation and microbial mats. Deposits from the middle member
of the Wood Canyon formation show a close association between wrinkle structures and
deep-penetrating vertical burrows (up to 10 cm in depth) confi ned to the same bedding
plane. These burrows have funnel and mound surface morphologies on bedding surfaces
and can exhibit vertical cylindrical, downward tapering, and U-shaped tubes that connect
190
to these surface features. While these burrows are deep-penetrating, they do not produce
high levels of bioturbation (ii = 2). These low levels of bioturbation may have allowed
for the preservation of these microbial mat structures, in spite of the depth of burrow
penetration.
Similar deep-penetrating burrows (up to 30 cm in depth) can be found in the
upper member of the Wood Canyon Formation. Burrows are found within bidirectional
cross-stratifi ed oolite with reactivation surfaces and sandy dolostone of a fl ood-tidal delta
and lagoonal environment, respectively. Burrows are assignable to three ichnogenera—
Bergaueria, Conichnus, and Dolopichnus—all of which are interpreted as recording the
behavior of anemones. Bergaueria consists of shallow hemispherical or plug-shaped
burrows with a blunt to gently rounded termination and represent a resting or dwelling
trace. Conichnus is a downward tapering conical burrow that can exhibit annulations
along its length and often defl ects adjacent laminae within the host rock. Conichnus
represents a deep-penetrating dwelling trace that in some instances includes equilibration
or escape components. Dolopichnus consists of a cylindrical tube that terminates
downward at a basal circular to elliptical expansion. Dolopichnus refl ects the burrowing
and anchoring behavior of an anemone as it inserted is physa into the sediment and
expanded it to anchor itself as it retracted into the burrow. These burrows exhibit much
higher levels of bioturbation (ii = 4) than the deep-penetrating vertical burrows of the
middle member of the Wood Canyon Formation; however, they are found in a different
depositional environment.
From these observations the following conclusions can be reached: (1)
Precambrian-Cambrian mixed carbonate-siliciclastic sedimentary rocks of the southern
Great Basin refl ect the interaction of shallow shelf and shoreface siliciclastic sediments
with marginal marine carbonate sediments of a barrier island coastline; (2) siliciclastic
and carbonate successions are commonly separated by erosional unconformities
191
that refl ect the destruction of subaerial barrier island deposits—including backshore
dunes—during progradation; (3) lower Cambrian Grand Cycles refl ect shifts between
shallow shelf and shoreface siliciclastics and back-barrier carbonates; (4) the transition
between siliciclastic rocks and carbonate rocks refl ects a shift in lateral depositional
environments, rather than a complete change in the depositional system, and this
transition occurs within the highstand systems tract, well above the maximum fl ooding
surface; (6) the progradation events recorded by these Grand Cycles allow for assessment
of paleoecological patterns over short timescales; (7) lower Cambrian wrinkle structures
show a preference for occurring within heterolithic deposits, primarily of offshore
transition and mixed tidal fl at environments; (8) these microbial mat features refl ect the
absence of a surface mixed layer, which is in accord with previous observations that
early Cambrian bioturbation was shallow penetrating and not extensively developed in
siliciclastic shelf environments; (9) tidal fl at environments from the middle member of
the Wood Canyon Formation record deep-penetrating burrows associated with wrinkle
structures; however, bioturbation levels are generally low (ii = 2); (10) lagoonal
environments of the upper member of the Wood Canyon Formation record similar
deep-penetrating burrows that are attributable to anemones and exhibit high levels of
bioturbation; (11) the nature of the lower Cambrian bioturbation and microbial mat record
shows that nearshore environments were characterized by deep-penetrating burrows in
lagoonal and tidal fl at environments, while shallow shelf environments exhibit shallow-
penetrating levels of bioturbation and a prevalence of microbial mat structures; (12)
this onshore-offshore pattern may parallel the observation that evolutionary novelties
originate in onshore environments and expand or migrate offshore over time; (13)
metazoan taxa and deep-penetrating bioturbation dominated in the onshore position
throughout much of this interval, while microbial mat substrates and shallow penetrating
bioturbation were typical for shelf environments; (14) The coexistence of microbial mats
192
and metazoans in the earliest Cambrian appears to be due to environmental restriction of
microbial mats primarily to the offshore position and metazoans primarily to the onshore
position, a distribution quite different to that found in the modern ocean today.
193
REFERENCES
Aigner, T., Reineck, H.-E., 1982. Proximality trends in modern storm sands from the
Helgoland Bight (North Sea) and their implications for basin analysis: Senckenbergiana
Maritima, v. 14, p. 183-215.
Aitken, J.D., 1966. Middle Cambrian to Middle Ordovician cyclic sedimentation,
southern Rocky Mountains of Alberta: Bulletin of Canadian Petroleum Geology, v. 14, p.
405-441.
Aitken, J.D., 1978. Revised models for depositional Grand Cycles, Cambrian of the
southern Rocky Mountains, Canada: Bulletin of Canadian Petroleum Geology, v. 26, p.
515-542.
Alpert, S.P., 1973. Bergaueria Prantl (Cambrian and Ordovician), a probable actinian
trace fossil: Journal of Paleontology, v. 47, p. 919-924.
Alpert, S.P., Moore, J.N., 1975. Lower Cambrian trace fossil evidence for predation on
trilobites: Lethaia, v. 8, p. 223-230.
Amos, C.L., Li, M.Z., Choung, K.-S., 1996. Storm-generated, hummocky stratifi cation on
the outer-Scotian Shelf: Geo-Marine Letters, v. 16, p. 85-94.
Ansell, A.D., Trueman, E.R., 1968. The mechanism of burrowing in the Anemone,
Peachia hastate Gosse: Journal of experimental marine biology and ecology, v. 2, p. 124-
134.
Armbruster, C.K., Stone, G.W., Xu, J.P., 1995. Episodic atmospheric forcing and bayside
foreshore erosion: Santa Rosa Island, Florida: Gulf Coast Association of Geological
Societies Transactions, v. 45, p. 31-37.
Armin, R.A., Mayer, L., 1983. Subsidence analysis of the Cordilleran miogeocline:
Implications for timing of late Proterozoic rifting and amount of extension: Geology, v.
11, p. 702-705.
Arnott, R.W., 1993. Quasi-planar-laminated sandstone beds of the Lower Cretaceous
Bootlegger Member, north-central Montana: Evidence of combined-fl ow sedimentation:
Journal of Sedimentary Petrology, v. 63, p. 488-494.
194
Arnott, R.W.C., Hein, F.J., Pemberton, S.G., 1995. Infl uence of the ancestral Sweetgrass
Arch on sedimentation of the Lower Cretaceous Bootlegger Member, north-central
Montana: Journal of Sedimentary Research, v. B65, p. 222-234.
Arnott, R.W., Southard, J.B., 1990. Exploratory fl ow-duct experiments on combined-
fl ow bed confi gurations, and some implications for interpreting storm-event stratifi cation:
Journal of Sedimentary Petrology, v. 60, p. 211-219.
Ausich, W.I., Bottjer, D.J., 1982. Tiering in suspension-feeding communities on soft
substrata throughout the Phanerozoic: Science, v. 216, p. 173-174.
Babcock, L.E., Peng, S., 2007. Cambrian chronostratigraphy: Current state and future
plans: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 254, p. 62-66.
Bailey, J.V ., Corsetti, F.A., Bottjer, D.J., Marenco, K.N., 2006. Microbially-mediated
environmental infl uences on metazoan colonization of matground ecosystems: Evidence
from the Lower Cambrian Harkless Formation: PALAIOS, v. 21, p. 215-226.
Banerjee, S., Jeevankumar, S., 2005. Microbially originated wrinkle structures on
sandstone and their stratigraphic context: Palaeoproterozoic Koldaha Shale, central India:
Sedimentary Geology, v. 176, p. 211-224.
Barnes, J.J., Klein, G.deV ., 1975. Tidal deposits in the Zabriskie Quartzite (Cambrian),
eastern California and western Nevada. In: Ginsburg, R.N. (ed.), Tidal Deposits: A
Casebook of Recent Examples and Fossil Counterparts, p. 163-169.
Bathurst, R.G.C., 1967. Subtidal gelatinous mat, sand stabilizer and food, Great Bahama
Bank: The Journal of Geology, v. 75, p. 736-738.
Baucon, A., 2008. Neoichnology of a microbial mat in a temperate, siliciclastic
environment: Spiaggia al Bosco (Grado, Northern Adriatic, Italy): Studi Trent. Sci. Nat.
Acta Geol., v. 83, p. 183-203.
Berelson, W.M., Heron, S.D., 1985. Correlations between Holocene fl ood tidal delta
and barrier island fi ll sequences: Back Sound—Shackleford Banks, North Carolina:
Sedimentology, v. 32, p. 215-222.
Berger, W.H., Ekdale, A.A., Bryant, P.P., 1979. Selective preservation of burrows in deep-
sea carbonates: Marine Geology, v. 32, p. 205-230.
195
Bernard, H.A., LeBlanc, R.J., Major, C.F., 1962. Recent and Pleistocene geology of
southeast Texas, fi eld excursion no. 3. In: Rainwater, E.H., Zingula, R.P. (Eds.), Geology
of the Gulf Coast and central Texas, guidebook of excursions. Geological Society of
America Annual Meeting Guidebook, p. 175-224.
Bertics, V .J., Ziebis, W., 2009. Biodiversity of benthic microbial communities in
bioturbated coastal sediments is controlled by geochemical microniches: The ISME
Journal, v. 3, p. 1269-1285.
Bjerstedt, T.W., 1988. Trace fossils from the early Mississippian Price Delta, southeast
West Virginia: Journal of Paleontology, v. 62, p. 506-519.
Bond, G.C., Christie-Blick, N., Kominz, M.A., Devlin, W.J., 1985. An early Cambrian
rift to post-rift transition in the Cordillera of western North America: Nature, v. 315, p.
742-746.
Bottjer, D.J., Ausich, W.I., 1986. Phanerozoic development of tiering in soft substrata
suspension-feeding communities: Paleobiology, v. 12, p. 400-420.
Bottjer, D.J., Hagadorn, J.W., Dornbos, S.Q., 2000. The Cambrian substrate revolution:
GSA Today, v. 10, p. 1-7.
Bottjer, D.J., Schubert, J.K., Droser, M.L., 1996. Comparative evolutionary
palaeoecology; assessing the changing ecology of the past. In: Hart, M.B. (ed.), Biotic
recovery from mass extinction events, Geological Society of London, p. 1-13.
Boyd, D.W., 1966. Lamination deformed by burrowers in Flathead Sandstone (Middle
Cambrian) of central Wyoming: Wyoming University, Contributions to Geology, v. 5, p.
45-53.
Boyd, D.W., Ore, H.T., 1963. Patterned cones in the Permo-Triassic redbeds of Wyoming
and adjacent areas: Journal of Sedimentary Petrology, v. 33, p. 438-451.
Boyd, R., Penland, S., 1981. Washover of deltaic barriers on the Louisiana coast: Gulf
Coast Association of Geological Societies Transactions, v. 31, p. 243-248.
Bromley, R.G., 1996. Trace Fossils: Biology, Taphonomy, and Applications: Second
Edition, Chapman & Hall, London, 361 p.
Bromley, R.G., Curran, H.A., Frey, R.W., Gutschick, R.C., Suttner, L.J., 1975. Problems
in interpreting unusually large burrows. In: Frey, R.W. (ed.), The Study of Trace Fossils,
Springer-Verlag, New York, p. 351-376.
196
Buatois, L.A., Mángano, M.G., 2003. Early colonization of the deep sea: Ichnologic
evidence of deep-marine benthic ecology from the Early Cambrian of northwest
Argentina: PALAIOS, v. 18, p. 572-581.
Buatois, L.A., Mángano, M.G., 2012. An early Cambrian shallow-marine ichnofauna
from the Puncoviscana Formation of Northwest Argentina: The interplay between
sophisticated feeding behaviors, matgrounds and sea-level changes: Journal of
Paleontology, v. 86, p. 7-18.
Buck, S.G., Goldring, R., 2003. Conical sedimentary structures, trace fossils or not?
