Carbonate geochemistry in primary, diagenetic and biological systems

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Content CARBONATE GEOCHEMISTRY IN PRIMARY, DIAGENETIC AND BIOLOGICAL
SYSTEMS

by

Sean Joseph Loyd

A Dissertation Presented to the
FACULTY OF THE USC GRADUATE SCHOOL
UNIVERSITY OF SOUTHERN CALIFORNIA
In Partial Fulfillment of the
Requirements for the Degree
DOCTOR OF PHILOSOPHY
(GEOLOGICAL SCIENCES)

August 2010

Copyright 2010 Sean Joseph Loyd
ii
ACKNOWLEDGEMENTS
This compilation of work would not have been possible (nor bearable!) without
the support and efforts of a multitude of people. Two individuals deserve special credit
above all others. The first is my graduate advisor Frank Corsetti who has overseen all of
my recent scientific endeavors with unfailing support and continued enthusiasm. I owe a
huge portion of my scientific success to him and can’t imagine a better scientific advisor
and colleague. The second is my dear fiancé (and soon to be wife!) Jennifer Bjelland
who has provided personal support in the form of love, devotion and most of all tolerance
during my development as a graduate student. There is not a single word I have written
in my tenure at USC that she has not read—she is my biggest fan and best critic. I also
owe a great deal of gratitude to my friends and family. I feel that individual achievement
can only occur with a strong personal support group. I would like to particularly
acknowledge my mother who has provided constant support regardless of circumstance.
Much of the included work would not have been possible without scientific efforts from
colleagues Pedro Marenco, Whitey Hagadorn, Will Berelson, Victoria Petryshyn, Jake
Bailey, Nate Lorentz, Dave Bottjer, Doug Hammond, Tim Lyons, Steven Bates, Chris
Evans, Miguel Rincon, Lowell Stott, Wiebke Ziebis, Jay Kaufman, Jim Boles and Jim
Kennett. Your past and continued support is greatly appreciated!

iii
TABLE OF CONTENTS
Acknowledgements ii
List of Tables iv
List of Figures v
Abstract ix
Chapter 1: Carbon and sulfur isotope chemostratigraphy of 1
Neoproterozoic to Cambrian carbonates from northwestern
Mexico and eastern California: Implications for marine sulfate
concentrations across the Precambrian-Cambrian boundary

Chapter 2: Local variability in ~580 Ma carbonates of the 37
Clemente Formation, Caborca, Mexico: Implications for
diagenesis and Neoproterozoic marine sulfate

Chapter 3: Carbonate-associated sulfate, δ
34
S and δ
13
C 80
analyses of dolomite concretions of the Miocene Monterey
Formation: Insights into formation environments

Chapter 4: Carbon and sulfur isotopic compositions of calcitic 121
concretions of the Upper Cretaceous Holz Shale: Catching
sulfide oxidation in the act

Chapter 5: The origin of the millimeter-scale lamination in the 153
Neoproterozoic lower Beck Spring Dolomite: Implications for
widespread, fine-scale, layer-parallel diagenesis in Precambrian
carbonates

Bibliography 182

iv
LIST OF TABLES
Table 1.1: Stratigraphic characterization of the Caborca and 7
Death Valley Successions

Table 1.2: Compilation of geochemical data from this study 12
(Caborca and upper Death Valley Successions) and others
(lower Death Valley Succession)

Table 2.1: Marker bed geochemistry 45

Table 3.1: Monterey geochemical data listed according to 99
site and sample

Table 3.2: Precipitation, remineralization and 107
other diagenetic reactions

Table 4.1: Geochemical data from Holz Shale concretionary 130
structures

Table 5.1: Carbon and oxygen isotope values of the Beck 170
Spring Dolomite separated by sample

v
LIST OF FIGURES
Figure 1.1: Locality map of the Death Valley and Caborca 4
Neoproterozoic to middle Cambrian successions

Figure 1.2: Stratigraphic correlation between units of 5
Caborca and Death Valley

Figure 1.3: Photographs of key lithostratigraphic markers 10
from Caborca and Death Valley

Figure 1.4: Chemostratigraphic trends of Cerro Rajon, 19
Caborca, Mexico

Figure 1.5: Chemostratigraphic trends of the Death Valley 21
Succession, eastern California, USA

Figure 1.6: Transposed plots of the Caborca and Death 23
Valley δ
34
S
CAS
, δ
34
S
pyrite
and Δ
34
S data

Figure 1.7: CAS concentrations compared to lithology 26

Figure 1.8: CAS concentrations for each of the three intervals 28

Figure 1.9: Cross plot of δ
34
S
CAS
and pyrite concentration 31
Figure 2.1: Map of study sites and other Neoproterozoic 41
facies in the Caborca region

Figure 2.2: Stratigraphic column of Neoproterozoic units 43
from Cerro Rajon

Figure 2.3: Photographs of marker bed components 46

Figure 2.4: Photomicrographs of oolitic and laminated 48
components of the marker bed

Figure 2.5: Photomicrographs showing the variability in 50
textures exhibited by the marker bed

Figure 2.6: Geochemical data shown stratigraphically for 52
all five marker bed sections

vi
Figure 2.7: Cross plot of δ
13
C
carbonate
and δ
18
O
carbonate
for 59
all marker bed samples

Figure 2.8: Composite δ
13
C
carbonate
, δ
34
S
CAS
and CAS 61
concentration curves for the five Caborca study sites and
those of the Rainstorm carbonates (DV)

Figure 2.9: Assessing pyrite oxidation in marker bed samples 63

Figure 2.10: Cross plots of CAS and traditional proxies of 66
meteoric diagenesis

Figure 2.11: Cross plots of δ
34
S
CAS
and traditional proxies 68
of meteoric diagenesis

Figure 2.12: Diagrammatic evolution of the Caborca 72
region ca. 580 Ma

Figure 2.13: Global comparison of δ
34
S
CAS
data 77

Figure 3.1: Map showing Monterey sample site locations 82

Figure 3.2: Concretions and stratigraphic column of the 84
Monterey Formation

Figure 3.3: Diagram depicting pore water profiles of major 87
compounds in marine sediments and predicted values of
δ
13
C, δ
34
S
CAS
and CAS concentration of concretions
precipitated at different relative depths

Figure 3.4: Modern pore water profiles of sulfate 88
concentration and δ
34
S
sulfate

Figure 3.5A: Photographs of Montana de Oro concretions 90
MMC1-MMC5

Figure 3.5B: Photographs of Montana de Oro concretions 91
MMC6-MMC9

Figure 3.6: Cemented layer photographs and stratigraphic 93
column for the Shell Beach locality

vii
Figure 3.7: Photographs of the two concretions sampled 95
from Naples Beach

Figure 3.8: Photomicrographs of Montana de Oro, Shell 97
Beach, MNC1 and MNC2

Figure 3.9: Weight percent dolomite in Monterey samples 101

Figure 3.10: Monterey geochemistry cross plots 103

Figure 3.11: Cross plot of Monterey δ
18
O and δ
13
C 106

Figure 3.12: Monterey sulfur isotope cross plots 114

Figure 3.13: Assessing pyrite oxidation in Monterey samples 117

Figure 4.1: Map showing location of the Holz Shale outcrop 123
examined in this study

Figure 4.2: Photographs of concretionary structures and host 125
rock (A-D) and stratigraphic column (E) of the Holz Shale

Figure 4.3: Photomicrographs and SEM images of Holz 128
Shale concretionary structures

Figure 4.4: Photographs and associated sketches showing 132
sampling pattern

Figure 4.5: Cross plot of Holz Shale δ
13
C and δ
18
O 135

Figure 4.6: Assessing pyrite contamination in the Holz Shale 137

Figure 4.7: CAS acidification test for five Holz Shale samples 139

Figure 4.8: CAS concentrations of concretionary structures 141
of the Holz Shale

Figure 4.9: Cross plot of Holz Shale δ
13
C and δ
34
S
CAS
143

Figure 4.10: Diagram depicting the relative concentration of 145
dissolved pore water sulfate and sulfide and the evolution of
δ
34
S
pws
with increasing depth

Figure 4.11: Holz Shale box model parameters and results 148
viii

Figure 5.1: Map of the Alexander Hills and other outcrops 156
of Proterozoic to early Cambrian miogeoclinal units

Figure 5.2: Stratigraphic column of the Pahrump Group and 158
overlying units of the Death Valley succession

Figure 5.3: The lamination of the lower Beck Spring 160
Dolomite viewed at increasing magnification

Figure 5.4: Cut and polished slabs of the lower Beck 162
Spring Dolomite

Figure 5.5: Morphologic variations in light and dark 164
laminae of the Beck Spring Dolomite

Figure 5.6: Image of merged photomicrographs showing 166
lateral variability and continuity of light/dark lamination

Figure 5.7: Merged photomicrographs showing relationship 168
of cavity structures to light laminae

Figure 5.8: Photomicrographs of the abnormal, light 173
laminae of BS7

Figure 5.9: Carbon and oxygen isotope values for each 176
Beck Spring sample

ix
ABSTRACT
The carbonate minerals calcite, aragonite and dolomite (and their rock-
counterparts) precipitate directly from fluids. The mineral-yielding fluids must contain
the necessary chemical constituents calcium, magnesium, carbon and oxygen. As the
carbonates precipitate they inherit and incorporate chemical signatures that are ultimately
governed by the nature of formation fluids. Therefore, carbonate rocks and minerals can
be treated as geologic reservoirs for information concerning past fluid chemistries and
very powerful geochemical databases.
All sedimentary rocks, including carbonates, form directly on or near Earth’s
surface typically in close spatial association with biology. It is well documented that
carbonate minerals can form via the influence of organisms (biotic) or independent of
biological processes (abiotic). In addition while some carbonates may precipitate
abiotically, biological processes that influence fluid conditions have the possibility of
being recorded in carbonate minerals. Therefore the study of carbonates is important not
only in geological and chemical contexts but also in a biological context as a direct result
and/or a passive recorder of such processes.
The following dissertation describes five projects that exploit this carbonate
mineral geochemical reservoir across a range of geologic settings. The carbonates
presented in the following chapters can be categorized based on their relative time of
formation compared to the deposition of the associated geologic unit. These categories
are 1) primary, or syndepositional carbonates and 2) diagenetic, or carbonates
precipitated post depositionally and prior to metamorphic temperature and pressure
x
ranges. Both sedimentary regimes are suitable for biology and diagenetic environments
in particular can be strongly controlled by microbial processes.
In the following studies isotopic, trace elemental, trace compound, petrographic,
stratigraphic and textural data are combined in order to determine the formation
environments of carbonates and the characteristics of carbonate-yielding fluids. Given
the major constituents of carbonate minerals listed above, isotopic analyses of carbon and
oxygen are critical measurements and can provide a great deal of evidence regarding the
sources and cycling of theses two elements. Oxygen isotopes in carbonates are
influenced by a multitude of processes (discussed in detail in the following chapters)
most of which are directly reflective of fluid temperatures and their isotopic
compositions, both of which can help distinguish between primary or diagenetic
environments. Trace element concentrations of iron, manganese and strontium are
exploited in chapter 2 in order to further constrain the influence of diagenesis on primary
chemical signatures.
Carbon is a major constituent of not only carbonate minerals but also organic
compounds, making is isotopic characterization in carbonate minerals a valuable tracer of
the sources of carbon in sedimentary environments (Claypool and Kaplan, 1974). As
with carbon, sulfur is strongly influenced by biologic processes. Carbonates incorporate
trace amounts of sulfate upon precipitation (Burdett et al., 1986) and this sulfate has been
shown to substitute directly for the carbonate ion within the crystal lattice (Pingitore et
al., 1995). This sulfate is referred to as carbonate-associated sulfate (CAS) and can be
extracted and analyzed for not only its bulk concentration but also its isotopic
xi
composition, both of which are dictated by the nature of carbonate-yielding fluids
(Burdett et al., 1986). Carbonate formation environment can be further characterized
when analysis of CAS (the oxidized sulfur species) is combined with isotopic and
abundance analyses of pyrite (the dominant reduced sulfur phase in geologic settings).
Therefore carbonate minerals act as geologic reservoirs for sulfur systematics as well.
The following chapters describe studies which employ all or some of the above
geochemical measurements in addition to more traditional geologic approaches such as
petrography and stratigraphy.
Chapters 1 and 2 deal with Neoproterozoic to middle Cambrian units from
northwestern Mexico and eastern California. These deposits are interpreted here as
largely primary, however the influence of diagenesis is identified and discussed in each.
This time interval is characterized by perhaps the most extreme evolutionary radiation
experienced in Earth history, the so-called “Cambrian Explosion”. It has long been
proposed that certain chemical conditions must have existed in order to support such a
drastic radiation, in particular increased marine oxygen concentrations. The data
presented here suggests that this time interval was characterized by oceans with low
sulfate concentrations—sulfate is a redox sensitive compound and is expected to increase
in tandem with oxygen. These results require a reevaluation of the Earth’s oceans during
such an evolutionary significant time interval.
Chapters 3 and 4 focus on carbonate concretions from the Miocene Monterey
Formation and the late Cretaceous Holz Shale. Concretions have long been known to
form within sediments and their distinction as diagenetic is not largely debated. In these
xii
chapters, I show that concretionary carbonates retain signatures consistent with particular
microbial processes, and directly identify that sulfur cycling plays a large role in past
marine sedimentary regimes. In most cases, the identified microbial processes are likely
directly responsible for carbonate mineralization. Concretions of the Monterey
Formation exhibit chemical characteristics consistent with formation in sediments
experiencing organic matter degradation via oxidation by nitrate and/or metal oxides and
sulfate. Methanogenesis and sulfide oxidation were also active in zones of concretion
precipitation in sediments of the Monterey Formation. Calcitic concretions of the Holz
Shale possibly formed in zones experiencing extensive sulfide oxidation. Sulfide
oxidation has, until now, not been identified as a reaction associated with carbonate
authigenesis.
Chapter 5 highlights the possibility of forming potentially primary-like
sedimentary textures through diagenetic processes in the Beck Spring Dolomite of
eastern California. In this deposit, an extensive, laminated texture is most reasonably
interpreted as diagenetic when geochemical, textural and petrographic data are
considered together. This study demonstrates that the distinction between primary and
diagenetic must not be made solely on field-based criteria.