Observations, experiments, and review: PALAIOS, v. 73, p. 338-353.
Burne, R.V ., 1970. The origin and signifi cance of sand volcanoes in the Bude Formation
(Cornwall): Sedimentology, v. 15, p. 211-228.
Cady, S.L., Noffke, N., 2009. Geobiology: Evidence for early life on Earth and the search
for life on other planets: GSA Today, v. 19, p. 4-10.
Carey, D.A., 1987. Sedimentological effects and palaeoecological implications of the
tube-building polychaete Lanice conchilega Palla: Sedimentology, v. 34, p. 49-66.
Cartwright, P., Halgedahl, S.L., Hendricks, J.R., Jarrard, R.D., Marques, A.C., Collins,
A.G., Lieberman, B.S., 2007. Exceptionally preserved jellyfi shes from the Middle
Cambrian: PLoS ONE, v. 2, p. 1-7.
Chamberlain, C.K., 1971. Biogenic mounds in the Dakota Sandstone of northwestern
New Mexico: Journal of Paleontology, v. 45, p. 641-644.
Chow, N., James, N.P., 1987. Cambrian Grand Cycles: A northern Appalachian
perspective: Geological Society of America Bulletin, v. 98, p. 418-429.
Clapham, M.E., Corsetti, F.A., 2005. Deep valley incision in the terminal Neoproterozoic
(Ediacaran) Johnnie Formation, eastern California, USA: Tectonically or glacially
driven?: Precambrian Research, v. 141, p. 154-164.
Clifton, H.E., 1967. Solution-collapse and cavity fi lling in the Windsor Group, Nova
Scotia, Canada. Geological Society of America Bulletin, v. 78, p. 819-832.
Clifton, H.E., 1969. Beach lamination: nature and origin. Marine Geology, v. 7, p. 553-
559.
197
Clifton, H.W., Hunter, R.E., Phillips, R.L., 1971. Depositional structures and processes
in the non-barred high-energy nearshore: Journal of Sedimentary Petrology, v. 41, p. 651-
670.
Cloud, P., Wright, L.A., Williams, E.G., Diehl, P., Walter, M.R., 1974. Giant stromatolites
and associated vertical tubes from the upper Proterozoic Noonday Dolomite, Death
Valley region, eastern California: Geological Society of America Bulletin, v. 85, p. 1869-
1882.
Cook, D.O., Gorsline, D.S., 1972. Field observations of sand transport by shoaling
waves: Marine Geology, v. 13, p. 31-55.
Cornish, F.G., 1986. The trace-fossil Diplocraterion: evidence of animal-sediment
interactions in Cambrian tidal deposits: PALAIOS, v. 1, p. 478-491.
Cornwall, H.R., Kleinhampl, F.J., 1961. Geology of the Bare Mountain quadrangle,
Nevada: U.S. Geological Survey Quadrangle Map GQ-157, scale 1:62,500.
Corsetti, C.L., 2007. Early Cambrian vertical burrows from the upper member of the
Wood Canyon formation, Death Valley area, California: implications for infaunal
ecospace utilization: Unpublished M.S. thesis, University of California, Santa Barbara,
158 p.
Corsetti, F.A., Grotzinger, J.P., 2005. Origin and signifi cance of tube structures in
Neoproterozoic post-glacial cap carbonates: example from Noonday Dolomite, Death
Valley, United States: PALAIOS, v. 20, p. 348-362.
Corsetti, F.A., Hagadorn, J.W., 2000. Precambrian-Cambrian transition: Death Valley,
United States: Geology, v. 28, p. 299-302.
Corsetti, F.A., Hagadorn, J.W., 2003. The Precambrian-Cambrian transition in the
southern Great Basin, USA: The Sedimentary Record, v. 1, p. 4-8.
Corsetti, F.A., Kaufman, A.J., 2005. The relationship between the Neoproterozoic
Noonday Dolomite and the Ibex Formation: New observations and their bearing on
‘snowball Earth’: Earth-Science Reviews, v. 73, p. 63-78.
Crimes, T.P., 1992. Changes in the trace fossil biota across the Proterozoic-Phanerozoic
boundary: Journal of the Geological Society, London, v. 149, p. 637-646.
198
Crimes, T.P., Anderson, M.M., 1985. Trace fossils from Late Precambrian-Early
Cambrian strata of southeastern Newfoundland (Canada): Temporal and environmental
implications: Journal of Paleontology, v. 59, p. 321-343.
Crimes, T.P., Droser, M.L., 1992. Trace fossils and bioturbation: The other fossil record:
Annual Review of Ecology and Systematics, v. 23, p. 339-360.
Curran, H.A., and Frey, R.W., 1977, Pleistocene trace fossils from North Carolina
(U.S.A.), and their Holocene analogues. In: Crimes, T.P., Harper, J.C. (eds.), Trace Fossils
2, Seel House Press, Liverpool, p. 139-162.
Curran, H.A., Martin, A.J., 2003. Complex decapods burrows and ecological
relationships in modern and Pleistocene intertidal carbonate environments, San Salvador
Island, Bahamas: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 192, p. 229-
245.
Dade, W.B., Davis, J.D., Nichols, P.D., Nowell, A.R.M., Thistle, D., Trexler, M.B.,
White, D.C., 1990. Effects of bacterial exopolymer adhesion on the entrainment of sand:
Geomicrobiology Journal, v. 8, p. 1-16.
Dalrymple, R.W., Knight, R.J., Zaitlin, B.A., Middleton, G.V ., 1990. Dynamics and facies
model of a macrotidal sand-bar complex, Cobequid Bay—Salmon River Estuary (Bay of
Fundy): Sedimentology, v. 37, p. 577-612.
Davies, D.K., Ethridge, F.G., Berg, R.R., 1971. Recognition of barrier environments. The
American Association of Petroleum Geologists Bulletin, v. 55, p. 550-565.
Dawson, W.C., Reaser, D.F., 1994. Biogenic mounds and associated trace fossils: Wolfe
City Formation (Upper Cretaceous), north-central Texas: Transactions of the Gulf Coast
Association of Geological Societies, v. 44, p. 155-160.
Dean, W.E., Davies, G.R., Anderson, R.Y ., 1975. Sedimentological signifi cance of
nodular and laminated anhydrite: Geology, v. 3, p. 367-372.
Demicco, R.V ., 1983. Wavy and lenticular-bedded carbonate ribbon rocks of the Upper
Cambrian Conococheague Limestone, central Appalachians: Journal of Sedimentary
Petrology, v. 53, p. 1121-1132.
DePatra, K.D., Levin, L.A., 1989. Evidence of the passive deposition of meiofauna into
fi ddler crab burrows: Journal of Experimental Marine Biology and Ecology, v. 125, p.
173-192.
199
Diehl, P.E., 1974. Stratigraphy and sedimentology of the Wood Canyon Formation, Death
Valley area, California. In: Guidebook, Death Valley Region, California and Nevada.
Death Valley Publishing Co, Shoshone, p. 38-48.
Diehl, P.E., 1979. The stratigraphy, depositional environments, and quantitative
petrography of the Precambrian-Cambrian Wood Canyon Formation, Death Valley
region, California: Unpublished Ph.D. thesis, University Park, Pennsylvania State
University, 365 p.
Dill, R.F., Shinn, E.A., Jones, A.T., Kelly, K., Steinen, R.P., 1986. Giant subtidal
stromatolites forming in normal salinity waters: Nature, v. 324, p. 55-58.
Dionne, J.-C., 1976. Miniature mud volcanoes and other injection features in tidal fl ats,
James Bay, Quebec: Canadian Journal of Earth Sciences, v. 13, p. 422-428.
Donaldson, D., Simpson, S., 1962. Chomatichnus, a new ichnogenus, and other trace-
fossils of Wegber Quarry: Liverpool and Manchester Geological Journal, v. 3, p. 73-81.
Dornbos, S.Q., Bottjer, D.J., 2000. Evolutionary paleoecology of the earliest
echinoderms: Helicoplacoids and the Cambrian substrate revolution: Geology, v. 28, p.
839-842.
Dornbos, S.Q., Bottjer, D.J., 2001. Taphonomy and environmental distribution of
helicoplacoid echinoderms: PALAIOS, v. 16, p. 197-204.
Dornbos, S.Q., Noffke, N., Hagadorn, J.W., 2007. Mat-decay features. In: Schieber, J.,
Bose, P.K., Ericksson, P.G., Banerjee, S., Sarkar, S, Altermann, W., Catuneanu, O. (eds.),
Atlas of microbial mat features preserved within the siliciclastic rock record, Elsevier,
Amsterdam, p. 106-110.
Dott, R.H., 1983. 1982 SEPM Presidential Address: Episodic sedimentation—How
normal is average? How rare is rare? Does it matter?: Journal of Sedimentary Petrology,
v. 53, p. 5-23.
Dott, R.H., Bourgeois, J., 1982. Hummocky stratifi cation: Signifi cance of its variable
bedding sequences: Geological Society of America Bulletin, v. 92, p. 663-680.
Draganits, E., Noffke, N., 2004. Siliciclastic stromatolites and other microbially
induced sedimentary structures in an Early Devonian barrier-island environment (Muth
Formation, NW Himalayas): Journal of Sedimentary Research, v. 74, p. 191-202.
Dravis, J.J., 1983. Hardened subtidal stromatolites, Bahamas: Science, v. 219, p. 385-386.
200
Droser, M.L., 1991. Ichnofabric of the Paleozoic Skolithos ichnofacies and the nature and
distribution of Skolithos piperock: PALAIOS, v. 6, p. 316-325.
Droser, M.L., Bottjer, D.J., 1986. A semiquantitative fi eld classifi cation of ichnofabric:
Journal of Sedimentary Research, v. 56, p. 558-559.
Droser, M.L., Bottjer, D.J., 1988. Trends in depth and extent of bioturbation in Cambrian
carbonate marine environments, western United States: Geology, v. 16, p. 233-236.
Droser, M.L., Bottjer, D.J., 1989. Ordovician increase in extent and depth of bioturbation:
Implications for understanding early Paleozoic ecospace utilization: Geology, v. 17, p.
850-852.
Droser, M.L., Bottjer, D.J., 1993. Trends and patterns of Phanerozoic ichnofabrics:
Annual Review of Earth and Planetary Science, v. 21, p. 205-225.
Droser, M.L., Jensen, S., Gehling, J.G., 2002. Trace fossils and substrates of the terminal
Proterozoic-Cambrian transition: Implications for the record of early bilaterians and
sediment mixing: PNAS, v. 99, p. 12572-12576.
Droser, M.L., Jensen, S., Gehling, J.G., 2004. Development of early Palaeozoic
ichnofabrics: evidence from shallow marine siliciclastics. In: McIlroy, D. (ed.), The
Application of Ichnology to Palaeoenvironmental and Stratigraphic Analysis, Geological
Society Special Publication 228, p. 383–396.
Duc, A.W., Tye, R.S., 1987. Evolution and stratigraphy of a regressive barrier/backbarrier
complex: Kiawah Island, South Carolina. Sedimentology, v. 34, p. 237-251.
Duke, W.L., 1990. Geostrophic circulation or shallow marine turbidity currents? The
dilemma of paleofl ow patterns in storm-infl uenced prograding shoreline systems: Journal
of Sedimentary Petrology, v. 60, p. 870-883.
Duke, W.L., Arnott, R.W.C., Cheel, R.J., 1991. Shelf sandstones and hummocky cross-
stratifi cation: New insights on a stormy debate: Geology, v. 19, p. 625-628.
Dumas, S., 2006. Origin of hummocky and swaley cross-stratifi cation—The controlling
infl uence of unidirectional current strength and aggradation rate: Geology, v. 34, p. 1073-
1076.
201
Duncan, P.B., 1987. Burrow structure and burrowing activity of the funnel-feeding
enteropneust Balanoglossus aurantiacus in Bogue Sound, North Carolina, USA: Marine
Ecology, v. 8, p. 75-95.
Dupraz, C., Visscher, P.T., 2005. Microbial lithifi cation in marine stromatolites and
hypersaline mats: TRENDS in Microbiology, v. 13, p. 429-438.
Fedo, C.M., Cooper, J.D., 1990. Braided fl uvial to marine transition: The basal Lower
Cambrian Wood Canyon Formation, southern Marble Mountains, Mojave Desert,
California: Journal of Sedimentary Petrology, v. 60, p. 220-234.
Fedo, C.M., Cooper, J.D., 2001. Sedimentology and sequence stratigraphy of
Neoproterozoic and Cambrian units across a craton-margin hinge zone, southeastern
California, and implications for the early evolution of the Cordilleran margin:
Sedimentary Geology, v. 141-142, p. 501-522.
Fenchel, T., 1998. Formation of laminated cyanobacterial mats in the absence of benthic
fauna: Aquatic Microbial Ecology, v. 14, p. 235-240.
Fenster, M.S., FitzGerald, D.M., 1996. Morphodynamics, stratigraphy, and sediment
transport patterns of the Kennebec River estuary, Maine, USA: Sedimentary Geology, v.
107, p. 99-120.
Fischer, A.G., 1961. Stratigraphic record of transgressing seas in light of sedimentation
on Atlantic coast of New Jersey: Bulletin of the American Association of Petroleum
Geologists, v. 45, p. 1656-1666.
Frey, R.W., 1970. Environmental signifi cance of recent marine lebensspuren near
Beaufort, North Carolina: Journal of Paleontology, v. 44, p. 507-519.
Frey, R.W., 1970. The lebensspuren of some common marine invertebrates near Beaufort,
North Carolina. II. Anemone burrows: Journal of Paleontology, v. 44, p. 308-311.
Frey, R.W., Howard, J.D., 1981. Conichnus and Schaubcylindrichnus: Redefi ned trace
fossils from the Upper Cretaceous of the Western Interior: Journal of Paleontology, v. 55,
p. 800-804.
Fritz, W.H., 1972. Lower Cambrian trilobites from the Sekwi Formation type section,
Mackenzie Mountains, northwestern Canada: Bulletin of the Geological Survey of
Canada, v. 212, p. 1-90.
202
Fritz, W.H., 1975. Broad correlations of some Lower and Middle Cambrian strata in the
North American Cordillera: Geological Survey of Canada Paper 75-1A, p. 145-159.
Fritz, W.H., 1993. New Lower Cambrian olenelloid trilobite genera Cirquella and
Geraldinella from southwestern Canada: Journal of Paleontology, v. 67, p. 856-868.
Fuller, D.R., 1972. Paleoenvironmental analysis of an Early Cambrian archaeocyathid
reef in the White-Inyo Mountains, California: Unpublished M.S. thesis, Idaho State
University, 53 p.
Genise, J.F., Poire, D.G., 2000. Fluidization in insect constructions in soils: Ichnos, v. 7,
p. 127-134.
Gerdes, G., Claes, M., Dunajtschik-Piewak, K., Riege, H., Krumbein, W.E., Reineck, H.-
E., 1993. Contribution of microbial mats to sedimentary surface structures: Facies, v. 29,
p. 61-74.
Gerdes, G., Klenke, T., Noffke, N., 2000. Microbial signatures in peritidal siliciclastic
sediments: a catalogue: Sedimentology, v. 47, p. 279-308.
Gill, W.D., Kuenen, P.H., 1957. Sand volcanoes on slumps in the Carboniferous of
Country Clare, Ireland: Quarterly Journal of the Geological Society, v. 113, p. 441-460.
Gingras, M.K., Bann, K.L., 2006. The bend justifi es the leans: Interpreting recumbent
ichnofabrics: Journal of Sedimentary Research, v. 76, p. 483-492.
Gingras, M.K., Hubbard, S.M., Pemberton, S.G., Saunders, T., 2000. The signifi cance of
Pleistocene Psilonichnus at Willapa Bay, Washington: PALAIOS, v. 15, p. 142-151.
Ginsburg, R.N., Lowenstam, H.A., 1958. The infl uence of marine bottom communities on
the depositional environment of sediments: The Journal of Geology, v. 66, p. 310-318.
Ginsburg, R.N., Planavsky, N.J., 2008. Diversity of Bahamian microbialite substrates.
In: Dilek, Y ., Furnes, H., Muehlenbachs, K. (eds), Links Between Geological Processes,
Microbial Activities & Evolution of Life, Springer, Berlin, p. 177-195.
Gonzalez, R., Eberli, G.P., 1997. Sediment transport and bedforms in a carbonate tidal
inlet; Lee Stocking Island, Exumas, Bahamas: Sedimentology, v. 44, p. 1015-1030.
Grant, S.W.F., 1990. Shell structure and distribution of Cloudina, a potential index fossil
for the terminal Proterozoic: American Journal of Science 290-A, 261-294.
203
Greenwood, B., Mittler, P.R., 1985. Vertical sequence and lateral transitions in the facies
of a barred nearshore environment: Journal of Sedimentary Petrology, v. 55, p. 366-375.
Greenwood, B., Sherman, D.J., 1986. Hummocky cross-stratifi cation in the surf zone:
fl ow parameters and bedding genesis: Sedimentology, v. 33, p. 33-45.
Griffi s, R.B., Suchanek, T.H., 1991. A model of burrow architecture and trophic modes in
thalassinidean shrimp (Decapoda: Thalassinidea): Marine Ecology Progress Series, v. 79,
p. 171-183.
Gutstadt, A.M., 1968. Petrology and depositional environments of the Beck Spring
Dolomite (Precambrian), Kingston Range, California: Journal of Sedimentary Petrology,
v. 38, p. 1280-1289.
Hagadorn, J.W., Belt, E.S., 2008. Stranded in upstate New York: Cambrian
scyphomedusae from the Potsdam Sandstone: PALAIOS, v. 23, p. 424-441.
Hagadorn, J.W., Bottjer, D.J., 1997. Wrinkle structures: Microbially mediated
sedimentary structures common in subtidal siliciclastic settings at the Proterozoic-
Phanerozoic transition: Geology, v. 25, p. 1047-1050.
Hagadorn, J.W., Bottjer, D.J., 1999. Restriction of a late Neoprotozoic biotope: Suspect-
microbial structures and trace fossils at the Vendian-Cambrian transition: PALAIOS, v.
14, p. 73-85.
Hagadorn, J.W., Dott, R.H., Damrow, D., 2002. Stranded on a Late Cambrian shoreline:
Medusae from central Wisconsin: Geology, v. 30, p. 147-150.
Hagadorn, J.W., McDowell, C., 2012. Microbial infl uence on erosion, grain transport and
bedform genesis in sandy substrates under unidirectional fl ow: Sedimentology, v. 59, p.
795-808.
Hagadorn, J.W., Schellenberg, S.A., Bottjer, D.J., 2000. Paleoecology of a large Early
Cambrian bioturbator: Lethaia, v. 33, p. 142-156.
Hagadorn, J.W., Waggoner, B., 2002. The Early Cambrian problematic fossil
Volborthella: new insights from the Basin and Range. In: Corsetti, F.A. (ed.), Proterozoic-
Cambrian of the Great Basin and Beyond: SEPM Pacifi c Section Guidebook, Santa Fe
Springs, p. 137-153.
Hallam, A., Swett, K., 1966. Trace fossils from the Lower Cambrian Pipe Rock of the
north-west Highlands: Scottish Journal of Geology, v. 2, p. 101-107.
204
Hammond, T.M., Collins, M.B., 1979. On the threshold of transport of sand-sized
sediment under the combined infl uence of unidirectional and oscillatory fl ow:
Sedimentology, v. 26, p. 795-812.
Hardie, L.A., Ginsburg, R.N., 1977. Layering: The origin and environmental signifi cance
of lamination and thin bedding. In: Hardie, L.A. (Ed.), Sedimentation on the modern
carbonate tidal fl ats of northwest Andros Islands, Bahamas, The John Hopkins University
Press, Baltimore, p. 50-123.
Harms, J.C., Southard, J.B., Spearing, D.R., Walker, R.G., 1975. Depositional
Environments as Interpreted from Primary Sedimentary Structures and Stratifi cation
Sequences: Society of Economic Paleontologists and Mineralogists, Short Course No. 2,
161 p.
Harwood, C.L., Sumner, D.Y ., 2011. Microbialites of the Neoproterozoic Beck Spring
Dolomite, southern California: Sedimentology, v. 58, p. 1648-1673.
Hayes, M.O., 1980. General morphology and sediment patterns in tidal inlets:
Sedimentary Geology, v. 26, p. 139-156.
Hazzard, J.C., 1937. Paleozoic section in the Nopah and Resting Springs Mountains, Inyo
County, California: California Journal of Mines and Geology, v. 33, p. 273-339.
Heaman, L.M., Grotzinger, J.P., 1992. 1.08 Ga diabase sills in the Pahrump Group,
California: Implications for development of the Cordilleran miogeocline: Geology, v. 20,
p. 637-640.
Hennessy, J.T., and Zarillo, G.A., 1987. The interrelation and distinction between fl ood-
tidal delta and washover deposits in a transgressive barrier island: Marine Geology, v. 78,
p. 35-56.
Héquette, A., Hill, P.R., 1993. Storm-generated currents and offshore sediment transport
on a sandy shoreface, Tibjak Beach, Canadian Beaufort Sea: Marine Geology, v. 113, p.
283-304.
Hill, P.R., Meulé, S., Longuépée, H., 2003. Combined-fl ow processes and sedimentary
structures on the shoreface of the wave-dominated Grande-Rivière-La-Baleine Delta:
Journal of Sedimentary Research, v. 73, p. 217-226.
205
Hiscott, R.N., James, N.P., Pemberton, S.G., 1984. Sedimentology and ichnology of
the Lower Cambrian Bradore Formation, coastal Labrador: Fluvial to shallow-marine
transgressive sequence: Bulletin of Canadian Petroleum Geology, v. 32, p. 11-26.
Hollingsworth, J.S., 2005. The earliest occurrence of trilobites and brachiopods in the
Cambrian of Laurentia: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 220, p.
153-165.
Holohan, B.A., Klos, E.G., Oviatt, C.A., 1998. Population density, prey selection,
and predator avoidance of the burrowing anemone (Ceriantheopsis americanus) in
Narragansett Bay, Rhode Island: Estuaries, v. 21, p. 466-469.
Horodyski, R.J., Bloeser, B., V onder Haar, S., 1977. Laminated algal mats from a coastal
lagoon, Laguna Mormona, Baja California, Mexico: Journal of Sedimentary Petrology, v.
47, p. 680-696.
Houser, C., Hapke, C., Hamilton, S., 2008. Controls on coastal dune morphology,
shoreline erosion and barrier island response to extreme storms: Geomorphology, v. 100,
p. 223-240.
Howard, J.D., 1972. Trace fossils as criteria for recognizing shorelines in the stratigraphic
record. In: Rigby, J.K., Hamblin, W.K. (eds.), Recognition of Ancient Sedimentary
Environments, SEPM Special Publication 16, p. 215-225.
Howard, J.D., Frey, R.W., 1984. Characteristic trace fossils in nearshore to offshore
sequences, Upper Cretaceous of east-central Utah: Canadian Journal of Earth Sciences, v.
21, p. 200-219.
Hoyt, J.H., 1967. Barrier island formation: Geological Society of America Bulletin, v. 78,
p. 1125-1136.
Hoyt, J.H., Henry, V .J., 1967. Infl uence of island migration on barrier-island
sedimentation: Geological Society of America Bulletin, v. 78, p. 77-86.
Hubbard, D.K., Oertel, G., Nummedal, D., 1979. The role of waves and tidal currents in
the development of tidal-inlet sedimentary structures and sand body geometry: Examples
from North Carolina, South Carolina, and Georgia: Journal of Sedimentary Petrology, v.
49, p. 1073-1092.
Hunt, D.L., 1990. Trilobite faunas and biostratigraphy of the Lower Cambrian Wood
Canyon Formation, Death Valley Region, California: Unpublished M.S. thesis, Davis,
University of California, 150 p.
206
Hunter, R.E., Clifton, H.E., Phillips, R.L., 1979. Depositional processes, sedimentary
structures, and predicted vertical sequences in barred nearshore systems, southern Oregon
coast: Journal of Sedimentary Petrology, v. 49, p. 711-726.
Imperato, D.P., Sexton, W.J., Hayes, M.O., 1988. Stratigraphy and sediment
characteristics of a Mesotidal ebb-tidal delta, north Edisto inlet, South Carolina: Journal
of Sedimentary Petrology, v. 58, p. 950-958.
Jackson, N.L., 1995. Wind and waves: Infl uence of local and non-local waves on
mesoscale beach behavior in estuarine environments: Annals of the Association of
American Geographers, v. 85, p. 21-37.
Jablonski, D., Sepkoski, J.J., Bottjer, D.J., Sheehan, P.M., 1983. Onshore-offshore
patterns in the evolution of Phanerozoic shelf communities: Science, v. 222, p. 1123-
1125.
Javor, B.J., Castenholz, R.W., 1984. Invertebrate grazers of microbial mats, Laguna
Guerrero Negro, Mexico. In: Cohen Y , Castenholz RW, Halvorson HO (eds.) Microbial
Mats: Stromatolites, MBL Lectures in Biology 3, Alan R. Liss Inc., New York, p. 85-94.
Johnson, H.D., 1977. Sedimentation and water escape structures in some late
Precambrian shallow marine sandstones from Finnmark, North Norway: Sedimentology,
v. 24, p. 389-411.
Jørgensen, B.B., Gallardo, V .A., 1999. Thioploca spp.: fi lamentous sulfur bacteria with
nitrate vacuoles: FEMS Microbiology Ecology, v. 28, p. 301-313.
Kerr, R.S., 1977. Facies, diagenesis, and porosity development in a Lower Cretaceous
bank complex, Edwards Limestone, north-central Texas. In: Bebout, D.G., and Loucks,
R.G. (eds.), Cretaceous Carbonates of Texas and Mexico: Applications to Subsurface
Exploration, Bureau of Economic Geology, Austin, Texas, p. 216-233.
Klein, G.D., 1971. A sedimentary model for determining paleotidal range: Geological
Society of America Bulletin, v. 82, p. 2585-2592.
Klein, G.D., 1975. Paleotidal range sequences, Middle Member, Wood Canyon
Formation (Late Pre-Cambrian), eastern California and western Nevada. In: Ginsburg,
R.N. (ed.), Tidal Deposits: New York, Springer-Verlag, p. 171-178.
Kochel, R.C., Dolan, R., 1986. The role of overwash on a mid-Atlantic coast barrier
island: Journal of Geology, v. 94, p. 902-906.
207
Knopf, A., 1918. A geologic reconnaissance of the Inyo Range and the eastern slope of
the southern Sierra Nevada, California, with a section on the stratigraphy of the Inyo
Range by Edwin Kirk: U.S. Geological Survey Professional Paper 110, 130 p.
Kraft, J.C., John, C.J., 1979. Lateral and vertical facies relations of transgressive barrier:
The American Association of Petroleum Geologists Bulletin, v. 63, p. 2145-2163.