1
CHAPTER 1: CARBON AND SULFUR ISOTOPE CHEMOSTRATIGRAPHY OF
NEOPROTEROZOIC TO CAMBRIAN CARBONATES FROM
NORTHWESTERN MEXICO AND EASTERN CALIFORNIA: IMPLICATIONS
FOR SULFATE CONCENTRATIONS ACROSS THE PRECAMBRIAN-
CAMBRIAN BOUNDARY

CHAPTER 1 ABSTRACT
The sulfur isotopic compositions of carbonate-associated sulfate (CAS) and pyrite
from correlative Ediacaran and Cambrian carbonates of Caborca, Mexico and Death
Valley, California exhibit high stratigraphic variability. Δ
34
S (the difference between
δ
34
S
CAS
and δ
34
S
pyrite
) values average ~10‰ and never exceed 28‰, consistent with
limited fractionation expression and suggestive of low sulfate concentrations. In
addition, stratigraphic variability in δ
34
S
CAS
over short stratigraphic intervals is
comparable to that of Mesoproterozoic successions (e.g., Kah et al. 2004), suggesting that
marine sulfate concentrations were ca. 1.5-4.5mM, 5-15% of modern levels, well into the
Cambrian. Although the δ
34
S
CAS
variability

suggests low sulfate levels, CAS
concentration data (if primary) suggests that a moderate increase in marine sulfate took
place across the Precambrian-Cambrian boundary.
These data indicate that marine sulfate levels were low not only in the
Neoproterozoic, but well into the Cambrian. Typically, marine sulfate concentrations are
taken as an oxygenation proxy—sulfate is primarily delivered to the ocean via oxic
weathering of continental sulfides. In addition, increased atmospheric and oceanic
2
oxygen concentrations would suppress sulfate reduction to the sediments, allowing
sulfate concentrations to build in the ocean. A moderate increase in CAS concentration
may suggest marginal oxygenation of the marine realm across the Precambrian-Cambrian
boundary, but given the presumed metabolic requirements of metazoa, it is surprising that
the oxygenation signal is not more severe. The results of this study indicate that either 1)
complex life evolved and thrived in low sulfate and relatively low oxygen environments
compared to the modern or 2) marine sulfate concentrations did not closely track oxygen
concentrations.

INTRODUCTION
The transition from the late Precambrian into the early Phanerozoic represents a
critical time interval in Earth’s history. During this period the Earth experienced massive
glacial episodes and crossed major evolutionary thresholds, including the advent of
biomineralization and the radiation of complex, multicellular life (Knoll and Carroll,
1999). The geologically instantaneous appearance of large body forms in the fossil
record suggests that Earth’s oceans underwent chemical changes, possibly associated
with increased oxygen concentrations. As there is no direct geologic reservoir for O
2
the
oxygenation hypothesis must be explored through indirect proxies. Recent studies (i.e.,
Fike et al., 2006; Kaufman et al., 2007; Halverson and Hurtgen, 2007; McFadden et al.,
2008; Fike and Grotzinger, 2008) have utilized sulfur isotope systematics as a proxy for
oxygenation. While sedimentary pyrites are common, the record of the oxidized forms of
sulfur (e.g., evaporites) is sparse in the geologic record. However, trace amounts of
3
sulfate are incorporated into carbonates at the time of precipitation and this sulfate is
referred to as carbonate-associated sulfate (CAS). CAS, although trace in quantity, can
be extracted, quantified and isolated for sulfur isotopic analysis. CAS provides a unique
opportunity to explore the sulfate characteristics of the Neoproterozoic ocean because
carbonates (limestones and dolostones) are both stratigraphically and spatially more
continuous than evaporite counterparts. The preservation of sulfate in carbonates allows
high-resolution sampling of a redox-sensitive compound which, when used in
conjunction with δ
13
C
carb
and δ
34
S
pyrite
, provides insight into Earth’s early marine
environments.
Here we present stratigraphic data consisting of δ
13
C
carb
, δ
34
S
CAS
, δ
34
S
pyrite
and
CAS concentration from Neoproterozoic to middle Cambrian carbonates of eastern
California and northwestern Mexico (Figures 1.1 and 1.2). Analyses from two somewhat
distant (~800 km apart), time-equivalent localities allow for a more complete
characterization of ocean chemistry across the Precambrian-Cambrian Boundary (PCB).
In addition, this study highlights the possibility of regional deviation in both CAS
concentration and δ
34
S
CAS
and raises questions about the likelihood of a “global” sulfur
signal.

4
Figure 1.1: Locality map of the Death Valley and Caborca Neoproterozoic to middle
Cambrian successions

Shaded regions highlight mountain ranges of sampled outcrops. Stars indicate precise
sample sites: CP = Chicago Pass, EP = Emigrant Pass, SNR = Southern Nopah Range,
WP = Winters Pass Hills, CR = Cerro Rajon. Samples of the SNR and WP are from
Hurtgen et al. (2004) and Kaufman et al. (2007), respectively.

5
Figure 1.2: Stratigraphic correlation between units of Caborca and Death Valley

Correlation is based on radiometric, bisotratigraphic, lithostratigraphic and chemostratigraphic criteria as discussed in text.
6
GEOLOGIC CONTEXT
Caborca Region, Northwestern Mexico
Neoproterozoic to Cambrian deposits crop out near the town of Caborca in
Sonora, Mexico, and represent predominantly shallow marine carbonate and siliciclastic
depositional environments. Cerro Rajón, the type section for Neoproterozoic units in
northwestern Mexico, is ~30 km to the southeast of Caborca (Figure 1.1) and
stratigraphic units of this mountain range are exposed as a moderately northward dipping
homocline (Stewart et al., 1984). As with Neoproterozoic and Cambrian units from
eastern California (Corsetti and Kaufman, 2003), those of Caborca are both siliciclastic
and carbonate in composition (Figure 1.2). Mixed siliciclastic/carbonate successions are
extremely useful for chronologic correlation as they provide a medium for both
chemostratigraphic (primarily δ
13
C of carbonates) and bisotratigraphic (some important
PCB fossils only occur in siliclastic strata, i.e., Treptichnus pedum) reference frames. In
contrast to the Death Valley Succession (discussed below), the Caborca Succession is
relatively unstudied aside from some seminal work by Stewart et al. (1984). The
stratigraphic section consists of the following units in ascending order: the El Arpa
Formation, the Caborca Formation, the Clemente Formation, the Pitiquito Quartzite, the
Gamuza Formation, the Papalote Formation, the Tecolote Quartzite, the La Cienega
Formation, the Puerto Blanco Formation, the Proveedora Quartzite, the Buelna
Formation, the Cerro Prieto Formation and the Arrojos Formation. Table 1.1 outlines the
stratigraphy, biostratigraphy and notable features of each stratigraphic unit.

7
Table 1.1: Stratigraphic characterization of the Caborca and Death Valley Successions

8
Death Valley Region, Southwestern United States
Mixed siliciclastic/carbonate units from the Death Valley region of eastern
California (Figure 1.2) are broadly equivalent to those of the Caborca region, as
discussed in Stewart et al. (1984). The Death Valley succession has been studied in
detail (see, for example, Corsetti and Kaufman, 2003), and consists of the following
units, in ascending order: the Crystal Spring Formation (intruded by a 1.08 Ga diabase
sill; Heaman and Grotzinger, 1992), the Beck Spring Dolomite, the (partially glaciogenic;
Corsetti and Kaufman, 2003) Kingston Peak Formation, the Noonday Dolomite (a “cap
carbonate” atop the glacial deposits of the Kingston Peak Formation), the Johnnie
Formation, the Stirling Quartzite, the Wood Canyon Formation, the Zabriskie Quartzite,
and the Cararra Formation. Table 1.1 outlines the stratigraphy, biostratigraphy and
notable features of each stratigraphic unit.

Caborca-Death Valley Correlation
Currently, it is not clear how the pre-Clemente Formation units correlate with the
pre-Johnnie Formation units in Death Valley. The Caborca succession lacks glacial or
cap carbonate deposits characteristic of Neoproterozoic sections worldwide, whereas they
are present in the Death Valley succession. The Caborca succession contains some
highly positive δ
13
C values in the El Arpa Formation that have no counterpart in the pre-
Johnnie Formation units in Death Valley. However, the overlying units correlate quite
well via litho-, chemo-, and biostratigraphy, and will be the focus of the geochemical
investigation presented here.
9
Two large-magnitude, negative carbon isotopic excursions likely representing the
older Wonoka-Shuram (W-S; magnitude ~10‰ VPDB) and the younger PCB (magnitude
~7‰) events are exhibited in both the Caborca and Death Valley successions. Lithologic
similarities add to the robustness of the correlation of W-S excursion facies between
Caborca and Death Valley, including a distinct oolite-pink carbonate marker bed and the
presence of formerly aragonitic seafloor fans within the W-S excursion facies (Table 1.1
and Figures 1.2 and 1.3). The lithologic similarity has long been known and was used as
a key tie point by Stewart et al. (1984). Here, we strengthen the correlation with δ
13
C
chemostratigraphy and more detailed lithostratigraphy (primarily the identification of
carbonate fans). Key biostratigraphic features present in both successions include
fossilized Cloudina (thought to represent an age of ~548 Ma, Corsetti and Hagadorn,
2000) and Treptichnus pedum (PCB; ~542 Ma). While T. pedum is the recognized
marker of the PCB, the actual boundary in Caborca likely occurs within the La Ciénega
Formation in conjunction with the negative carbon excursion. The apparent “late”
appearance of the trace fossil probably reflects the life habits of T. pedums creator (i.e.,
the organism lived on siliciclastic and not carbonate substrates). Additional
bisotratigraphic correlations include the first appearance of trilobites/trilobite debris,
archeocyaths, Skolithos and morphologically similar oncoids (Table 1.1 and Figure 1.2;
Stewart et al, 1984).
Although abundant chronostratigraphic tie-points exist, the differences in
lithology between Caborca and Death Valley make dual analysis complementary.
Specifically, the relatively carbonate-rich Caborca succession allows more complete

10
Figure 1.3: Photographs of key lithostratigraphic markers from Caborca and Death Valley

Photographs of key lithostratigraphic markers from Caborca (A, C and E) and Death
Valley (B, D and F). A, B) Outcrop photos of Wonoka-Shuram facies. C, D)
Photomicrographs of basal oolite. E, F) Photomicrographs of formerly aragonitic crystal
fans from upper W-S facies pink carbonates.

11
development of carbon isotope chemostratigraphy. This is extremely useful for
completing δ
13
C trends in the underrepresented post W-S, PCB and lower/middle
Cambrian intervals in Death Valley. The Gamuza/Papalote, La Cienega, upper Puerto
Blanco, Buelna and Cerro Prieto Formations of Caborca shed light on ocean chemistry
through these blank intervals. Thus, the multiple chronostratigraphic markers between
and complementary lithologies of Caborca and Death Valley make these two successions
ideal for a regional study of sulfur systematics.

METHODS
Carbonate samples were taken in stratigraphic context from the entire
Neoproterozoic and Cambrian portions of the Cerro Rajon succession and from the
previously unsampled PCB and Cambrian facies of the Death Valley succession.
Previous studies by Hurtgen et al. (2004) and Kaufman et al. (2007) provide data from
the Noonday Dolomite (Southern Nopah Range) and Rainstorm Member carbonates
(Winters Pass Hills) of Death Valley, respectively (Figure 1.1). Our analyses include
isotopic (δ
13
C
carbonate
, δ
18
O
carbonate
, δ
34
S
CAS
and δ
34
S
pyrite
) and sulfur phase concentration
(CAS and pyrite) analyses (Table 1.2).

Carbonate Associated Sulfate Concentration
The extraction method used here is a modified version of that from Burdett et al.
(1989). Rock samples of at least 400g were powdered using a RockLabs rock crusher
that was meticulously cleaned between samples. Splits of 150g were removed and
12
Table 1.2: Compilation of geochemical data from this study (Caborca and upper Death
Valley Successions) and others (lower Death Valley Succession)

13
Table 1.2 continued

14
Table 1.2 continued

15
Table 1.2 continued

16
washed four times in ultrapure (18mΩ), de-ionized water and once in a sodium
hypochlorite (NaClO, bleach) solution. The washes remove non-structural sulfur phases
including soluble sulfates (anhydrite, gypsum) and organic sulfur. Washed powders were
dried, weighed and acidified in 3M HCl in teflon beakers at room temperature in order to
liberate lattice-bound sulfate (CAS) into solution as SO
4
2-
. We typically used 1L of acid
for 150 grams of sample which will allow for stoichiometrically complete dissolution, yet
with some insoluble residue, an excess of acid is assured. The samples were filtered
(0.45µm) to remove insoluble residues which were then quantified gravimetrically. The
mass of insoluble residue was subtracted from the initial powder mass in order to
quantify weight percent carbonate (assuming that all of the dissolved material was pure
carbonate). The supernatant fluids were heated to ~70°C and a 30% BaCl
2
solution
added to induce BaSO
4
(barite) precipitation. Precipitation progressed at room
temperature for 72 hours in order to ensure reaction completion. The barite-containing
solutions were filtered and barite concentration determined gravimetrically. Sulfate
concentration was determined based on the mass of barite and CAS reported in parts per
million (ppm) compared to the amount of pure carbonate in each sample. Duplicate
measurements of CAS concentration were within +/–35 ppm of reported values. Two
samples exhibit replicates that do not agree within +/–35 ppm, as shown by error bars in
Figure 1.4.