Kristensen, E., Aller, R.C., Aller, J.Y ., 1991. Oxic and anoxic decomposition of tubes
from the burrowing sea anemone Ceriantheopsis americanus: Implications for bulk
sediment carbon and nitrogen balance: Journal of Marine Research, v. 49, p. 589-617.
Kumar, N., Sanders, J.E., 1974. Inlet sequence: a vertical succession of sedimentary
structures and textures created by the lateral migration of tidal inlets: Sedimentology, v.
21, p. 491-532.
Kuwahara, V .S., Toda, T., Hamasaki, K., Kikuchi, T., Taguchi, S., 2000. Variability in the
relative penetration of ultraviolet radiation to photosynthetically available radiation in
temperate coastal waters, Japan: Journal of Oceanography, v. 56, p. 399-408.
Labotka, T.C., 1978. Geology of the Telescope Peak Quadrangle, California and late
Mesozoic regional metamorphism, Death Valley area, California: Unpublished Ph.D.
thesis, Pasadena, California Institute of Technology, 392 p.
Labotka, T.C., Albee, A.L., Lanphere, M.A., McDowell, S.D., 1980. Stratigraphy,
structure, and metamorphism in the central Panamint Mountains (Telescope Peak
quandrangle), Death Valley area, California: Summary: Geological Society of America
Bulletin, v. 91, p. 125-129.
Langille, G.B., 1974. Earliest Cambrian – latest Proterozoic ichnofossils and problematic
fossils from Inyo County, California: Unpublished Ph.D. thesis, Binghamton, State
University of New York, 194 p.
Lanphere, M.A., 1964. Geochronologic studies in the eastern Mojave Desert, California:
The Journal of Geology, v. 72, p. 381-399.
Leatherman, S.P., 1979. Barrier dune systems: A reassessment: Sedimentary Geology, v.
24, p. 1-16.
Leckie, D., 1985. The Lower Cretaceous Notikewin Member (Fort St. John Group),
northeastern British Columbia: A progradational barrier island system: Bulletin of
Canadian Petroleum Geology, v. 33, p. 39-51.
208
Leckie, D.A., Krystinik, L.F., 1989. Is there evidence for geostrophic currents preserved
in the sedimentary record of inner to middle-shelf deposits: Journal of Sedimentary
Petrology, v. 59, p. 862-870.
Leatherman, S.P., 1983. Barrier dynamics and landward migration with Holocene sea-
level rise: Nature, v. 301, p. 415-417.
Leatherman, S.P., Zaremba, R.E., 1987. Overwash and aeolian processes on a U.S.
northeast coast barrier: Sedimentary Geology, v. 52, p. 183-206.
Leckie, D., 1985. The Lower Cretaceous Notikewin Member (Fort St. John Group),
northeastern British Columbia: A progradational barrier island system: Bulletin of
Canadian Petroleum Geology, v. 33, p. 39-51.
Levy, M., Christie-Blick, N., 1989. Pre-Mesozoic palinspastic reconstruction of the
eastern Great Basin (western United States): Science, v. 245, p. 1454-1462.
Levy, M., Christie-Blick, N., 1991. Tectonic subsidence of the early Paleozoic passive
continental margin in eastern California and southern Nevada: Geological Society of
America Bulletin, v. 103, p. 1590-1606.
Lipps, J.H., Sylvester, A.G., 1968. The enigmatic Cambrian fossil Volborthella and its
occurrence in California: Journal of Paleontology, v. 42, p. 329-336.
Lock, B.E., Roberts, J.L., 1999. Sedimentological features within the Edwards Group of
west-central Texas: Gulf Coast Association of Geological Societies Transactions, v. 49, p.
310-320.
Logan, B.W., 1961. Cryptozoon and associate stromatolites from the Recent, Shark Bay,
Western Australia: The Journal of Geology, v. 69, p. 517-533.
Logan, B.W., Hoffman, P.F., Gebelein, C.D., 1974. Algal mats, cryptalgal fabrics,
and structures, Hamelin Pool, Western Australia: American Association of Petroleum
Geologists Memoirs, v. 22, p. 140-194.
MacEachern, J.A., Pemberton, S.G., 1992. Ichnological aspects of Cretaceous shoreface
successions and shoreface variability in the Western Interior Seaway of North America:
In: Pemberton, S.G. (ed.), Applications of Ichnology to Petroleum Exploration, SEPM
Core Workshop, v. 17, p. 57-84.
209
MacNaughton, R.B., Narbonne, G.M., 1999. Evolution and ecology of Neoproterozoic-
Lower Cambrian trace fossils, NW Canada: PALAIOS, v. 14, p. 91-115.
Maliva, R.G., Knoll, A.H., Simonson, B.M., 2005. Secular change in the Precambrian
silica cycle: Insights from chert petrology: Geological Society of America Bulletin, v.
117, p. 835-845.
Mángano, M.G., Buatois, L.A., 2004. Ichnology of Carboniferous tide-infl uenced
environments and tidal fl at variability in the North American Midcontinent. In: McIlroy,
D. (ed.), The Application of Ichnology to Palaeoenvironmental and Stratigraphic
Analysis, Geological Society, London, Special Publications 228, p. 157-178.
Männil, R.M., 1966. O vertikalnykh norkakh zaryvaniya v Ordovikskikh izvestnyakakh
Pribaltiki. In: Hecker, R.F. (eds.), Organizm i sreda v geologischeskom proshlom.
Akademiya Nauk SSSR. Paleontologicheskij Institut, p. 200-207.
Marenco, K.N., Bottjer, D.J., 2008. The importance of Planolites in the Cambrian
substrate revolution: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 258, p. 189-
199.
Mariscal, R.N., Conklin, E.J., Bigger, C.H., 1977. The ptychocyst, a major new category
of cnida used in tube construction by a cerianthid anemone: The Biological Bulletin, v.
152, p. 392-405.
Mata, S.A., Bottjer, D.J., 2009a. Development of Lower Triassic wrinkle structures:
Implications for the search for life on other planets: Astrobiology, v. 9, p. 895-906.
Mata, S.A., Bottjer, D.J., 2009b. The paleoenvironmental distribution of Phanerozoic
wrinkle structures: Earth-Science Reviews, v. 96, p. 181-195.
Mata, S.A., Bottjer, D.J., 2011. Microbes and mass extinctions: paleoenvironmental
distribution of microbialites during times of biotic crisis: Geobiology, v. 10, p. 3-24.
Mata, S.A., Bottjer, D.J., in review. The self-destructive prograding barrier island model:
Evidence from the Precambrian-Cambrian transition of the southern Great Basin, United
States: Sedimentary Geology.
Mata, S.A., Corsetti, C.L., Corsetti, F.A., Awramik, S.M., and Bottjer, D.J., in review.
Early Cambrian anemone burrows from the upper member of the Wood Canyon
Formation, Death Valley region, United States: Paleoecological and paleoenvironmental
signifi cance: PALAIOS.
210
Maxson, J.H., 1935. Pre-Cambric stratigraphy of the Inyo Range (abstract): Proceedings
of the Geological Society of America for 1934, p. 314.
Maynard, A.K., Suter, J.R., 1983. Regional variability of washover deposits on the south
Texas coast: Gulf Coast Association of Geological Societies Transactions, v. 33, p. 339-
346.
McKee, E.H., Gangloff, R.A., 1969. Stratigraphic distribution of archaeocyathids in
the Silber Peak Range and the White and Inyo Mountains, western Nevada and eastern
California: Journal of Paleontology, v. 43, p. 716-726.
McKee, E.H., Moila, R.J., 1962. Precambrian and Cambrian rocks of south-central
Esmeralda County, Nevada: American Journal of Science, v. 260, p. 530-538.
McIlroy, D., Logan, G.A., 1999. The impact of bioturbation on infaunal ecology and
evolution during the Proterozoic-Cambrian transition: PALAIOS, v. 14, p. 58-72.
McIlroy, D., Garton, M., 2004. A worm’s eye view of the Early Palaeozoic sea fl oor:
Geology Today, v. 20, p. 224-230.
Meadows, P.S., Tait, J., 1989. Modifi cation of sediment permeability and shear strength
by two burrowing invertebrates: Marine Biology, v. 101, p. 75-82.
Middleton, G.V ., 1961. Evaporite solution breccias from the Mississippian of southwest
Montana: Journal of Sedimentary Petrology, v. 31, p. 189-195.
Miller, M.F., Byers, C.W., 1984. Abundant and diverse early Paleozoic infauna indicated
by the stratigraphic record: Geology, v. 12, p. 40-43.
Miller, M.C., McCave, I.N., Komar, P.D., 1977. Threshold of sediment motion under
unidirectional currents: Sedimentology, v. 24, p. 507-527.
Moore, J.N., 1976. Depositional environments of the Lower Cambrian Poleta Formation
and its stratigraphic equivalents, California and Nevada: Brigham Young University
Geology Studies, v. 23, p. 23-38.
Morgan, N., 1976. The Montenegro bioherms: their paleoecology, relation to other
archaeocyathid bioherms and to Early Cambrian sedimentation in the White and Inyo
mountains, California. In: Moore, J.N., Fritsche, A.E. (eds.), Depositional Environments
of Lower Paleozoic Rocks in the White-Inyo Mountains, Inyo County, California: SEPM
Pacifi c Coast Paleogeography Field Guide 1, p. 13-17.
211
Morton, R.A., 2002. Factors controlling storm impacts on coastal barriers and beaches:
A preliminary basis for near real-time forecasting: Journal of Coastal Research, v. 18, p.
486-501.
Moslow, T.F., Tye, R.S., 1985. Recognition and characterization of Holocene tidal inlet
sequences: Marine Geology, v. 63, p. 129-151.
Mount, J.F., 1982. Storm-surge-ebb origin of hummocky cross-stratifi ed units of the
Andrews Mountain Member, Campito Formation (Lower Cambrian), White-Inyo
Mountains, eastern California: Journal of Sedimentary Petrology, v. 52, p. 941-958.
Mount, J.F., Bergk, 1998. Depositional sequence stratigraphy of Lower Cambrian Grand
Cycles, southern Great Basin, U.S.A.: International Geology Review, v. 40, p. 55-77.
Mount, J.F., Hunt, D.L., Greene, L.R., Dienger, J., 1991. Depositional systems,
biostratigraphy and sequence stratigraphy of Lower Cambrian Grand Cycles,
southwestern Great Basin. In: Cooper, J.D., Stevens, C.H. (eds.), Paleozoic
paleogeography of the Western United States—II: Bakersfi eld, CA, Pacifi c Section
SEPM, p. 209-226.
Mount, J.F., Rowland, S.M., 1981. Grand Cycle A (Lower Cambrian) of the southern
Great Basin: A product of differential rates of sea level rise. In: Taylor, M.E. (ed.),
Short papers for the Second International Symposium on the Cambrian System: U.S.
Geological Survey Open File Report 81-743, p. 143-146.
Mount, J.F., Signor, P.W., 1985. Early Cambrian innovation in shallow subtidal
environments: Paleoenvironments of Early Cambrian shelly fossils: Geology, v. 13, p.
730-733.
Murray, S.P., 1970. Bottom currents near the coast during Hurricane Camille: Journal of
Geophysical Research, v. 75, p. 4579-4582.
Narbonne, G.M., Myrow, P.M., Landing, E., Anderson, M.M., 1987. A candidate
stratotype for the Precambrian-Cambrian boundary, Fortune Head, Burin Peninsula,
southeastern Newfoundland: Canadian Journal of Earth Sciences, v. 24, p. 1277-1293.
Nelson, C.A., 1962. Lower Cambrian-Precambrian succession, White-Inyo Mountains,
California: Bulletin of the Geological Society of America, v. 73, p. 139-144.
Nelson, C.A., 1976. Late Precambrian-Early Cambrian stratigraphic and faunal
succession of eastern California and the Precambrian-Cambrian boundary: Pacifi c Coast
paleogeography fi eld guide 1, SEPM, p. 31-42.
212
Nelson, C.A., 1978. Late Precambrian-Early Cambrian stratigraphic and faunal
succession of eastern California and the Precambrian-Cambrian boundary: Geological
Magazine, v. 115, p. 121-126.
Nelson, C.H., 1982. Modern shallow-water graded sand layers from storm surges, Bering
Shelf: A mimic of Bouma sequences and turbidite systems: Journal of Sedimentary
Petrology, v. 52, p. 537-545.
Nelson, J.R., Eckman, J.E., Robertson, C.Y ., Marinelli, R.L., Jahnke, R.A., 1999. Benthic
microalgal biomass and irradiance at the seafl oor on the continental shelf of the South
Atlantic Bight: Spatial and temporal variability and storm effects: Continental Shelf
Research, v. 19, p. 477-505.
Neumann, A.C., Gebelein, C.D., Scoffi n, T.P., 1970. The composition, structure and
erodability of subtidal mats, Abaco, Bahamas: Journal of Sedimentary Petrology, v. 40, p.
274-297.
Neumann-Mahlkau, P., 1976. Recent sand volcanoes in the sand of a dike under
construction: Sedimentology, v. 23, p. 421-425.
Niedoroda, A.W., Swift, D.J.P., Figueiredo, A.G., Freeland, G.L., 1985. Barrier island
evolution, middle Atlantic shelf, U.S.A. Part II: Evidence from the shelf fl oor: Marine
Geology, v. 63, p. 363-396.
Noffke, N., 1998. Multidirected ripple marks rising from biological and sedimentological
processes in modern lower supratidal deposits (Mellum Island, southern North Sea):
Geology, v. 26, p. 879-882.
Noffke, N., 1999. Erosional remnants and pockets evolving from biotic-physical
interactions in a Recent lower supratidal environment: Sedimentary Geology, v. 123, p.
175-181.
Noffke, N., 2000. Extensive microbial mats and their infl uences on the erosional and
depositional dynamics of a siliciclastic cold water environment (Lower Arenigian,
Montagne Noire, France): Sedimentary Geology, v. 136, p. 207-215.
Noffke, N., 2009. Turbulent lifestyle: Microbial mats on Earth’s sandy beaches—Today
and 3 billion years ago: GSA Today, v. 18, p. 4-9.
213
Noffke, N., Eriksson, K.A., Hazen, R.M., Simpson, E.L., 2006. A new window into Early
Archean life: Microbial mats in Earth’s oldest siliciclastic tidal deposit (3.2 Ga Moodies
Group, South Africa): Geology, v. 34, p. 253-256.
Noffke, N., Gerdes, G., Klenke, T., Krumbein, W.E., 2001. Microbially induced
sedimentary structures—a new category within the classifi cation of primary sedimentary
structures: Journal of Sedimentary Research, v. 71, p. 649-656.
Noffke, N., Hazen, R., Nhleko, N., 2003. Earth’s earliest microbial mats in a siliciclastic
marine environment (2.9 Ga Mozaan Group, South Africa): Geology, v. 31, p. 673-676.
Noffke, N., Knoll, A.H., Grotzinger, J.P., 2002. Sedimentary controls on the formation
and preservation of microbial mats in siliciclastic deposits: A case study from the upper
Neoproterozoic Nama Group, Namibia: PALAIOS, v. 17, p. 533-544.
Nolan, T.B., 1929. Notes on the stratigraphy and structure of the northwest portion of
Spring Mountain, Nevada: American Journal of Science, v. 217, p. 461-472.
Oertel, G.F., 1973. Examination of textures and structures of mud in layered sediments at
the entrance of a Georgia tidal inlet: Journal of Sedimentary Petrology, v. 43, p. 33-41.
Pacze śna, J., 2010. Ichnological record of the activity of Anthozoa in the Early Cambrian
succession of the Upper Silesian Block (southern Poland): Acta Geologica Polonica, v.
60, p. 93-103.
Paerl, H.W., Pinckney, J.L., 1996. A mini-review of microbial consortia: Their roles in
aquatic production and biogeochemical cycling: Microbial Ecology, v. 31, p. 225-247.
Palmer, A.R., Halley, R.B., 1979. Physical stratigraphy and trilobite biostratigraphy of
the Carrara Formation (Lower and Middle Cambrian) in the southern Great Basin: U.S.
Geological Survey Professional Paper 1047, 131 p.
Palmer, A.R., Repina, L.D., 1993. Through a glass darkly: Taxonomy, phylogeny, and
biostratigraphy of the Olenellina: University of Kansas Paleontological Contributions, v.
3, 35 p.
Paphitis, D., Velegrakis, A.F., Collins, M.B., Muirhead, A., 2001. Laboratory
investigations into the threshold of movement of natural sand-sized sediments under
unidirectional, oscillatory and combined fl ows: Sedimentology, v. 48, p. 645-659.
Peach, B.N., Horne, J., 1884. Report on the geology of the north-west of Sutherland:
Nature, v. 31, p. 31-34.
214
Pemberton, S.G., Flach, P.D., Mossop, G.D., 1982. Trace fossils from the Athabasca Oil
Sands, Alberta, Canada: Science, v. 217, p. 825-827.
Pemberton, S.G., Frey, R.W., Bromley, R.G., 1988. The ichnotaxonomy of Conostichus
and other plug-shaped ichnofossil: Canadian Journal of Earth Sciences, v. 25, p. 866-892.
Pemberton, S.G., Kobluk, D.R., 1978. Oldest known brachiopod burrow: the Lower
Cambrian of Labrador: Canadian Journal of Earth Sciences, v. 15, p. 1385-1389.
Pemberton, S.G., MacEachern, J.A., 1997. The ichnological signature of storm deposits:
The use of trace fossils in event stratigraphy: In: Brett, C.E. (ed.), Paleontological event
horizons: ecological and evolutionary implications, Columbia University Press, p. 73-
109.
Pemberton, S.G., Magwood, J.P.A., 1990. A unique occurrence of Bergaueria in the
Lower Cambrian Gog Group near Lake Louise, Alberta: Journal of Paleontology, v. 64, p.
436-440.
Pemberton, S.G., Risk, M.J., 1982. Middle Silurian trace fossils in the Hamilton, Ontario
region: Their identifi cation, abundance, and signifi cance: Northeastern Geology, v. 4, p.
98-104.
Pemberton, S.G., Van Wagoner, J.C., Wach, G.D., 1992. Ichnofacies of a wave-dominated
shoreline: In: Pemberton, S.G. (ed.), Applications of Ichnology to Petroleum Exploration,
SEPM Core Workshop, v. 17, p. 339-382.
Peng, S., Babcock, L., 2008. Cambrian Period. In: Ogg, J.G., Ogg, G., Gradstein, F.M.
(eds.), The Concise Geologic Time Scale, Cambridge University Press, Cambridge, p. 37-
46.
Penland, S., Boyd, R., 1981. Shoreline changes on the Louisiana barrier coast: IEEE
Oceans, v. 81, p. 209-219.
Pfl üger, F., 1999. Matground structures and redox facies: PALAIOS, v. 14, p. 25-39.
Picha, F., Gibson, R.I., 1985. Cordilleran hingeline: Late Precambrian rifted margin of the
North American craton and its impact on the depositional and structural history, Utah and
Nevada: Geology, v. 13, p. 465-468.
Pierre, C., Rouchy, J.M., 1988. Carbonate replacements after sulfate evaporites in the
middle Miocene of Egypt: Journal of Sedimentary Petrology, v. 58, p. 446-456.
215
Pilkey, O.H., Neal, W.J., Monteiro, J.H., Dias, J.M.A., 1989. Algarve Barrier Islands: A
noncoastal-plain system in Portugal: Journal of Coastal Research, v. 5, p. 239-261
Plint, A.G., 1984. A regressive coastal sequence from the upper Eocene of Hampshire,
southern England: Sedimentology, v. 31, p. 213-225.
Plint, A.G., Walker, R.G., 1987. Cardium Formation 8. Facies and environments of the
Cardium shoreline and coastal plain in the Kakwa Field and adjacent areas, northwestern
Alberta: Bulletin of Canadian Petroleum Geology, v. 35, p. 48-64.
Prantl, F., 1945. Two new problematic trails from the Ordovician of Bohemia: Académie
tchéque des sciences, Bulletin international, Classe des sciences mathematiques et
naturelles et de la médecine, v. 46, p. 49-59.
Prave, A.R., 1992. Depositional and sequence stratigraphic framework of the Lower
Cambrian Zabriskie Quartzite: Implications for regional correlations and the Early
Cambrian paleogeography of the Death Valley region of California and Nevada:
Geological Society of America Bulletin, v. 104, p. 505-515.
Prave, A.R., 1999. Two diamictites, two cap carbonates, two δ13C excursions, two rifts:
The Neoproterozoic Kingston Peak Formation, Death Valley, California: Geology, v. 27,
p. 339-342.
Pruss, S.B., Corsetti, F.A., Fischer, W.W., 2008. Seafl oor-precipitated carbonate fans in
the Neoproterozoic Rainstorm Member, Johnnie Formation, Death Valley region, USA:
Sedimentary Geology, v. 207, p. 34-40.
Pryor, W.A., 1975. Biogenic sedimentation and alteration of argillaceous sediments in
shallow marine environments: Geological Society of America Bulletin, v. 86, p. 1244-
1254.
Rankey, E.C., Riegl, B., Steffen, K., 2006. Form, function and feedbacks in a tidally
dominated ooid shoal, Bahamas: Sedimentology, v. 53, p. 1191-1210.
Reddering, J.S.V ., 1983. An inlet sequence produced by migration of a small microtidal
inlet against longshore drift: the Keurbooms Inlet, South Africa: Sedimentology, v. 30, p.
201-218.
Reineck, H.-E., 1967. Layered sediments of tidal fl ats, beaches and shelf bottoms
of the North Sea. In: Lauff, G.H. (Ed.), Estuaries, 83. American Association for the
Advancement of Science, p. 191–206.
216
Reineck, H.-E., 1968. Classifi cation and origin of fl aser and lenticular bedding:
Sedimentology, v. 11, p. 99-104.
Reineck, H.-E., Singh, I.B., 1972. Genesis of laminated sand and graded rhythmites in
storm-sand layers of shelf mud: Sedimentology, v. 18, p. 123-128.
Reinson, G.E., 1979. Barrier island systems. In: Walker, R.G. (Ed.), Facies Models.
Geoscience Canada, p. 51-68.
Reinson, G.E., 1984. Barrier-island and associated strand-plain systems. In: Reading,
H.G. (ed.), Facies Models, Geoscience Canada Reprint Series 1, Geological Association
of Canada, p. 119-140.
Ricketts, E.F., Calvin, J., Hedgpath, J.W., 1985. Between Pacifi c Tides: Fifth Edition,
Stanford University Press, Stanford, 652 p.
Riley, C.M., Byrne, J.V ., 1961. Genesis of primary structures in anhydrite: Journal of
Sedimentary Petrology, v. 31, p. 553-559.
Rowden, A.A., Jones, M.B., 1995. The burrow structure of the mud shrimp Callianassa
subterranean (Decapoda: Thalassinidea) from the North Sea: Journal of Natural History,
v. 29, p. 1155-1165.
Rowland, S.M., 1978. Environmental stratigraphy of the Lower Member of the Poleta
Formation (Lower Cambrian), Esmeralda County, Nevada: Unpublished Ph.D. thesis,
University of California, Santa Cruz, 116 p.
Rowland, S.M., 1984. Were there framework reefs in the Cambrian?: Geology, v. 12, p.
181-183.
Rowland, S.M., Gangloff, R.A., 1988. Structure and paleoecology of Lower Cambrian
reefs: PALAIOS, v. 3, p. 111-135.
Ryer, T.A., 1977. Patterns of Cretaceous shallow-marine sedimentation, Coalville and
Rockport areas, Utah: Geological Society of America Bulletin, v. 88, p. 177-188.
Saito, Y ., 1989. Modern storm deposits in the inner shelf and their recurrence intervals,
Sendai Bay, northeast Japan. In: Taira, A., Masuda, F. (eds.), Sedimentary Facies in the
Active Plate Margin, TERRAPUB, Tokyo, p. 331-344.
217
Savrda, C.E., 2002. Equilibrium response refl ected in a large Conichnus (Upper
Cretaceous Eutaw Formation, Alabama, USA): Ichnos, v. 9, p. 33-40.
Schäfer, W., 1962. Ecology and Palaeoecology of Marine Environments: Chicago,
University of Chicago Press, 568 p.
Schieber, J., 1999. Microbial mats in terrigenous clastics: The challenge of identifi cation
in the rock record: PALAIOS, v. 14, p. 3-12.
Schubert, J.K., Bottjer, D.J., 1992. Early Triassic stromatolites as post-mass extinctions
disaster forms: Geology, v. 20, p. 883-886.
Seilacher, A., 1999. Biomat-related lifestyles in the Precambrian: PALAIOS, v. 14, p. 86-
93.
Seilacher, A., 2007. Trace Fossil Analysis: Springer-Verlag, Berlin, 226 p.
Seilacher, A., Buatois, L.A., Mángano, M.G., 2005. Trace fossils in the Ediacaran-
Cambrian transition: Behavioral diversifi cation, ecological turnover and environmental
shift: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 227, p. 323-356.
Seilacher, A., Pfl üger, F., 1994. From biomats to benthic agriculture: A prehistoric
revolution. In: Krumbein, W.E., Paterson, D.M., Stal, L.J. (eds.), Biostabilization of
Sediments, Bibliotheks
und Informationssystem der Carl von Ossietzky Universität Oldenburg, p. 97–105.
Shanley, K.W., McCabe, P.J., Hettinger, R.D., 1992. Tidal infl uence in Cretaceous fl uvial
strata from Utah, USA: a key to sequence stratigraphic interpretation: Sedimentology, v.
39, p. 905-930.
Shinn, E.A., 1968. Burrowing in recent lime sediments of Florida and the Bahamas:
Journal of Paleontology, v. 42, p. 879-894.
Shinn, E.A., Lloyd, R.M., Ginsburg, R.N., 1969. Anatomy of a modern carbonate tidal-
fl at, Andros Island, Bahamas: Journal of Sedimentary Petrology, v. 39, p. 1202-1228.
Signor, P. W., 1994. Proterozoic-Cambrian boundary trace fossils: Biostratigraphic
signifi cance of Harlaniella, in the Lower Cambrian Wood Canyon Formation, Death
Valley, California: Bulletin of the New York State Museum, v. 481, p. 317-322.
Signor, P.W., Ryan, D.A., 1993. Lower Cambrian fossil Volborthella: The whole truth or
just a piece of the beast?: Geology, v. 21, p. 805-808.
218
Soutar, A., Crill, P.A., 1977. Sedimentation and climatic patterns in the Santa Barbara
Basin during the 19th and 20th century: Geological Society of America Bulletin, v. 88, p.