17
Pyrite Concentration
Pyrite concentration was determined via the chromium reduction method as
described by Canfield et al. (1986). Two gram splits of the insoluble residues acquired
from CAS extraction were isolated from filter paper and reacted in a 1M CrCl
2
/HCl
solution under a N
2
atmosphere. The chromium solution converted any sulfide phases
(strictly pyrite in geologic samples as other reduced phases convert to FeS
2
over geologic
timescales) to gaseous H
2
S which was passed into a trap containing a 3% AgNO
3
/10%
NH
4
OH solution. Reaction of H
2
S with AgNO
3
produced solid-phase Ag
2
S which was
filtered from solution and quantified gravimetrically. Pyrite concentration was then
determined stochiometrically and reported in weight % (wt%) compared to the original
powder mass. Standard analyses with known FeS
2
indicate that this method and the
extraction apparatus are consistent (to ±15%) and accurate to within ±10% of reported
values.

Isotopic Analyses
Carbon and Oxygen: Carbonate carbon and oxygen isotopic analyses were conducted
on a VG Prism II isotope ratio mass spectrometer (IRMS) at the University of Southern
California. Samples were microdrilled from thin section billets after being examined
under a petrographic microscope to ensure sampling of the most pristine phases.
Reproducibility of these measurements is better than 0.1‰ for both carbon and oxygen.
Oxygen and carbon isotopic values are recorded in the standard δ notation in comparison
to the VPDB standard.
18
Sulfur: Sulfur isotopic analyses were conducted on a ThermoScientific Delta V Plus
IRMS at the University of California, Riverside. The IRMS is interfaced with a Costech
Analytical Technologies, Inc. elemental combustion system via a Thermoscientific
CONFLO III interface. Sulfur isotope values are reported in the standard δ notation in
comparison to the VCDT standard. Interlaboratory (compared to values generated at the
University of Maryland) and replicate analyses yield values consistent within +/–1.0‰.

RESULTS
CAS Concentration
In general, carbonates from Death Valley contain higher concentrations of CAS
(avg. 159ppm) than carbonates from Caborca (54ppm), however stratigraphic similarities
occur between the two localities (Figures 1.4 and 1.5). Three distinct groupings can be
recognized based on relative concentrations; 1) the W-S facies (moderately high CAS), 2)
the pre-PCB interval (low CAS), excluding W-S facies and 3) the PCB-Cambrian interval
(moderately high CAS).
CAS concentrations in carbonates of Caborca exhibit values between 1) 0 and
512ppm (avg. 209ppm) in the W-S facies, 2) 0 and 131 ppm (avg. 26ppm) in the pre-
PCB interval and 3) 0 and 466 ppm (avg. 54ppm) in the PCB-Cambrian interval. In
Death Valley, carbonate CAS concentrations range from 1) 6 to 1,491ppm (avg. 363ppm)
in the W-S facies, 2) 0 to 272ppm (avg. 51ppm) in the pre-PCB interval and 3) 0 to
1,086ppm (avg. 314ppm) in the PCB-Cambrian interval. In both sections, carbonates of

19
Figure 1.4: Chemostratigraphic trends of Cerro Rajon, Caborca, Mexico

20
the PCB-Cambrian interval contain higher concentrations of CAS than those of the pre-
PCB interval and the W-S facies contain the highest concentrations exhibited out of all of
the sampled intervals.

CAS δ
34
S
In Caborca and Death Valley carbonates, δ
34
S
CAS
values exhibit high stratigraphic
variability (Figures 1.4 and 1.5). Isotopic swings of ~10‰ in less than 100m of section
are relatively common. The ranges in isotopic values are –1.5 to 31.6‰ in Caborca and
11 to 37.3‰ in Death Valley. Interestingly, when small-scale variability is ignored, both
sections display sulfur isotopic values oscillating around that of modern seawater sulfate
(~21‰). Average δ
34
S
CAS
values are 21.6 and 22.5‰ in Caborca and Death Valley,
respectively.

Pyrite δ
34
S
δ
34
S
pyrite
values also display high stratigraphic variability. The δ
34
S
pyrite
values
from Caborca range from –12.2 to 30.5‰ and average 13.3‰. Death Valley samples
exhibit δ
34
S
pyrite
values that range from 2.9 to 30‰ and average 15.6‰. It is interesting
to note that most of these values are mildly to highly positive, a feature that will be
discussed, below.

21
Figure 1.5: Chemostratigraphic trends of the Death Valley Succession, eastern California, USA.

22
Δ
34
S
The relative enrichment in
34
S between contemporaneous δ
34
S
CAS
and δ
34
S
pyrite
is
denoted by Δ
34
S, such that Δ
34
S = δ
34
S
CAS
– δ
34
S
pyrite
. Δ
34
S values of Caborca carbonates
range from –22.9 to 27.1‰ and average 8.7‰. Death Valley carbonates exhibit Δ
34
S
values that range from 2.0 to 22.4‰ and average 12.9‰. In carbonates of Caborca, Δ
34
S
exhibits high variability with generally increasing values through the Neoproterozoic, a
maximum near the PCB and decreasing values progressively up section. Although
incomplete stratigraphically, Δ
34
S values of Death Valley fit well with the Caborca data
(Figure 1.6).

INTERPRETATION
It is important to note that, while similar, the trends in CAS concentration and
δ
34
S
CAS
exhibited by Caborca and Death Valley are not identical (Figure 1.6). Assuming
these proxies reflect primary signals, the differences in two relatively proximal
environments suggest that heterogeneity in Neoproterozoic and Cambrian marine sulfate
could have existed. Without additional information across the PCB the possibility of
regional or global heterogeneity cannot be discounted. Despite the fact that they are not
identical, similar features exist and will be highlighted here. First, trends in CAS
concentration will be discussed in detail, followed by an interpretation of the sulfur
isotopic data.

23
Figure 1.6: Transposed plots of the Caborca and Death Valley δ
34
S
CAS
, δ
34
S
pyrite
and Δ
34
S
data

Notice how the δ
34
S
CAS
and δ
34
S
pyrite
trends are somewhat similar but not identical. Data
points plotted based on correlation via δ
13
C. Neoproterozoic and Cambrian δ
13
C from
Fike et al. (2006) and Zhu et al. (2006), respectively. CAS and δ
34
S
CAS
data of Death
Valley from Hurtgen et al. (2004), Kaufman et al. (2007) and this study; remaining
geochemical data from this study. Radiometric dates obtained by comparing δ
13
C to
alternate localities including Oman [O] (Fike et al., 2006), South China [SC] (McFadden
et al, 2008) and Morocco [M] (Maloof et al., 2005). Timescale-δ
13
C correlation as
presented in Ogg et al., 2008.

24
Trends in CAS Concentration
Modern carbonates precipitated in equilibrium with seawater record CAS
concentrations between ~1000 and 4500 ppm and average ~2400 ppm (Lyons et al.,
2004; Gellatly and Lyons, 2005). However, the incorporation of CAS into the carbonate
lattice is not well understood, and may be dependent on mineralogy as well as
crystallization rate (Busenberg and Plummer, 1985). For example, biogenic carbonates
typically incorporate significantly more CAS versus abiotic precipitates, and aragonite
incorporates significantly more than calcite. Meteoric diagenesis is known to strip sulfate
from the carbonate lattice, while leaving the δ
34
S relatively unchanged (Gill et al., 2008).
Thus, great care must be exercised when interpreting ancient CAS concentrations.
As noted above, the CAS concentration record can be subdivided into a
Precambrian interval with CAS values significantly below modern values, a W-S interval
with anomalously elevated CAS concentrations, and a Cambrian interval with CAS
concentrations approaching modern values. It is tempting to interpret the CAS
concentration to reflect changes in seawater sulfate concentration, especially since the
lower increase in CAS concentration coincides with the W-S anomaly, and the upper
increase with the PCB. However, it is more prudent to dissect the CAS trends, focusing
on their stratigraphic and diagenetic context first. Diagenetic considerations are obvious,
but mineralogic and biogenic factors will be discussed, as well.

Diagenetic considerations: Meteoric diagenesis removes CAS from carbonates (Gill et
al., 2008), but it is not clear that the Precambrian section was preferentially altered

25
compared to the remaining succession. Indeed, similar facies are present in all portions,
and thus one would expect similar exposure to meteoric systems. Dolomitization, on the
other hand, may play a larger role. It is clear from Figure 1.7 that, on average, limestones
record more CAS in both the Caborca and Death Valley successions versus dolostones.
In order to isolate the artifacts of dolomitization, it is useful to examine CAS
concentrations in carbonates of the same mineralogy (i.e., compare limestones to
limestones and dolostones to dolostones). Comparison of this nature allows recognition
of primary trends across the PCB. As Figure 1.8A demonstrates, a large proportion of
Cambrian dolostones (~20%) contain elevated concentrations of CAS, in contrast to
Precambrian dolostones. In addition, a significant proportion of Cambrian limestones
(~25%) exhibit high CAS concentrations (Figure 1.8B), although no Precambrian
limestones are available (aside from the W-S facies limestones) for comparison.

Mineralogic Considerations: Interestingly, the units with the highest CAS
concentrations associated with the W-S interval are limestones and dolomites exhibiting
formerly aragonitic phases (crystal fans and fibrous ooids; see Figure 1.3). As discussed
above, the fans are currently calcite and the ooids are currently dolomitic but they are
pseudomorphs after aragonite. It is possible that they record more CAS because
aragonite typically incorporates more CAS than calcite, a feature that may be retained
during neomorphism, similar to the way formerly aragonite phases record higher Sr
concentrations (e.g., Sandberg, 1983).

26
Figure 1.7: CAS concentrations compared to lithology

Notice how some limestones and dolostones exhibit near zero CAS concentrations, likely
indicative of at least minor diagenetic removal in both lithologies. Limestones dominate
the high CAS concentration grouping (250+ ppm).

27
Increased Biogenic Component in Cambrian Carbonates: Some modern biogenic
carbonates have been found to incorporate more CAS than abiotic counterparts (Burdett
et al., 1989; Lyons et al., 2004) and early Cambrian biogenic carbonates may have been
similarly enriched in CAS. Thus, up section changes in CAS concentration may be
caused by increases in the relative contribution of biogenic material. In Caborca,
limestones of the middle to upper Puerto Blanco Formation are composed of densely
packed archeocyath debris. However, samples from this interval exhibit CAS
concentrations from 0 to 167ppm, indistinguishable from Neoproterozoic concentrations.
In fact, dominantly micritic and oolitic limestones (abiotic) stratigraphically above the
archeocyath accumulations exhibit relatively elevated CAS concentrations of up to
260ppm. Additionally, limestones from the Cambrian units of Death Valley are
dominantly composed of micrite that is essentially devoid of biogenic material, yet these
contain the highest concentrations of CAS recognized in this study (aside from those of
the W-S facies). Therefore, while biogenic material can complicate the interpretation of
stratigraphic trends in CAS concentration, it cannot reasonably account for the trends of
Death Valley and Caborca.

Is there a CAS concentration trend?: If we combine the caveats presented above, and
compare non-formerly aragonite normal marine facies to like facies, it is clear that CAS
concentrations increase moderately from the PCB to the Middle Cambrian (see Figure
1.8). This is consistent with increased oxygenation of the marine realm, but frankly, the
magnitude of the change is unremarkable considering the importance given to

28
Figure 1.8: CAS concentrations for each of the three intervals

Intervals are highlighted in Figures 4 and 5 and are separated into A) dolostones and B)
limestones.

29
oxygenation and the Cambrian Explosion. The significance of this will be combined with
the isotopic data and discussed, below.

Sulfur Isotopes and Marine Sulfate
Where CAS concentration may be subject to many factors, the δ
34
S of CAS and
coeval pyrite offers another window into the Ediacaran-Cambrian sulfur cycle and
oxygenation. In the following sections the variability in δ
34
S
CAS
and Δ
34
S will be
discussed and related to marine sulfate concentrations.

High Stratigraphic Variability in δ
34
S
CAS
: The high variability in δ
34
S
CAS
throughout
both the Caborca and Death Valley successions is likely indicative of a low sulfate ocean.
In the modern, oceanic δ
34
S
sulfate
is relatively invariant (Paytan et al., 1998; 2004a;
2004b), largely due to the high concentration and long residence time of marine sulfate.
Isotopic mass balance dictates that large changes in δ
34
S of seawater sulfate can only
occur if the sulfate concentration is low, and thus easy to perturb. Kah et al. (2004)
demonstrated that major swings in δ
34
S over relatively short stratigraphic intervals (on
the order observed here) would require low sulfate concentrations, generally between 1
and 4 mM (versus 28 mM in the modern ocean). The high stratigraphic variability in the
data presented here suggests that oceanic sulfate was low not only in the Neoproterozoic,
but also well into the Cambrian.
It is important to mention that δ
34
S
CAS
values may be contaminated by pyrite
oxidation during samples analysis (Marenco et al., 2008). This contamination should

30
produce a negative correlation between δ
34
S
CAS
and pyrite concentration. Figure 1.9
demonstrates that no such correlation exists, suggesting that the samples presented here
are not contaminated by pyrite oxidation.