1161-1172.
Stamhuis, E.J., Schreurs, C.E., Videler, J.J., 1997. Burrow architecture and turbative
activity of the thalassinid shrimp Callianassa subterranean from the central North Sea:
Marine Ecology Press Series, v. 151, p. 155-163.
Stephenson, T., 1935. The British Sea Anemones, V ol. 2: Ray Society, London, 426 p.
Stewart, J.H., 1966. Correlation of Lower Cambrian and some Precambrian strata in the
southern Great Basin, California and Nevada: U.S. Geological Survey Professional Paper
550-C, p. C66-C72.
Stewart, J.H., 1970. Upper Precambrian and Lower Cambrian strata in the southern Great
Basin, California and Nevada: U.S. Geological Survey Professional Paper 620, 206 p.
Stewart, J.H., 1972. Initial deposits in the Cordilleran Geosyncline: Evidence of a late
Precambrian (<850 m.y.) continental separation: Geological Society of America Bulletin,
v. 83, p. 1345-1360.
Stewart, J.H., 1982. Regional relations of Proterozoic Z and Lower Cambrian rocks in
the western United States and northern Mexico: In: Cooper, J.D., Troxel, B.W., Wright,
L.A. (eds.), Geology of selected areas in the San Bernardino Mountains, western Mojave
Desert, and southern Great Basin, California, Geological Society of America annual
meeting, Cordilleran section fi eld trip no. 9, p. 171-181.
Stewart, J.H., Poole, F.G., 1974. Lower Paleozoic and uppermost Precambrian
Cordilleran Miogeocline, Great Basin, western United States. SEPM Special Publication
22, p. 28-57.
Stewart, JH, and Suczek, CA, 1977. Cambrian and latest Precambrian paleogeography
and
tectonics in the western United States. In: Stewart, J. H., Stevens, C. H., Fritsche, A. E.
(eds.), Paleozoic paleogeography of the western United States: Pacifi c Section SEPM
Symposium 1, p. 1-17.
Stone, G.W., Liu, B., Pepper, D.A., Wang, P., 2004. The importance of extratropical and
tropical cyclones on the short-term evolution of barrier islands along the northern Gulf of
Mexico, USA: Marine Geology, v. 210, p. 63-78.
219
Summa, C.L., 1993. Sedimentologic, stratigraphic, and tectonic controls on a mixed
carbonate-siliciclastic succession: Neoproterozoic Johnnie Formation, southeast
California: Unpublisehed Ph.D. thesis, Massachusetts Institute of Technology, 615 p.
Swinbanks, D.D., 1981. Sediment reworking and the biogenic formation of clay laminae
by Abarenicola pacifi ca: Journal of Sedimentary Petrology, v. 51, p. 1137-1145.
Tucker, M.E., 1976. Replaced evaporites from the late Precambrian of Finnmark, Arctic
Norway: Sedimentary Geology, v. 16, p. 193-204.
V ossler, S.M., Pemberton, S.G., 1988. Skolithos in the Upper Cretaceous Cardium
Formation: an ichnofossil example of opportunistic ecology: Lethaia, v. 21. P. 351-362.
Walcott, C.D., 1908. Cambrian geology and paleontology No. 5, Cambrian sections in the
Cordilleran area: Smithsonian Miscellaneous Collections, v. 53, p. 166– 230.
Walker, R. G., 1985. Geological evidence for storm transportation and deposition on
ancient shelves: In: Tillman, R. W., Swift, D.P.J., Walker, R. G. (eds.), Shelf sands and
Sandstone Reservoirs: Society of Economic Paleontologists and Mineralogists Short
Course Notes, v. 13, p. 243-302.
Wasserburg, G.J., Wetherill, G.W., Wright, L.A., 1959. Ages in the Precambrian terrane
of Death Valley, California: Journal of Geology, v. 67, p. 702-708.
Wells, J.T., Adams, C.E., Park, Y .-A., Frankenberg, E.W., 1990. Morphology,
sedimentology and tidal channel processes on a high-tide-range mudfl at, west coast of
South Korea: Marine Geology, v. 95, p. 111-130.
Wheeler, H.E., 1948. Late Pre-Cambrian-Cambrian stratigraphic cross section through
southern Nevada: University of Nevada Bulletin, v. 42, p. 1-58.
Widdows, J., Blauw, A., Heip, C.H.R., Herman, P.M.J., Lucas, C.H., Middelburg, J.J.,
Schmidt, S., Brinsley, M.D., Twisk, F., Verbeek, H., 2004. Role of physical and biological
processes in sediment dynamics of a tidal fl at in Westerschelde Estuary, SW Netherlands:
Marine Ecology Progress Series, v. 274, p. 41-56.
Wright, L., Williams, E.G., Cloud, P., 1978. Algal and cryptalgal structures and platform
environments of the late pre-Phanerozoic Noonday Dolomite, eastern California:
Geological Society of America Bulletin, v. 89, p. 321-333.
220
Wunderlich, F., 1970. Genesis and environment of the “Nellenköpfchenschichten”
(lower Emsian, Rheinian Devon) at locus typicus in comparison with modern coastal
environment of the German Bay: Journal of Sedimentary Petrology, v. 40, p. 102-130.
Xue, Z., Feng, A., Yin, P., Xia, D., 2009. Coastal erosion induced by human activities: a
Northwest Bohai Sea case study: Journal of Coastal Research, v. 253, p. 723-733.
Young, H.R., 1979. Evidence of former evaporites in the Cambro-Ordovician Durness
Group, northwest Scotland: Sedimentary Geology, v. 22, p. 287-303.
Ziebis, W., Forster, S., Huettel, M., Jørgensen, B.B., 1996. Complex burrows of the mud
shrimp Callianassa truncata and their geochemical impact in the sea bed: Nature, v. 382,
p. 619-622.
Zonneveld, J.-P., Gingras, M.K., Pemberton, S.G., 2001. Trace fossil assemblages in
a Middle Triassic mixed siliciclastic-carbonate marginal marine depositional system,
British Columbia: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 166, p. 249-
276.
221
APPENDIX
Locality Information
Death Valley Region
Emigrant Pass (upper member, Wood Canyon Formation): From Baker, California, take
CA-127 N/Death Valley Road north for 48 miles. Turn right onto Old Spanish Trail
toward Tecopa, California. The Emigrant Pass locality is found approximately 13 miles
down Old Spanish Trail. Exposed at this locality on the north side of the road is the up-
per member of the Wood Canyon Formation, which is found on a low-lying hill adjacent
to the road and consists of dark brown sandstones and siltstones with a prominent tan
dolostone unit near the top of the hill. Beds dip to the northeast. The measured section
begins at N 35° 53.577 ′, W 116° 04.800 ′.
Chicago Pass (lower member, Wood Canyon Formation): From Baker, California, take
CA-127 N/Death Valley Road north for 56 miles. Turn right onto CA-178 E/Charles
Brown Hwy toward Pahrump, Nevada. The Chicago Pass locality is found approximate-
ly 15.7 miles down CA-178 on the right-hand side. The locality is marked by a low-lying
hill consisting of dark brown rocks with three prominent tan dolostone units that form
ledges. The measured section is through the lowermost dolostone unit and begins at N
36° 08.565 ′, W 116° 09.175 ′.
Southern Salt Spring Hills (middle member, Wood Canyon Formation): From Baker,
California, take CA-127 N/Death Valley Road north for approximately 28 miles.
Heading north there will be a paved turnout just north of a salt fl at south of the road; park
at this turnout. Exposed on the south side of the road is the middle and upper members of
the Wood Canyon Formation that are overlain by the Zabriskie Quartzite. Beds dip to the
northeast, and the middle member forms prominent dark brown dip-slopes on the hillside.
The measured section begins at N 35° 36.709 ′, W 116° 16.061 ′ and is found in the low-
lying saddle that separates the prominent peaks represented by the conglomeratic lower
interval of the middle member and the quartzites of the Zabriskie.
White-Inyo Region
Cedar Flat (Harkless Formation): Heading north on US-395 through Big Pine, California,
turn right onto CA-168 near the northernmost end of town. Travel east on CA-168 for
approximately 13 miles. At the top of the climb, when the road fl attens out, turn right
into the CARMA Observatory. Follow the CARMA dirt road toward the telescope array
and observatory building. Turn left onto the un-gated dirt road just before the building
and follow this road past the telescope array to the northeast and up a low-lying ridge
222
that trends northeast-southwest. Halfway up the ridge will be a picnic table; park here.
Follow the ridge further to the northeast until it leads down in elevation and fl attens
out before running into a dark brown peak with cliff exposures. These strata belong to
the Harkless Formation and dip to the northwest. The measured section contains the
uppermost portion of the peak and begins at N 37° 17.111 ′, W 118° 08.035 ′.
North of Cedar Flat (Campito and Poleta Formations): Heading north on US-395 through
Big Pine, California, turn right onto CA-168 near the northernmost end of town. Travel
east on CA-168 for approximately 14 miles. Turn right onto the dirt road on the east
side of the road at park. Follow CA-168 south for approximately 0.2 miles and head
down into the wash on the east side of the road. This wash leads northeast into a narrow
canyon. Follow this canyon for approximately 800 meters. Exposed on the east side
of the canyon are dark brown siltstones and sandstones of the Montenegro Member of
the Campito Formation overlain by gray limestones of the Lower Member of the Poleta
Formation. Beds dip to the east and the measured section begins at N 37° 18.070 ′, W
118° 08.834 ′ and contains the Campito-Poleta contact.
Hines Ridge (Hines Tongue, Reed Dolomite and Lower Member, Deep Spring Forma-
tion): Heading north on US-395 through Big Pine, California, turn right onto CA-168
near the northernmost end of town. Travel east on CA-168 for approximately 2.3 miles.
Turn right onto Death Valley Road/Waucoba Road and travel east for approximately 11
miles. Park along the south side of the road, approximately 0.4 miles west of the turn-
off for Hines Road. South of the road is exposed a northeast-southwest trending ridge.
Proceed down the wash that parallels the east side of Hines Ridge for approximately 2.5
kilometers. Up the side of the ridge to the west is the dark brown to tan sandstones of
the Hines Tongue of the Reed Dolomite. The beds dip to the southwest and the measured
section begins at (coordinates). Approximately 350 meters to the southwest is the basal
contact of the Deep Spring Formation consisting of dark brown siltstones and sandstones
overlying tan dolostones of the Reed Dolomite. The measured section of the Lower
Member of the Deep Spring Formation begins at N 37° 06.173 ′, W 118° 05.628 ′.
Poleta Folds (Poleta Formation): Heading north on US-395 through Big Pine, California,
turn right onto CA-168 near the northernmost end of town. Travel east on CA-168 for
approximately 18.5 miles. Just after passing the boundary of the National Forest, there is
a dirt road on the south side of the road that is blocked by an electric fence. Turn off onto
this dirt road and unhook the electric fence to pass through into a roundabout. Park just
off the road in this roundabout. Travel approximately 500 meters to the northeast along
the fl at before heading down a shallow and narrow valley into the main wash. On the
southeast side of the wash is exposed dark brown sandstone and siltstone of the Middle
Member of the Poleta Formation. Cross the wash and head up the slope made up by
the Middle Member. At the top of the wash is exposed the Upper Member of the Poleta
223
Formation, which is incised by a narrow canyon that cuts down through both the Middle
and Upper Members. Beds dip slightly to the southeast. The entirety of this canyon was
measured beginning at N 37° 18.877 ′, W 118° 05.444 ′. For off-road vehicles that can
handle dirt roads, a shorted route can be achieved by making a right turn off CA-168 at N
37° 19.105 ′, W 118° 05.489 ′ and following the dirt road passed the electric fence. Park at
N 37° 19.011 ′, W 118° 05.436 ′ and the canyon will be to the Southwest. The canyon can
be accessed at its bottom, or from the top by climbing the hill in which it resides.

Abstract (if available)

Abstract The Precambrian-Cambrian transition (~542 Ma) is a pivotal time in Earth’s history and is notable for marking the first skeletonized occurrences of many major groups of marine organisms. Concurrent with the appearance of these taxa are significant changes to the marine substrate, shifting from a seafloor dominated by microbial mats to one dominated by metazoan bioturbation. In the southern Great Basin, United States, these changes, however, appear gradual and the early Cambrian is unlike either the Precambrian or the subsequent Phanerozoic in terms of the unusual coexistence between seafloor microbial mats and marine organisms. Essential to understanding these interactions is a strong paleoenvironmental framework to work within to examine which environments microbial mats occur in and whether these are or are not the same environments that metazoan fossils and trace fossils occur in. The purpose of this dissertation is to develop a rigorous depositional model of Precambrian-Cambrian strata within the southern Great Basin to document the paleoenvironmental distribution of fossil taxa, trace fossils, and microbial mat structures to examine the nature of this coexistence during the early Cambrian. This may shed light on whether metazoans and microbial mats existed in similar environments or if each occupied a preferred environmental niche apart from the other. ❧ The Precambrian-Cambrian transition of the southern Great Basin, United States, consists of mixed-carbonate siliciclastic sedimentary rocks. Prior studies have emphasized that carbonate rocks were deposited offshore to siliciclastic rocks