Δ
34
S
The stratigraphic trend in Δ
34
S from Caborca carbonates is shown in Figure 1.4.
The general increase toward the PCB is similar to that observed by Fike et al. (2006),
however the Caborca data set displays much higher variability and Δ
34
S values that never
exceed 28‰. An increase in Δ
34
S could reflect an increased isotopic fractionation
between sulfate and sulfide, possibly associated with increased sulfate concentrations
(Habicht et al., 2002). Alternatively, the increase could reflect the preferential influence
of an alternate pathway of sulfate reduction that imparted a more extreme isotopic
fractionation. Bacterial sulfur disproportionation (BSD) is known to impose a more
severe fractionation with expressed Δ
34
S values up to 70‰ (Canfield and Thamdrup,
1994; Canfield and Teske, 1997). BSD is dependent on the presence of intermediate
sulfur species (i.e., elemental sulfur, thiosulfate, polysulfides, etc.) and it is typically
assumed that an increase in oxygenation state would promote the production of these
intermediates. However, an increase in oxygenation state does not have to be
accompanied by a significant increase in sulfate concentration, particularly in an
environment experiencing vigorous sulfate reduction. In addition, recent work by
Canfield et al. (2010) demonstrates that isotopic fractionations up to ~70‰ can be
achieved at low sulfate concentrations (1.1-2mM) via bacterial sulfate reduction alone

31
Figure 1.9: Cross plot of δ
34
S
CAS
and pyrite concentration.

Contamination by pyrite oxidation should produce a negative correlation, which is not
exhibited by these samples.

32
(no BSD required). Therefore, the increase in Δ
34
S near the PCB need not be the result of
or accompanied by a large increase in dissolved sulfate concentrations.
The overall low Δ
34
S values in both Caborca and Death Valley carbonates (below
28‰) suggest that while sulfate concentrations may have fluctuated, they remained
relatively low. Modeling by Habicht et al. (2002) suggests that Δ
34
S values of this
magnitude are consistent with sub-millimolar sulfate concentrations. However,
variability in Δ
34
S cannot be attributed solely to primary marine processes as sulfate
reduction dominantly occurs in sediments in marine environments. Low oxygen
concentrations may have allowed water column sulfate reduction, however little is known
about the degree of oceanic anoxia in the PCB system. Therefore, while the δ
34
S
CAS

values of pristine carbonates likely record water column conditions, δ
34
S
pyrite
(and thus
Δ
34
S) values could reflect processes in the water column or the sediments.

The Marine Sulfate Reservoir and Implications for Oxygenation
All of the data presented here indicate an overall relatively low (at least
regionally) marine sulfate reservoir. Increased CAS concentrations in Cambrian
carbonates suggest a marginal increase in sulfate, however the retention of highly
variable δ
34
S
CAS
and low Δ
34
S are difficult to explain without low concentrations of
dissolved sulfate.
The isotopic variability in δ
34
S
CAS
exhibited by Caborca and Death Valley
carbonates is comparable to trends recognized in the Mesoproterozoic Society Cliffs
Formation and Dismal Lakes Group (Kah et al. 2004). Model results suggest that

33
isotopic variability of ~ ±10‰ over ~200-300m is consistent with Mesoproterozoic
marine sulfate concentrations of ~1.5-4.5mM. Although chronostratigraphic constraints
are not as rigorous in the Caborca and Death Valley successions, the comparable
variability suggests that the Neoproterozoic to middle Cambrian oceans exhibited similar
sulfate concentrations.
A lack of significant sulfate increase across the PCB is at odds with the proposed
oxygenation required to support the Cambrian Explosion. A marine oxygenation event
should produce a contemporaneous increase in marine dissolved sulfate (reflected by
reduced stratigraphic variability in δ
34
S
CAS
). In addition, oxygenation should produce
decreases in the δ
34
S of both CAS and pyrite, as a consequence of increased continental
weathering. The geochemical characteristics of Caborca carbonates have implications for
the state of the Earth’s ocean during the evolution and proliferation of complex life
forms. These data suggest that: 1) early Cambrian organisms of Death Valley and
Caborca did not evolve in an environment with high sulfate and oxygen or 2) sulfate did
not track oxygen closely during this time interval.

Elevated δ
34
S
pyrite
in Neoproterozoic to Middle Cambrian Carbonates
The sulfur isotopic compositions of sedimentary pyrites have varied significantly
over geologic time (Canfield, 1998). In general, the sulfur isotopic composition of
pyrites is depleted compared to coeval sulfates in units younger than ~2.7 Ga (Thode and
Goodwin, 1983), reflecting the advent of microbial sulfate reduction. However, some
Neoproterozoic units exhibit elevated δ
34
S
pyrite
values that, in some circumstances, exceed

34
the sulfur isotopic composition of contemporaneous sulfate minerals (that is Δ
34
S < 0).
These isotopic characteristics can only be achieved if the sulfate and sulfide (or their
respective reservoirs, i.e., carbonates for sulfate and pyrite for sulfide) are forming in
environments distinct from one another. In modern marine environments, pyrite
dominantly forms in sediments where closed-system behavior promotes a tandem
increase in pore water δ
34
S
sulfate
and δ
34
S
sulfide
through progressive bacterial sulfate
reduction. Carbonate, on the other hand, forms dominantly in the water column such that
any incorporated CAS records the isotopic composition of water column sulfate.
Perhaps the increase in δ
34
S
pyrite
towards the PCB recognized here and by Canfield
(1998) corresponds to a retreat of sulfate-reducing communities into the sediments. In
this situation, the end Neoproterozoic represents a transitional period between marine
environments dominated by water column sulfate reduction and dominated by
sedimentary sulfate reduction.
What would drive sulfate reduction into the sediments? Many modern sulfate-
reducing communities are obligate anaerobes that cannot survive in oxygenated
environments. Therefore, increased marine oxygen concentrations could have driven
sulfate reducers into the sediments, where oxygen concentrations were kept low through
reactions with organic compounds (Claypool and Kaplan, 1974). Thus the elevated
δ
34
S
pyrite
and negative Δ
34
S values recognized in this study and other successions globally
could reflect an increase in oxygen concentrations approaching the PCB. Such a situation
can still account for generally low sulfate oceans (as interpreted above) if sedimentary
sulfate reduction were sufficiently vigorous during this time interval.

35
The return to relatively low δ
34
S
pyrite
and positive Δ
34
S values in the Late
Cambrian and remainder of the Phanerozoic may represent the demise of strong
sedimentary, closed-system behavior due to increased bioturbation. Mixing by
burrowing organisms would enhance bioirrigation and preclude severe isotopic
enrichments via bacterial sulfate reduction.

CONCLUSIONS
Comparative isotopic and trace sulfate analyses of late Neoproterozoic to Middle
Cambrian carbonates from Death Valley and Caborca reveal both broad similarities and
moderate differences between the two localities. Units from both locations exhibit
increasing CAS concentrations coincident with the transition from the Neoproterozoic to
the Cambrian. This transition is also characterized by a lithologic transition from
dolostones to limestones. In all likelihood, the lithologic transition is at least partly
responsible for the increase in CAS concentration across the PCB. δ
34
S
CAS
values display
high stratigraphic variability in Death Valley and Caborca and lack precise agreement
between the two locations. The magnitude of variability in δ
34
S
CAS
is consistent with the
presence of relatively low sulfate conditions well into the Cambrian, perhaps comparable
to Mesoproterozoic sulfate levels ca. 1.5-4.5mM. High δ
34
S
pyrite
and negative Δ
34
S values
in the late Neoproterozoic suggest that sulfate reduction was restricted to the sediments
near the PCB, perhaps due to increased oxygen concentrations. Ultimately, this
restriction may have attributed to the increase in CAS across the PCB although the

36
retention of highly variable δ
34
S
CAS
indicates the persistence of relatively low sulfate
concentrations.
The results of this study suggest that heterogeneity in δ
34
S
CAS
could have existed
in the Neoproterozoic and Cambrian and that data from any one locality should not be
attributed to global phenomena. In addition, relatively low sulfate concentrations through
the PCB is seemingly at odds with a drastic oxygenation event contemporaneous with the
“Cambrian Explosion”, as has been proposed. In light of these findings we must consider
that either 1) the proliferation and diversification of complex life did not occur in oceans
exhibiting near modern oxygen and sulfate concentrations or 2) sulfate did not closely
track oxygen concentrations across the PCB.

37
CHAPTER 2: LOCAL VARIABILITY IN CARBONATE-ASSOCIATED
SULFATE AND δ
34
S IN ~580 MA CARBONATES OF THE CLEMENTE
FORMATION, CABORCA, MEXICO: IMPLICATIONS FOR DIAGENESIS AND
NEOPROTEROZOIC MARINE SULFATE

CHAPTER 2 ABSTRACT
The Neoproterozoic δ
13
C record from carbonates has been reproduced at many
stratigraphic sections around the world, and is generally considered a robust stratigraphic
tool as well as a window into the past carbon cycle. Recently, δ
34
S records have been
produced from carbonate-associated sulfate (CAS) in an attempt to understand the coeval
sulfur cycle, and thus better understand the oxidation state of the Neoproterozoic oceans.
Here, we select one stratigraphic interval—globally well defined by a uniquely negative
excursion in δ
13
C, commonly termed the Wonoka-Shuram (W-S) excursion—to explore
the sulfur isotopic record. First, we examine the δ
34
S record from multiple, closely
spaced sections near Caborca, Mexico, to explore local heterogeneities, and then compare
the Caborcan record with sections from around the world, to investigate the global sulfur
cycle.
The W-S excursion is located in the Clemente Formation and coincides with a
locally extensive carbonate marker bed and associated strata. Five sections over 20 km of
lateral distance were compared and record (in total) moderate variability in δ
34
S
CAS

(range from +18.6 to +27.6‰ VCDT) and significant variability in CAS concentration
(range from <30 to >1,200 ppm). While the δ
34
S values revealed no correlation with

38
diagenetic indicators, the CAS concentrations show strong negative correlation with
Mn/Sr ratios and Fe
carb
concentrations, consistent with CAS removal upon meteoric
recrystallization. We interpret the variability in isotopic composition to arise either from
local heterogeneity in marine δ
34
S
sulfate
or non-meteoric diagenetic overprinting.
When compared to coeval strata in Death Valley (~800 km to the north), both
sites record an overall average decrease in δ
34
S and an increase in CAS concentration, but
the absolute values, the detailed isotopic profiles, and the magnitude of the changes are
distinct from one another.
The heterogeneity noted on the local (Caborca) and regional (Caborca-Death
Valley) scales is mirrored in the global δ
34
S record. Similar trends in δ
34
S
CAS
are not
recognized from W-S excursion-containing units of South China or Oman, suggesting
that 1) the Neoproterozoic oceans were heterogeneous with respect to δ
34
S
sulfate
, 2) some
of these localities do not record “open-ocean” conditions due to basin restriction and/or
3) diagenesis has altered carbonates at some or all of these locations. The disagreement
in δ
34
S
CAS
has not as of yet been explicitly addressed, however it must be considered in
order to rigorously interpret the state of Neoproterozoic oceans ca. 580 Ma and
demonstrates that δ
34
S
CAS
cannot be used as a correlation

tool.

INTRODUCTION
Stable isotopic compositions (δ
13
C and δ
34
S) of Neoproterozoic marine
precipitates (i.e., limestones, dolostones, and evaporites) display extreme stratigraphic
trends (Halverson et al., 2005 and references therein; Hurtgen et al., 2004; Fike et al.,

39
2006; Kaufman et al., 2007; McFadden et al., 2008). The similarity in δ
13
C
carbonate
values
among temporally equivalent rock units has prompted the idea of a globally
homogeneous ocean with respect to δ
13
C
DIC
and has led to the adaptation of carbon
isotope chemostratigraphy as a correlation tool in the absence of biostratigraphy
(Halverson et al., 2005 and references therein). Unlike δ
13
C
carbonate
, the δ
34
S of evaporites
and trace sulfate in carbonate rocks has not been shown to be globally homogeneous, and
yet recent Neoproterozoic studies seek to attribute trends in both δ
13
C and δ
34
S to global
phenomena (c.f., Fike et al., 2006; Kaufman et al., 2007; Halverson and Hurtgen, 2007;
McFadden et al., 2008; Fike and Grotzinger, 2008). In order to better understand the
Neoproterozoic ocean system it is necessary to evaluate the lateral extensiveness of
isotopic and chemical trends, particularly those of carbonate-associated sulfate (CAS)
concentration and δ
34
S
CAS
, prior to spatial interpretation.
The most distinctive carbon isotopic event in Neoproterozoic time (and perhaps
all time), is the so-called Wonoka-Shuram (W-S) event, where δ
13
C values in carbonates
plummet to ~ –11‰, ca. 580-550 Ma (see Halverson et al., 2005 and references therein).
The W-S event provides a unique tie point, as it has been identified in many successions
around the world, including Oman, Namibia, Australia, India, South China, Eastern
China, and the southwestern United States (Burns et al., 1994; Calver, 2000; Amthor et
al., 2003; Corsetti and Kaufman, 2003; Le Guerroue et al., 2006; Kaufman et al., 2006;
Zhou and Xiao, 2007). Although some have recently attributed the W-S event to either
meteoric (Knauth and Kennedy, 2009) or burial diagenesis (Derry, 2010), the ubiquity of
the excursion across multiple depositional facies within single basins, and its conspicuous

40
stratigraphic position above Neoproterozoic glacial deposits and below the Precambrian-
Cambrian boundary would seem to preclude a purely diagenetic origin. Here, we will use
the presence of the W-S as a stratigraphic tie point from which to compare the δ
34
S
record locally, regionally, and globally, in order to better understand the nature of the
sulfur cycle in Neoproterozoic time.
The carbonates of the Clemente Formation, Caborca, Mexico, will constitute our
test case, as they record a very large magnitude negative carbon isotopic excursion (down
to ~ –10‰ VPDB), which we correlate to the W-S event. The strata in question extend
~20 km laterally among three mountain ranges: Cerro Rajon (CR), Cerro Clemente (CC)
and Cerro Calaveras (CCv) (Figure 2.1). Geochemical analyses among the three
localities provide insight into the extensiveness and variability of chemical signatures
during this climactic time interval. Here, we explore the lateral marker bed CAS
concentration and δ
34
S
CAS
variability as well as their relationship to traditional proxies of
carbonate diagenesis. The primary objectives of this study are to 1) examine the impact
meteoric diagenesis on CAS proxies (concentration and δ
34
S) at the local scale, 2)
comment about the “global” nature of CAS proxies at ~580 Ma and 3) provide an
environmental interpretation of the most unaltered data.