Linked assets

University of Southern California Dissertations and Theses

Conceptually similar

PDF

Bioturbation in Cambrian siliciclastic shelf strata: paleoecological, paleoenvironmental, and temporal patterns

PDF

Benthic and pelagic marine ecology following the Triassic/Jurassic mass extinction

PDF

The geobiology of fluvial, lacustrine, and marginal marine carbonate microbialites (Pleistocene, Miocene, and Late Triassic) and their environmental significance

PDF

Community paleoecology and global diversity patterns during the end-Guadalupian extinction (middle-late Permian) and the transition from the Paleozoic to modern evolutionary faunas

PDF

Evolution of the Indian Monsoon and rise of C₄ photosynthesis in the Miocene and Pliocene

PDF

Ecological recovery dynamics of the benthic and pelagic fauna in response to extreme temperature events and low oxygen environments developed during the early Triassic

PDF

Benthic paleoecology and macroevolution during the Norian Stage of the Late Triassic

PDF

Stromatolites in the ancient and modern: new methods for solving old problems

PDF

Carbonate geochemistry in primary, diagenetic and biological systems

PDF

Paleoenvironmental and paleoecological trends leading up to the end-Triassic mass extinction event

PDF

The geobiological role of bioturbating ecosystem engineers during key evolutionary intervals in Earth history

PDF

Paleoecology of Upper Triassic reef ecosystems and their demise at the Triassic-Jurassic extinction, a potential ocean acidification event

PDF

Viral, archaeal and bacterial colonization of native marine rock substrata

PDF

The early Triassic recovery period: exploring ecology and evolution in marine benthic communities following the Permian-Triassic mass extinction

PDF

Preservation of gas-related textures in microbialites: Evidence for ancient metabolisms and environments

PDF

Great Salt Lake ooids: insights into rate of formation, potential as paleoenvironmental archives, and biogenicity

PDF

New perspectives on ancient microbes and microbialites: from isotopes to immunology

PDF

Annual pattern and response of the bacterial and microbial eukaryotic communities in an aquatic ecosystem restructured by disturbance

PDF

Evolution & ecology of Mesozoic birds: a case study of the derived Hesperornithiformes and the use of morphometric data in quantifying avian paleoecology

PDF

Evolution of upper crustal magmatic plumbing systems: comparisons of geochronological, petrological, geochemical and structural records in the Cretaceous Fangshan pluton, Beijing, China, the Meso...

Asset Metadata

Creator Mata, Scott Andrew (author)

Core Title Paleoenvironments and the Precambrian-Cambrian transition in the southern Great Basin: Implications for microbial mat development and the Cambrian radiation

School College of Letters, Arts and Sciences

Degree Doctor of Philosophy

Degree Program Geological Sciences

Publication Date 07/09/2012

Defense Date 04/26/2012

Publisher University of Southern California (original), University of Southern California. Libraries (digital)

Tag barrier island,Death Valley,OAI-PMH Harvest,sedimentology,stratigraphy,White-Inyo

Language English

Contributor Electronically uploaded by the author (provenance)

Advisor Bottjer, David J. (committee chair), Corsetti, Frank A. (committee member), Ziebis, Wiebke (committee member)

Creator Email scottamata@gmail.com

Permanent Link (DOI) https://doi.org/10.25549/usctheses-c3-53424

Unique identifier UC11289311

Identifier usctheses-c3-53424 (legacy record id)

Legacy Identifier etd-MataScottA-923.pdf

Dmrecord 53424

Document Type Dissertation

Rights Mata, Scott Andrew

Type texts

Source University of Southern California (contributing entity), University of Southern California Dissertations and Theses (collection)

Access Conditions The author retains rights to his/her dissertation, thesis or other graduate work according to U.S. copyright law. Electronic access is being provided by the USC Libraries in agreement with the a...

Repository Name University of Southern California Digital Library

Repository Location USC Digital Library, University of Southern California, University Park Campus MC 2810, 3434 South Grand Avenue, 2nd Floor, Los Angeles, California 90089-2810, USA