GEOLOGIC SETTING
Neoproterozoic units crop out in northwestern Sonora, Mexico, primarily to the
south of the town of Caborca (Figure 2.1). These strata consist of mixed siliciclastic-
carbonate successions that span the late Neoproterozoic to the Early Cambrian.

41
Figure 2.1: Map of study sites and other Neoproterozoic facies in the Caborca region

Sampled ranges (gray) and other ranges containing Neoproterozoic facies in the Caborca
region. Modified from Sour-Tovar et al. (2007).

42
Dominant sedimentary lithologies include dolomite, shale and quartzite with minor chert,
basalt and limestone. Miogeoclinal deposits of Cerro Rajon, the type section for
Neoproterozoic units in northwestern Mexico (Stewart et al., 1984), unconformably
overlie the Aibo Granite. The Aibo Granite is basement for many of the passive margin
deposits of the Caborca region and has been dated at ~1.11 Ga (Anderson et al., 1979;
Rodrigues-Castaneda, 1994). Further chronological constraint is provided
stratigraphically higher in the La Cienega Formation where Cloudina has been reported
(McMenamin, 1984; 1996; Sour-Tovar et al., 2007), representing an age of ~548 Ma
(Grotzinger et al., 1995; Corsetti and Hagadorn, 2000). In addition, Treptichnus pedum,
the fossil that marks the base of the Cambrian, has been located in the lowermost member
of the overlying Puerto Blanco Formation (see Figure 2.2 for stratigraphic context; Sour-
Tovar et al., 2007).
Above the Aibo Granite lies ~200 m of massive to thinly bedded dolomite and
sandy dolomite of the El Arpa and Caborca Formations. The El Arpa-Aibo Granite
contact is erosional and in portions exhibits meter-scale incision. In addition, the
lowermost El Arpa contains clasts of the Aibo Granite (also recognized by Damon et al.,
1962; Anderson et al., 1979). The Clemente Formation lies above the Caborca Formation
at Cerro Rajon and consists of an additional ~200 m of section dominated by siliciclastic
facies with minor dolomite, sandy dolomite and limestone (Stewart et al., 1984).
Approximately 133 m above the Clemente-Caborca contact is a 2.6 m thick carbonate
marker bed composed of a basal oolite and an upper, finely-laminated micrite (Figures

43
Figure 2.2: Stratigraphic column of Neoproterozoic units from Cerro Rajon

Zoom-in provides stratigraphic context of marker bed facies. The marker bed contains
carbon isotopic values consistent with the ~580 Ma Wonoka-Shuram global negative
excursion. Scale bar corresponds to the left hand column. Age constraints are discussed
in detail in text.

44
2.2 and 2.3). Within these carbonates a large magnitude negative carbon isotopic
excursion expresses δ
13
C values down to ~ –10‰ and is likely correlative to the ~580 Ma
Wonoka-Shuram excursion (see Figures 14 and 15 in Halverson et al., 2005). Above the
marker bed, the Clemente Formation returns to a primarily siliciclastic succession of
alternating shales and quartzites with minor thin sandy dolomite beds. The Clemente
Formation is conformably overlain by the ~75 m thick fine- to medium-grained and
commonly crossbedded Pitiquito Quartzite. The carbonate marker bed described above is
the focus of this study.
METHODS
Five stratigraphic sections of the carbonate marker bed were measured from three
localities in the Caborca region: Cerro Rajon (CR-1 and CR-2), Cerro Clemente (CC-1
and CC-2) and Cerro Calaveras (CCv-1). Samples taken at regular intervals were
analyzed for major and trace elemental concentration (Ca, Mg, Sr, Mn, Fe), trace sulfate
concentration (CAS: carbonate-associated sulfate), pyrite concentration and carbon,
oxygen and sulfur isotopic (δ
13
C
carbonate
, δ
18
O
carbonate
, δ
34
S
CAS
) analysis. Refer to Table 2.1
for geochemical data. Microscopic examination was conducted in order to confirm
marker bed correlations using rare textures and structures and to document textural
changes possibly associated with diagenesis.
Refer to chapter 1 for methods concerning CAS concentration, pyrite
concentration and isotopic analysis. Duplicate measurements of CAS concentration were
within +/–8% of reported values.

45
Table 2.1: Marker bed geochemistry.

46
Figure 2.3: Photographs of marker bed components

A) Marker bed at Cerro Rajon (CR-1). B) Close-up of lowermost oolite component. The
buff colored clasts are composed of oolite and float in an orange ooid matrix (pen for
scale). C) Flat pebble conglomerate facies. Clasts are highly elongate and are steeper in
lower, brownish-orange portion and dip more shallowly upward (pen for scale). D) Flat
pebble layer between laminated dolomite/limestone. Notice the high variability in clast
dips. E) Scanned slab of FPC. F) Uppermost pink and purplish laminated facies.

47
RESULTS
Marker Bed Petrography
The carbonate marker bed (Figure 2.3A) of the Clemente Formation exhibits a
laterally consistent stratigraphic pattern with basal oolite (mimetic dolomite) overlying
reddish-brown quartzite in an erosional contact (Figure 2.3B). Overlying the oolite in
places is a buff to pinkish-gray dolomite containing highly elongate clasts of dolomite
and/or pinkish-gray limestone (referred to here as flat-pebble conglomerate or FPC;
Figure 2.3C-E). Above the FPC lies a fine-scale, wavy-laminated dolomite/limestone
(Figure 2.3F). Finally, the laminated component is overlain by purple shale in
depositional contact. At sample sites CCv-1 and CC-2 the FPC is absent and the
laminated dolomite/limestone directly tops the oolite. At CC-1, the FPC occurs between
two discrete laminated horizons. Each major lithologic component is described in detail
below.

Oolite component: The lowermost buff-colored oolite ranges in thickness from ~60 to
270 cm. Horizons of matrix-supported conglomerate occur within the oolite.
Conglomerate clasts are typically 1-10 cm in diameter, subrounded to rounded and are
composed of dolomitic oolite or buff-colored, featureless dolomite (Figure 2.3B).
Ooids range in diameter from ~100 to 700 µm and are circular to oval in cross
section (Figures 2.4A, 2.5A). Ooid interiors display intricate concentric banding and
interstices are filled by anhedral sparry or bladed cements. Bladed cements exist as
isopachous rims where present and are followed by a sparry, second generation cement.

48
Figure 2.4: Photomicrographs of oolitic and laminated components of the marker bed

A) Oolite of CC-2. Interstices are filled by sparry cement. B) Crystal fan grouping
showing radiating habit (CC-1). C) Cluster of crystal fan blades of CR-1. D) Close-up of
fan blade from CC-1. Notice square termination, indicative of an aragonitic precursor.
E) Bedding-parallel thin section photomicrograph from laminated component of CR-1.
Notice polygonal morphology. F) Possibly recrystallized fan grouping of CC-2. The size
and upward doming shape is reminiscent of well-preserved fans from CR-1, CR-2 and
CC-1.

49
Some samples do not exhibit bladed cements and are cemented entirely by spar.
Boundaries and laminae of CCv-1 ooids are less clear than the other localities and appear
ghost-like (Figure 2.5B).

Flat Pebble Conglomerate Component: The middle member of the marker bed
sequence is composed of buff to reddish-gray, matrix supported, limestone/dolomite
conglomerate. The clasts are primarily composed of reddish-gray limestone, highly
elongate and have length-to-width aspect ratios >10:1. In addition, clasts display a wide
range of spatial orientations with inclinations between 0 and 90° compared to bedding
(Figure 2.3D-E). The matrix consists of buff-colored dolomicrite and interlocking
microspar mosaic. The thickness of the FPC component is variable and ranges from
completely absent to ~ 1 m thick. Lateral examination reveals that the FPC exhibits a
lensoidal morphology. In rare cases, the upper laminated facies are interrupted by
relatively thin intervals of FPC.

Laminated Component: The uppermost member of the marker bed consists of finely-
laminated, purple-gray to buff-gray micritic dolomite and limestone (Figure 2.3F).
Laminae thickness and purple coloration tend to decrease upward such that lamination is
difficult to distinguish in outcrop in the uppermost centimeters. The lamination exhibits a
hummocky-cross stratified to tangentially-cross stratified fabric and is defined by
horizontal distributions of opaque inclusions or by differences in crystal size.

50
Figure 2.5: Photomicrographs showing the variability in textures exhibited by the marker bed

A) Ooids of CR-1 with sharp boundaries and laminae. B) Ooids of CCv-1 with blurred boundaries and laminae. Notice ghost-like
appearance. C) Well defined fans of CR-2. D) Doming, coarse-grained structures of CC-1. E) Fan-free, psuedospar mosaic of CCv-
1.

51
In study sites CR-1, CR-2 and CC-1, the laminated facies contain upwardly domed
structures with fan-like morphology (Figure 2.4B-C, 2.5C). These “fans” consist of ~100
to 1500 µm tall, radiating crystals with square terminations (Figure 2.4D). In rare cases,
clusters of opaque inclusions accumulate near basal portions of larger crystal projections.
Bedding-parallel sectioning exposes crystal tops with pseudohexagonal morphologies
(Figure 2.4E). Fans tend to occur in high abundance along particular horizons; however
isolated fans also exist. In samples from CC-2, clear square terminating crystals are
absent, however similarly sized and shaped upward-doming features are distinguishable.
These features are comparatively coarse crystalline (crystals up to 100 µm in diameter)
and are more discontinuous in nature (Figure 2.4F, 2.5D). Samples from CCv-1 do not
exhibit fans and the lamination is coarse crystalline and difficult to recognize in thin
section. Crystals in CCv-1 range in size from ~50 to 200 µm in diameter and are
arranged in an interconnected mosaic (Figure 2.5E).

Geochemistry
Carbon and Oxygen Isotopes: All five sample sites display similar trends in carbon and
oxygen isotopic values (Figure 2.6). Oxygen isotopic compositions begin at ~ –8‰
(VPDB) in the basal oolite and decrease to a minimum of ~ –13‰ in the uppermost
laminated component. Carbon isotopes show a similar trend except minima are typically
exhibited near the basal or middle portions of the laminated component. Minimum δ
13
C
values in all sections fall below –8‰ (VPDB). δ
13
C
carbonate
values from CR-1 and CR-2

52
Figure 2.6: Geochemical data shown stratigraphically for all five marker bed sections.
Refer to Figure 2.2 for stratigraphic key. Three fan icons indicate abundant fan
groupings. A single fan icon indicates rare fan groupings. A single fan icon with an
overlying R indicates recrystallized fan horizons. Strontium concentrations are elevated
10X to enhance visualization of stratigraphic trends. Mg/Ca diagrams are separated as
follows: light gray = Mg/Ca ratio consistent with low Mg calcite (LMC), gray = high Mg
calcite (HMC), far right of dark gray = dolomite (DOL; Mg/Ca = 1). Notice similarity in
stratigraphic trends of δ
13
C
carbonate
and δ
18
O
carbonate
and lack thereof in δ
34
S
CAS
. The
negative trend in δ
13
C up section is likely the Wonoka-Shuram excursion, recognized in
~580 Ma carbonates elsewhere (refer to Figure 1.2 for δ
13
C from units above and below
the marker bed).

53
Figure 2.6 continued

54
Figure 2.6 continued

55
Figure 2.6 continued

56
Figure 2.6 continued

57
Figure 2.6 continued

58
include the lowest measured, with minima as low as –9.5‰. δ
13
C
carbonate
and δ
18
O
carbonate
show moderate positive correlation (Figure 2.7).

Elemental Concentration: Stratigraphic analysis of the trace elements Sr, Mn and Fe
(referred to as Sr
carb
, Mn
carb
and Fe
carb
, respectively) and major elements Ca and Mg
reveals general similarities among the five sample sites (Figure 2.6). Strontium
concentrations are low and range from 80 to 208 ppm in the basal oolite component and
increase to maximum concentrations between 134 and 323 ppm in the middle to
uppermost portions of the laminated component. The higher resolution sample sites (CR-
1 and CR-2) show a pronounced increase in Sr
carb
coincident with the transition into the
laminated member. Sample site CCv-1 does not exhibit an increase in Sr
carb
and instead
concentrations decrease slightly up section. In all sections Mn
carb
concentrations fall
between 289 and 2,636 ppm. CR-1 displays elevated Mn
carb
concentrations (between
1,139 and 2,636 ppm) within the FPC component and a sharp decline (from 1,918 to 982
ppm) coincident with the transition into the laminated component. A less pronounced
increase is evident at 120 cm in the FPC component of CR-2, where Mn
carb

concentrations reach 2,250 ppm.
In general Fe
carb

concentrations decrease up section with the exception of CCv-1.
Sample sites CR-1 and CR-2 show pronounced increases in the FPC component. Basal
Fe
carb
concentrations are typically high (except CR-1), with values >3,000 ppm and up to
5,582 ppm. Concentrations are lower in the laminated component (with the exception of
the uppermost CCv-1 data point), with values between 449 and 1,594 ppm.

59
Figure 2.7: Cross plot of δ
13
C
carbonate
and δ
18
O
carbonate
for all marker bed samples

60
Aside from CCv-1, Mn/Sr ratios are relatively low in the upper laminated
component (between 0.7 and 5.4). Sample sites CR-1, CR-2 and CC-1 show a
pronounced increase near the middle of the section coincident with FPC (CR-1 and CR-
2) or lower laminated (CC-1) members. CCv-1 Mn/Sr values increase from 4.5 in the
basal oolite to 18.2 in the laminated member.
The lowermost samples of each section yield relatively high Mg/Ca (molar)
ratios, approaching 1, the value of stochiometric dolomite. These values generally
decrease up section and laminated facies exhibit ratios near zero. One exception is site
CCv-1 where the uppermost data point has a Mg/Ca of 0.47.

Carbonate-Associated Sulfate and Pyrite: CAS concentrations are highly variable
among the different stratigraphic sections (Figure 2.6 and 2.8). Sites CC-2 and CCv-1
exhibit extremely low values, falling below 30 ppm. CR-1 CAS concentrations show a
minimum near the basal FPC component and elevated concentrations of ~400 ppm in the
basal and uppermost horizons of the section. CR-2 displays high CAS concentrations
(~500 ppm) in the upper portion of the laminated component and uppermost
oolite/lowermost FPC (~300 ppm), and low values (<49 ppm) elsewhere. CC-1 exhibits
the highest CAS values, with the laminated component containing CAS in excess of
1,000 ppm.
Pyrite concentrations are low with all samples containing less that 0.013 wt%
(Figure 2.9).

61
Figure 2.8: Composite δ
13
C
carbonate
, δ
34
S
CAS
and CAS concentration curves for the five
Caborca study sites and those of the Rainstorm carbonates (DV)

Here, sulfur curves are time-correlated via δ
13
C (same-sample geochemical data plotted
in stratigraphic relation to those from other localities based on δ
13
C). Notice the
moderate variability in δ
34
S
CAS
and high variability in CAS concentration. δ
34
S
CAS
and
CAS concentration are significantly reduced and elevated, respectively, compared to
underlying and overlying carbonates at Cerro Rajon.

62
Sulfur Isotopic Composition: The sulfur isotopic composition of CAS varies
moderately within individual and among different sections (Figure 2.6). In total, δ
34
S
CAS

values range from +18.6 to +27.6‰ (VCDT). Correlation of individual data points via
δ
13
C shows that mid-section samples show less deviation in δ
34
S
CAS
between sections
compared to samples from the uppermost and lowermost portions of the marker bed
(Figure 2.8). Lowermost samples range from +18.9 (CR-1) to +27.6‰ (CC-1), mid-
section samples range from +21.5 (CR-2) to +23.3‰ (CR-1) and uppermost samples
range from +20.6 (CR-1) to +25.5‰ (CC-1). Sample site CC-2 did not contain sufficient
CAS to perform isotopic analyses and only one sample from CCv-1 yielded sufficient
CAS, thus stratigraphic trends could not be analyzed at these sites.

DISCUSSION
Pyrite Oxidation and the Reliability of CAS in Carbonates
It has been demonstrated by Marenco et al. (2008) that oxidation of pyrite during
the CAS extraction procedure can alter primary values of both CAS concentration and
δ
34
S
CAS
. The low pyrite concentrations in marker bed carbonates (all samples <0.013
wt.%) make alteration of this sort unlikely. In addition, CAS concentration and δ
34
S
CAS
do not show positive and negative correlations with pyrite concentration, respectively
(Figure 2.9A-B); trends expected to occur if pyrite oxidation were the source of sulfate
during sample processing (Marenco et al., 2008). Thus, while pyrite oxidation can
obscure CAS data in carbonates, it does not significantly alter those of the Clemente
marker bed.

63
Figure 2.9: Assessing pyrite oxidation in marker bed samples

A) Cross plot of pyrite concentration and CAS. Contamination by pyrite oxidation would be expressed as a positive correlation.
B) Cross plot of pyrite concentration and δ
34
S
CAS
. Contamination by pyrite oxidation would be expressed as a negative correlation.

64
Diagenetic Indicators
Trace element concentrations of Sr
carb
, Mn
carb
and Fe
carb
are traditionally used to
assess the degree of meteoric alteration in carbonate systems due to the divergent
partition coefficients of these elements, their relative concentrations in marine and
meteoric fluids and their affinity to the carbonate crystal lattice (Bodine et al., 1965;
Kinsman, 1969; Turekian, 1972; Brand and Veizer, 1980). Progressive meteoric
alteration of originally marine carbonates leads to a decrease in Sr
carb
and increases in
both Mn
carb
and Fe
carb
concentrations (Brand and Veizer, 1980). In addition, meteoric
alteration can produce decreases in both carbon and oxygen isotopic compositions
(Turekian, 1972; Allan and Mathews, 1982; Banner and Hanson, 1990).
Textural changes are also expected during progressive meteoric alteration. In
general, crystal size increases with increasing recrystallization such that the
crystallographic progression should follow the trend of micrite to microspar to
psuedospar to spar in carbonates experiencing meteoric alteration (Brand and Veizer,
1980). This crystal growth has been termed aggrading neomorphism (Folk, 1965;
Bathurst, 1975) and has been recognized in many carbonate systems. Crystal growth of
this nature tends to produce large crystals at the expense of smaller ones. In addition,
intricate textural features can be destroyed upon recrystallization. However, reports of
Precambrian carbonates demonstrate that recrystallization can preserve some textures;
through a process known as mimetic recrystallization (Tucker, 1983; Sibley, 1991;
Zempolich and Baker, 1993; Corsetti et al., 2006).

65
Recent work by Gill et al. (2008) provides a first step toward understanding the
effects of meteoric diagenesis on CAS proxies (concentration and δ
34
S
CAS
). In this study,
the authors show that meteoric recrystallization of Pleistocene coralline aragonite to low-
Mg calcite is accompanied by significant decreases in δ
18
O
carbonate
, Sr, Na and CAS.
δ
34
S
CAS
values are relatively invariant between primary and recrystallized phases. These
findings led the authors to conclude that meteoric recrystallization can cause a reduction
in CAS concentration but that δ
34
S
CAS
is relatively robust and retained in the carbonate.

Marker Bed Diagenesis and CAS Concentration
CAS concentration exhibits negative correlations with Fe
carb
and Mn/Sr (Figure
2.10A-B). These trends are consistent with meteoric alteration of an initially marine-
precipitated carbonate as meteoric waters will be relatively depleted in sulfate and
strontium and enriched in iron and manganese compared to marine waters. These results
are in broad agreement with those of Gill et al. (2008) in that CAS concentrations
decrease upon meteoric recrystallization.
Samples and sections exhibiting increased textural maturity (i.e., increasing
crystal/grain size) and lacking well-defined intricate features (crystal fans and lamination)
exhibit low to negligible CAS concentrations. Specifically, CC-2 and CCv-1 display
CAS values below 30 ppm and do not contain well-pronounced crystal fan groupings.
CC-2 contains “fan-like” features (Figures 2.4F and 2.5D), however these are
characterized by a coarse-crystalline, anhedral, mosaic, likely indicative of aggrading
neomorphism. Fan blades do not exhibit well-pronounced, straight edges as in CR-1,

66
Figure 2.10: Cross plots of CAS and traditional proxies of meteoric diagenesis

A) Fe
carb
and B) Mn/Sr. Notice how samples with high CAS correspond to decreased
Fe
carb
and Mn/Sr. Samples with elevated Fe
carb
and Mn/Sr and reduced CAS are
interpreted as preferentially altered by meteoric fluids.

67
CR-2 and CC-1, also indicative of recrystallization. CCv-1 contains no fans or fan-like
features and lacks a well-defined lamination (Figure 2.5E). The lamination can been seen
in outcrop, however lamina transitions are obscure and boundaries diffuse. In
photomicrograph the laminated facies are composed of psuedospar, likely reflecting
aggrading neomorphism of a finer-grained precursor. In addition, the oolitic component
of CCv-1 is composed of a pseudospar mosaic and ooids are identifiable but appear
ghost-like (Figure 2.5B), suggestive of mimetic recrystallization.
The above observations indicate removal of CAS during meteoric
recrystallization in agreement with the findings of Gill et al. (2008). Elevated Mn/Sr and
Fe
carb
and increased textural maturity show strong correlation with reduced CAS
concentrations. The overall low concentration of Sr
carb
(142-374 ppm) and high
concentrations of Mn
carb
(289-2,636 ppm) and Fe
carb

(486-5,699 ppm) are indicative of
meteoric diagenetic influence in all components of the marker bed. The samples with the
lowest Mn/Sr ratios and Fe
carb
and highest CAS concentrations should thus exhibit a
geochemistry closest to the primary marine depositional conditions.

Diagenetic Effects on δ
34
S
CAS

The isotopic composition of CAS is moderately variable among the sample sites.
δ
34
S
CAS
shows no appreciable correlation with Fe
carb
, Mn/Sr or CAS concentration
(Figure 2.11A-C), further supporting the hypothesis that meteoric diagenesis does not
affect the sulfur isotopic composition of CAS (Lyons et al., 2004; Gill et al., 2008).
Incidentally, diagenetic modification (removal) of CAS concentration can obscure an

68
Figure 2.11: Cross plots of δ
34
S
CAS
and traditional proxies of meteoric diagenesis

A) Fe
carb
and B) Mn/Sr, and C) CAS concentration. Notice the lack of significant
correlation in all three.

69
initial covariation with δ
34
S
CAS
, introducing complications when attempting to interpret
the primary depositional regime. Ultimately, the variability in δ
34
S
CAS
in marker bed
sites may result from primary local heterogeneity in seawater δ
34
S
sulfate
or later, non-
meteoric diagenetic modification.

Local Seawater Sulfate and δ
34
S

Although diagenetic overprinting removes CAS, it is encouraging that samples
from Neoproterozoic units retain at least some residual CAS as this implies that
Neoproterozoic seawater contained appreciable sulfate to be preserved in solid-phase
carbonate. Deriving precise sulfate concentration in Neoproterozoic seawater (liquid
concentration, mol/L) is impossible given only CAS concentration in parts per million
(solid concentration g/10
6
g), especially without well-known partition coefficients.
However, one can comment on relative changes in concentrations of seawater sulfate
using both CAS concentration and δ
34
S
CAS
to develop reasonable conclusions concerning
secular ocean chemistry.
CAS concentrations within the marker bed at CR-1, CR-2 and CC-1 are
significantly elevated compared to units stratigraphically above and below (see previous
chapter). Non-marker bed Neoproterozoic carbonates of Cerro Rajon exhibit CAS
concentrations <200 ppm, in striking contrast to marker bed values of up to 1,200 ppm.
A similar increase in CAS is observed in time equivalent carbonates of the Rainstorm
Member (Johnnie Formation) in the Death Valley region (Figure 2.8; Kaufman et al.,
2007). The temporal equivalence inference is justified because the Rainstorm carbonates

70
also yield carbon isotopic values consistent with the W-S excursion. Incidentally, these
Rainstorm carbonates also contain similar lithologic transitions (oolite overlain by finely-
laminated carbonate) and enigmatic fabrics (FPC and formerly aragonitic crystal fans;
Corsetti and Kaufman, 2003; Pruss et al., 2008). The increased CAS in ~580 Ma Death
Valley and Caborca carbonates is consistent with a transient increase in seawater sulfate.
The spatial scale of this increase is difficult to constrain and extrapolation to a global
increase in marine sulfate is far from justifiable given the lack of CAS concentration data
from additional, more distant localities. However, the CAS increase in Caborca and
Death Valley carbonates does indicate some degree of continuity over ~800 km of lateral
distance (or less, if one accepts the Sonora-Mojave megashear hypothesis which suggests
that Caborca and Death Valley were subsequently displaced from one another along a
major, now obscured, fault zone; c.f., Anderson and Silver, 2005; Stewart, 2005).
Sulfur isotopic values of CAS exhibited by the Clemente marker bed show
differing stratigraphic trends (Figure 2.6 and 2.8). Section CC-1 has a δ
34
S
CAS
trend
similar to the Rainstorm Member carbonates of Death Valley, with isotopic values
decreasing ~7-10‰ near the middle to upper portions of the section (Kaufman et al.,
2007). CR-2 also exhibits a mid-section decrease in δ
34
S
CAS
, however a less severe
increase of ~3‰ occurs in the lower 60 cm of the marker bed. Finally, section CR-1
displays an overall increase in δ
34
S
CAS
with a mid-section maximum of ~23‰. The
deviation among the Clemente marker bed sections demonstrates that local variability in
δ
34
S
CAS
exists. This is in striking contrast to trends recognized by Hurtgen et al. (2006)
in cap carbonates of the Maieberg Formation (Marinoan equivalent; ~635 Ma). In these

71
units the authors report lateral agreement in δ
34
S
CAS
across ~180 km of distance, however
they also report a significant change in δ
34
S
CAS
when compared to an additional section
200 km to the east of their N-S transect. The younger (~580 Ma) carbonates of the
Clemente marker bed show significant variability over a mere 20 km of lateral distance,
suggestive of localized controls on seawater δ
34
S or diagenetic resetting of an initial
value.

CAS and δ
34
S
CAS
of Clemente and Rainstorm Carbonates: A New Interpretation
Despite noticeable variability in δ
34
S
CAS
among the Clemente marker bed sites,
sections from Cerro Rajon exhibit sulfur isotopic values that are up to ~8‰ depleted
compared to carbonates stratigraphically below (~26‰) and above (~28‰) (see Figure
1.4). This decreased isotopic value and an increase in CAS of ~1000 ppm, suggest an
input of isotopically light sulfate to or a decrease in its removal from (i.e., diminished
BSR and pyrite burial) local seawater. The ultimate extent of this input/output is poorly
constrained, however given the broad similarities between the Clemente marker bed and
the Rainstorm carbonates of Death Valley, this perturbation could have impacted ~800
km of Neoproterozoic shelf and developed from regional, but not global (as discussed
below), oceanic conditions or diagenetic modification.
The data presented here and by Kaufman et al. (2007) can be explained by
localized changes in depositional environment (see Figure 2.12). Under restrictive,
evaporitic conditions, salinity and alkalinity increase, producing expectable changes in
the geologic record. Concentration of seawater through evaporation will increase both

72
Figure 2.12: Diagrammatic evolution of the Caborca region ca. 580 Ma. Time progresses
alphabetically from A to D and the depositional environment, stratigraphic progression
and chemical evolution are shown for each time slice. The stratigraphic progression
indicates a deepening sequence consistent with transgression. Flat pebble facies are
laterally discontinuous and not shown here. In order to maintain basin restriction during
transgression, an outboard barrier must grow vertically with sea level rise, perhaps by
carbonate accretion as shown. Upon establishment of a restricted basin, sulfate and
alkalinity (and strontium as discussed in text) increase as a result of evaporation
concentration. Low iron limits removal of sulfide as pyrite and as a result nearly all
sulfide is reoxidized to sulfate. As a result, the basin becomes progressively enriched in
sulfate that is depleted in δ
34
S relative to the open ocean; these values are recorded in
precipitating carbonate. Ultimately, alkalinity increases significantly yielding
precipitation of seafloor fans (C). In D, the lithologic transition into shale indicates
further deepening, and the lack of carbonate prohibits continued analysis of CAS,
δ
34
S
CAS
, etc. Even if the restricted basin is entirely cut off from the ocean, global δ
13
C
communication can be maintained through atmospheric CO
2
isotopic equilibrium.

73
Figure 2.12 continued

74
strontium and sulfate, yielding carbonates elevated in trace amounts of both (trends
recognized in fan-bearing carbonates of both Caborca and Death Valley; Vogel et al.,
2002; Hurtgen et al., 2004). Increased evaporation can also produce high alkalinity
waters, perhaps sufficient to reach a saturation threshold conducive to aragonitic seafloor
fan precipitation. Although not recognized here, an increase in δ
18
O should accompany
the transition to an evaporitic regime. However given the many factors (both primary
and diagenetic) that can influence oxygen isotopic compositions in carbonates, it is
difficult to evaluate the δ
18
O signal. In particular, diagenetic overprinting can alter δ
18
O
values of carbonate rocks and obscure initial isotopic compositions.
In order to produce all of the observed trends, restrictive evaporation must be
coupled with iron limitation. Under iron-limiting conditions, sulfide produced via
bacterial sulfate reduction (BSR) would not be sequestered as sedimentary pyrite and
instead be re-oxidized to sulfate (assuming availability of an external oxidant). If sulfide
burial does not occur (in the form of pyrite), seawater δ
34
S
sulfate
would decrease toward
the isotopic value of the riverine input (δ
34
S
riverine
). Additionally, a lack of pyrite burial
would lead to increased sulfate concentrations and decreased pyrite concentrations, both
exhibited by the most unaltered marker bed carbonates.
These interpretations require specific initial conditions for the Caborca region and
the Neoproterozoic ocean during this time interval. In order to develop local
perturbations in δ
34
S
sulfate
, the initial sulfate reservoir must have been low, a condition
supported by the relatively low concentrations of CAS in Neoproterozoic carbonates. In
addition, low sulfate concentrations must have been sustained by continuous removal via

75
BSR in order to account for a decrease in δ
34
S
sulfate
upon initiation of iron-limiting
conditions. Finally, these inferences imply that pyrite burial was a sufficient sink for
32
S
during most of the Neoproterozoic.
When combined, the data suggest that Death Valley and Caborca were
experiencing similar depositional conditions at ~580 Ma. Both locations were either part
of one continuous, restrictive basin or perhaps individual, isolated basins encountering
similar processes. This is geologically reasonable, given the fact that the western edge of
North America was transitioning from a rifted to passive margin at this time (e.g., Levy et
al., 1994). The Clemente-Rainstorm basin never produced evaporites, but it may have
been somewhat isolated from the global ocean at the time, perhaps similar to the Red Sea,
today. The local deviation in δ
34
S
CAS
in the Clemente marker bed and Rainstorm
carbonates can be explained by differences in the degree of iron limitation or by local
deviations in δ
34
S
riverine
. Variations in CAS concentrations may reflect differences in
localized sulfate increases (based on iron limitation and concentration by evaporation)
and/or differences in diagenetic removal. Strong correlation with Fe
carb
would suggest
influence by the latter, as discussed above, such that those samples with the highest CAS
are most reflective of the primary environment. Ultimately, the interpretation given above
does not require a global increase in O
2
and is thus more parsimonious in light of the new
data and decoupled δ
13
C

and δ
34
S.

76
BROAD IMPLICATIONS: GLOBAL HETEROGENEITY IN δ
34
S
CAS

Using the W-S as a global tie point, we have plotted the known δ
34
S records from
Oman, China, Death Valley, and Caborca. The results are somewhat striking, in that none
of the data follow the same trend and magnitude. The lack of agreement in the absolute
value of δ
34
S
CAS
in excursion facies of Oman and South China (Figure 2.13) indicate that
the Neoproterozoic marine realm was heterogeneous with respect to δ
34
S
sulfate
, in contrast
to δ
13
C. The discrepancy between δ
34
S

and δ
13
C in ~580 Ma carbonates suggests that the
two were somewhat decoupled, such that models explaining both parameters must allow
for global homogeneity of δ
13
C

and heterogeneity of

δ
34
S
CAS
, or diagenetic overprinting
of δ
34
S
CAS
but not of δ
13
C. Although diagenesis has not been conclusively identified as a
δ
34
S
CAS
modifying process (c.f., Lyons et al., 2004; Gill et al., 2008) it seems likely that
diagenetic processes could preferentially affect δ
34
S
CAS
versus δ
13
C
carbonate
, given the
rock-buffering tendencies of carbonates (Banner and Hanson, 1990). Nevertheless, the
data presented here imply heterogeneity in δ
34
S
CAS
and modify our understanding of
Neoproterozoic oceanic δ
34
S
sulfate
. In addition, it is evident that δ
34
S
CAS
cannot be used as
a chemostratigraphic correlation tool in Neoproterozoic carbonates.
Li et al. (2010) report δ
34
S
CAS
variability across a depth transect in the late
Neoproterozoic succession of southern China. The authors propose that this variability
arises from a stratified ocean basin characterized by a high sulfate shallow ocean and a
low sulfate deep ocean. While the environmental interpretation drawn for the marker bed
carbonates involves isolation from the open ocean, a similarly stratified basin may have
existed in the Caborca and Death Valley marker bed depocenters. Stratigraphic variation

77
Figure 2.13: Global comparison of δ
34
S
CAS
data

Correlation determined via δ
13
C. Notice how data from Oman (Fike et al., 2006) and
South China (McFadden et al., 2008) do not agree well with Caborca (this study), Death
Valley (Kaufman et al., 2007) nor one another, indicative of heterogeneity in oceanic
δ
34
S
sulfate
in the Neoproterozoic.

78
in δ
34
S
CAS
may have arisen due to chemocline fluctuation and the associated changes in
the availability of dissolved sulfate and its isotopic composition. Regardless, the
observations of Li et al. (2010) provide an alternate mechanism for the local δ
34
S
CAS

heterogeneity in ~580 Ma carbonates.

CONCLUSIONS
A locally extensive, carbonate marker bed of the Clemente Formation, Caborca,
Mexico exhibits high variability in CAS concentration and moderate variability in
δ
34
S
CAS
over ~20 km of lateral distance. Negative correlation of CAS concentration with
Mn/Sr ratios and Fe
carb
concentrations in all samples is consistent with removal of CAS
during meteoric recrystallization. Given meteoric diagenetic removal, samples exhibiting
high CAS are likely reflective of the most primary values. δ
34
S
CAS
does not show
correlation with these traditional proxies for meteoric diagenesis, suggesting an alternate
source of variability, possibly arising from local variations in seawater δ
34
S
sulfate
or from
alternate modes of diagenetic alteration.
As a whole, marker bed sulfur systematics are similar to those from fan-bearing
carbonates of the Johnnie Formation, Death Valley, California. Both sites show an
increase in CAS concentration and a decrease in δ
34
S
CAS
compared to overlying and
underlying carbonates. These trends, in addition to increased Sr
carb
, decreased pyrite and
the presence of seafloor precipitates, are consistent with an evaporitic depositional regime
in which iron is limited.

79
The marker bed of Caborca and equivalent units from Death Valley contain a
high-magnitude negative δ
13
C excursion likely equivalent to the ~580 Ma Wonoka-
Shuram event. Expression of this excursion allows comparisons to be made to other
excursion-bearing units worldwide. Similar trends in and values of δ
34
S
CAS
are not
recognized from temporally equivalent units from South China or Oman, suggesting that:
1) the Neoproterozoic oceans were not globally homogeneous with respect to δ
34
S
sulfate
,
2) some of these localities were restricted environments and thus do not record “open
ocean” conditions or 3) diagenesis has preferentially altered some or all locations. This
study highlights the global heterogeneity in Neoproterozoic δ
34
S
CAS
, which must be
considered in order to characterize oceanic δ
34
S
sulfate
, and demonstrates that unlike δ
13
C,
δ
34
S
CAS
cannot be used as

a correlation tool.

80
CHAPTER 3: CARBONATE-ASSOCIATED SULFATE, δ
34
S AND δ
13
C
ANALYSES OF DOLOMITE CONCRETIONS OF THE MIOCENE MONTEREY
FORMATION: INSIGHTS INTO FORMATION ENVIRONMENTS

CHAPTER 3 ABSTRACT
Dolomite concretions are a significant component of the Miocene Monterey
Formation, California. Concretion growth is commonly thought to proceed as organic
matter is progressively degraded in the subsurface, increasing pore water alkalinity.
Principally, sulfate reduction is hypothesized as the dominant organic matter
remineralization pathway and is interpreted as the primary concretion-producing
mechanism. Previous studies have used carbon and oxygen isotope signatures to deduce
pore water conditions during concretion growth; however these proxies remain equivocal
with respect to the primary organic matter degradation pathways in marine sediments and
do not help distinguish diagenetic zones dominated by organic matter oxidation by
oxygen, nitrate, metal oxides and sulfate, and degradation by thermal decarboxylation.
Here we employ carbonate-associated sulfate (CAS) and δ
34
S
CAS
in order to more
uniquely determine the mechanisms of concretion authigenesis in the Monterey
Formation as well as identify contemporaneous diagenetic reactions.
Combining the CAS concentration, δ
34
S
CAS
, and δ
13
C reveals that concretions
from the Monterey Formation formed above, within, and below the zone of sulfate
reduction, depending on locality. One nodular concretion from the Phosphatic Shale
Member at Naples Beach yields δ
34
S
CAS
near Miocene seawater sulfate (~22‰),

81
relatively high CAS (ca. 1000 ppm) and depleted δ
13
C, values consistent with shallow
formation in association with organic matter degradation by oxygen, nitrate and/or metal
oxides, and only minor sulfate reduction. Cemented, concretionary layers of the
Phosphatic Shale Member at Shell Beach display elevated δ
34
S
CAS
(up to ~37‰), CAS
concentrations ca. 600ppm and mildly depleted δ
13
C (ca. –6‰), indicative of formation
in sediments influenced by sulfate reduction. Finally, the Siliceous Member concretions
of Montana de Oro and Naples Beach show depleted δ
34
S
CAS
(less than Miocene
seawater), low CAS concentrations and positive δ
13
C values, consistent with formation in
sediments experiencing sulfide oxidation as well as methanogenesis. Combining CAS
analysis with more traditional techniques reveals that concretion formation occurs across
a range of diagenetic conditions, and may represent the first reported occurrence of the
importance of sulfide oxidation in the formation of some concretions.

INTRODUCTION
Concretions, preferentially-cemented regions within sediments or sedimentary
rocks, are recognized all over the world in units of nearly all ages and depositional
environments (Mozley and Burns, 1993), including deposits of the Miocene Monterey
Formation (see locality map Figure 3.1 and Figure 3.2). Despite common occurrence,
little is known about the precise mechanisms responsible for concretion precipitation in
sediments. Outcrop relationships indicate that concretions are diagenetic in origin and
form largely through passive cement precipitation in open pore space, although some

82
Figure 3.1: Map showing Monterey sample site locations

83
authors have suggested invasive cementation (i.e., Mozley, 1989; Raiswell and Fisher,
2000). In many cases the cementing mineral is carbonate in composition, and based on
stable isotopic analyses (primarily δ
13
C) it is clear that the degradation of organic
compounds plays a role in authigenesis (i.e., Claypool and Kaplan, 1974). Alone, δ
13
C
analysis cannot distinguish among the primary organo-diagenetic environments
encountered in marine sediments. As Figure 3.3 demonstrates, multiple horizons exhibit
negative δ
13
C values, including those of oxygen, nitrate, metal oxide and sulfate
reduction, anaerobic oxidation of methane (AOM) and abiotic thermal decarboxylation.
In fact, only positive δ
13
C values signify concretion growth in a particular organo-
diagenetic zone (methanogenesis), thus in order to distinguish among the other
mechanisms another proxy must be utilized.
Here we use carbonate-associated sulfate (CAS), trace sulfate that is incorporated
into a carbonate mineral upon precipitation, as a diagenetic indicator in the concretions
and cemented layers of the Monterey Formation. Sulfate substitutes for the carbonate ion
within the crystal lattice (Pingitore et al., 1995). Although trace in quantity, sufficient
CAS is present in natural carbonates to extract for measurement of its concentration and
its isotopic composition (δ
34
S
CAS
) (Burdett et al., 1986). CAS concentration and δ
34
S
CAS

are related to the dissolved sulfate concentration and δ
34
S
sulfate
of carbonate-precipitating
fluids, respectively (Burdett et al., 1989) and therefore carbonates act as a geologic
reservoir for dissolved sulfate.
Since CAS mimics dissolved sulfate, one can use modern sulfate profiles to
predict how CAS should behave in a pore water environment (see Figure 3.4 for pore

84
Figure 3.2: Concretions and stratigraphic column of the Monterey Formation

Concretions come from the siliceous member at Montana de Oro (A and B), the phosphatic shale member at Naples Beach (C) and the
phosphatic shale member at Shell Beach (D). Also included is a stratigraphic column of the Monterey Formation (E).

85
water sulfate characteristics). In diffusion-dominated marine sediments, sulfate is
removed with increasing depth as a result of microbial sulfate reduction (Jorgensen,
1983; Berner, 1984; Canfield and Thamdrup, 1994; Canfield, 2001). As sulfate is
progressively depleted, the residual pool of sulfate becomes isotopically enriched as a
result of preferential reduction of
32
SO
4
(and subsequent sulfide burial as pyrite) by
microbes (Berner, 1984; Canfield and Thamdrup, 1994; Canfield, 2001). Ultimately,
sulfate is depleted to negligible concentrations at depth and sulfate reduction can no
longer be supported. These characteristics are applied to CAS in concretions such that,
for example, those exhibiting reduced CAS concentrations and elevated δ
34
S
CAS

(compared to primary marine-precipitated carbonate) are most reasonably interpreted to
have precipitated in the sulfate reduction zone (Figure 3.3). As Figure 3.4 demonstrates,
the depth scale of sulfate removal can vary over orders of magnitude from centimeters to
hundreds of meters, thus the depth axis in Figure 3.3 is displayed as relative depth.
When CAS is used in conjunction with δ
13
C, each of the major organo-diagenetic
environments exhibit a unique signature (Figure 3.3) and the concretion formation
mechanism(s) can be identified. In this study, we analyzed concretions from the
Phosphatic Shale (PSM) and Siliceous Members (SM) of the Miocene Monterey
Formation in order to determine which organo-diagenetic environment(s) can be
attributed to concretion growth.

86
GEOLOGIC CONTEXT
The Monterey Formation, California is a hemipelagic to pelagic, organic-rich,
largely siliceous deposit dominated by diatomite, porcelanite, chert, dolomite and
phosphatic shale (Schwalbach and Bohacs, 1991; Schwalbach, 1992; Behl and Garrison,
1994) (Figure 3.2E). Outcrops occur at multiple locations along the California coast from
south of Palos Verdes to north of Monterey Bay. The Monterey Formation is an
important source and reservoir rock for petroleum in California (Surdam and Stanley,
1981; Mertz et al., 1983; Isaacs, 1984) and its stratigraphy and paleoenvironmental
setting are well documented (Bramlette, 1946; Isaacs, 1981; Svadra and Bottjer, 1986;
Burns and Baker, 1987; Behl et al., 1991; Schwalbach and Bohacs, 1991; Schwalbach,
1992; Behl and Garrison, 1994; Eichhubl and Behl, 1998; Behl, 1999).
Carbonate concretions and cemented layers are abundant in the Monterey
Formation (Bramlette, 1946; Kushnir and Kastner, 1984; Hennessy and Knauth, 1985;
Compton and Siever, 1986; Burns and Baker, 1987; Eichhubl and Boles, 2000). Carbon
isotopic variation among individual concretions has been well documented from the
Monterey Formation (Kushnir and Kastner, 1984; Hennessy and Knauth; 1985; Burns
and Baker, 1987), making it an excellent candidate in the assessment of the new CAS
diagenetic approach.

METHODS
Concretions and concretionary layers of the Monterey Formation were sampled
from the SM and PSM. SM samples were collected from Montana de Oro State Beach

87
Figure 3.3: Diagram depicting pore water profiles of major compounds in marine sediments and predicted values of δ
13
C, δ
34
S
CAS
and
CAS concentration of concretions precipitated at different relative depths.

Also listed are the primary organic matter degradation horizons with depth in marine sediments. Major compound pore water profiles
from Jorgensen (1983), δ
13
C profile from Claypool and Kaplan (1974) and pore water δ
34
S (and δ
34
S
CAS
) inferred from modern pore
water profiles (Berner, 1984; Canfield and Thamdrup, 1994; Canfield, 2001; Bottcher et al., 2001). *Me-Ox (metal oxide) refers to
Fe- and Mn-oxides. **SMTZ = sulfate-methane transition zone.

88
Figure 3.4: Modern pore water profiles of sulfate concentration and δ
34
S
sulfate

Notice that as sulfate is progressively depleted, the residual pool becomes progressively
enriched in
34
S indicative of bacterial sulfate reduction. Profiles from the Gulf of
California (Goldhaber and Kaplan, 1980), offshore New Zealand (Bottcher et al., 2004)
and offshore Namibia (Dale et al., 2009). Dashed lines correspond to unmodified
seawater δ
34
S
sulfate
(~21‰; coarse dash) and sulfate concentration (~28mM; fine dash).
Left hand depth axis corresponds to Goldhaber and Kaplan (1980) and Dale et al. (2009),
right hand corresponds to Bottcher et al. (2004).

89
and PSM samples were collected from Shell Beach and Naples Beach (Figure 3.1). For
significantly large structures (> ~100cm
3
), multiple samples were analyzed in order to
explore variability within individual concretions. Samples were analyzed for CAS
concentration, pyrite concentration, weight percent carbonate (dolomite), δ
13
C
carbonate
,
δ
18
O
carbonate
, δ
34
S
CAS
and δ
34
S
pyrite
.
Refer to chapter 1 for methods concerning pyrite concentration and isotopic
analyses. Since these concretions likely contain relatively high concentrations of organic
matter, powdered samples were washed twice in the sodium hypochlorite solution. Other
than this difference, the CAS extraction procedure is identical to that outlined in chapter
1. Duplicate measurements of CAS concentration were performed on 8 samples and the
precision of replicates was within +/–10%.

RESULTS
Nature of Concretions
Photographs of the analyzed concretions are displayed in Figures 3.2, 3.5, 3.6 and
3.7 and the latter three pertain to Montana de Oro, Shell Beach and Naples Beach,
respectively. Figure 3.8 contains photomicrographs of the key features recognized from
the different study areas. These sites are discussed in detail below.

Montana de Oro: Dolomites of the Siliceous Member at Montana de Oro consist
primarily of elliptical to irregular concretions that are preferentially distributed along
bedding planes (Figures 3.2A, B, and 3.5A, B). Concretions are pale yellow and

90
Figure 3.5A: Photographs of Montana de Oro concretions MMC1-MMC5

Also shown are the spatial distributions of CAS concentration (ppm), δ
34
S
CAS
(‰) and
δ
13
C (‰) within individual concretions.

91
Figure 3.5B: Photographs of Montana de Oro concretions MMC5-MMC9

Also shown are the spatial distributions of CAS concentration (ppm), δ
34
S
CAS
(‰) and
δ
13
C (‰) within individual concretions.

92
relatively resistant compared to the siliceous host rock. Concretions range from 25 to
60cm thick in cross-section and can be as much as 4.5m across in plan view. Spacing
along bedding planes is fairly regular with ~2 to 5m between individual concretions. In
all instances, external laminae exhibit deflection around concretions (see, for example
Figure 3.2A). These samples have moderately variable cement concentrations, spanning
from 51.9 to 89.9 wt% (Figure 3.9). Nine individual concretions were sampled from this
locality, identified as MMC1-MMC9.
The dolomite cements of the Montana de Oro concretions are composed of
relatively fine-grained mosaics of interlocking anhedral to subhedral crystals (Figure
3.8A). Most crystals are less than 5µm across yet some exhibit diameters up to 20µm.
Larger crystals typically exhibit rhombohedral morphology consistent with a dolomitic
mineralogy (Figure 3.8A). Diatom frustules are common, range up to ~500µm across
and most have been reminerallized to dolomite (Figure 3.8B). Insoluble components
include non-reminerallized diatoms, chert and quartz grains, fine opaque grains
(including pyrite and possibly organic matter) and clays.

Shell Beach: Dolomites of the Shell Beach site dominantly occur as resistant, gray to
brownish-gray, continuous cemented layers (Figure 3.6) within the PSM. Fifteen
individual layers were identified within ~15m of section. Some layers exhibit a pinch
and swell morphology and range in thickness from ~10 to ~75cm. The thickest layer was
sampled three times along a vertical transect (MSC1-3); all others were sample once.
The Shell Beach samples have a wide range in dolomite content, from 18.0 to 89.8 wt%

93
Figure 3.6: Cemented layer photographs and stratigraphic column for the Shell Beach
locality

Cemented layer photographs (A and B) and stratigraphic column (C) for the Shell Beach
locality. A) Upper portion of the stratigraphic section containing samples MSC2-MSC6.
B) Mid section, thick cemented layer (samples MSC1-1 to MSC1-3). As in Figures 3.5A
and B, spatial distributions of CAS concentration, δ
34
S
CAS
and δ
13
C are shown.

94
(Figure 3.9). The discontinuous fabric of the host rock and layering of concretionary
carbonate made identification of laminae deflection difficult. Seven horizons were
sampled from Shell Beach and are identified as MSC1-MSC7.
The cements of Shell Beach concretionary layers are composed of densely
packed, coarse-crystalline dolomite rhombs (Figure 3.8C). Crystals exhibit diameters up
to 120µm and average diameters of ~50µm. In contrast to Montana de Oro and Naples
Beach concretions, Shell Beach concretionary layers contain relatively few microfossils,
dominated by foraminifera and diatoms (Figure 3.8D). Insoluble components include
partially dolomitized phosphate nodules and chert grains, diatom frustules, fine-grained
opaque material (pyrite and perhaps organic matter) and clays.

Naples Beach: Two concretions were sampled from the PSM at the Naples Beach
locality. MNC1 is pale-yellow, irregular and similar in morphology to concretions of the
Montana de Oro site (Figure 3.7). This concretion occurs within the transition zone
between the Lower Shaley and Phosphatic Shale Members of the Monterey Formation,
thus the host lithology ranges from thinly-bedded shale and porcelanite to more thickly-
bedded phosphatic shale. The host rock exhibits minor laminae deflection. Concretion
MNC2 is a white, elliptical nodule, ~80cm across and 30cm thick (Figures 3.2C and
3.7). Along the same bedding plane as MNC2, multiple similarly sized concretions
occur, all of which exhibit clear laminae deflection. MNC2 was sampled ~50m up
section from MNC1 and near the middle of the PSM. The two concretions from the

95
Figure 3.7: Photographs of the two concretions sampled from Naples Beach

Top images correspond to MNC1 and bottom images correspond to MNC2. As in
Figures 3.5A, 3.5B and 3.6, spatial distributions of CAS concentration, δ
34
S
CAS
and δ
13
C
are shown.

96
Naples Beach site contain high concentrations of dolomite cement, ranging from 83.2 to
90.3 wt% (Figure 3.9).
Naples Beach concretion MNC1 is composed of relatively fine-crystalline
dolomite cement. Crystals average ~5µm in diameter and exhibit maximum diameters of
~20µm (Figure 3.8E and 3.8F). Foraminifera and diatom frustules are abundant and are
typically filled with or recrystallized to a coarse-crystalline mosaic. In fact, nearly all
microfossils have been reminerallized to dolomite (Figure 3.8F). Insoluble material
includes partially dolomitized phosphate nodules, fine opaque grains (pyrite and perhaps
organic matter) and clays.
MNC2 is composed of fine-crystalline dolomite. Crystals exhibit an average
diameter of <5µm and a maximum diameter of ~100µm. Patches of relatively coarse-
crystalline cement are disseminated throughout the concretion (Figure 3.8G).
Microfossils comprise between 5 and 10% of thin section area and, in many cases, retain
intricate features such as well-preserved tests (Figure 3.8F). In addition, microfossils are
typically filled with sediment or coarse-crystalline cement.

CAS Concentration, δ
34
S
CAS
and δ
13
C
carb

Analyses of the concretions and concretionary layers yield four primary groupings
with regard to CAS concentration, δ
34
S
CAS
, and δ
13
C
carb
(Figure 3.10A, B and Table 3.1).
These groupings are 1) high CAS [greater than 800ppm], near seawater δ
34
S
CAS
(~22‰:
Paytan et al., 1998), very negative δ
13
C
carb
[less than –9‰], 2) moderate to high CAS
[400 to 800ppm], seawater to high δ
34
S
CAS
[20 to 37‰], moderately negative δ
13
C
carb
[–4

97
Figure 3.8: Photomicrographs of Montana de Oro, Shell Beach, MNC1 and MNC2

Photomicrographs of Montana de Oro (A,B), Shell Beach (C,D), MNC1 (E,F) and MNC2
(G,H). In A, C, E, and G arrows highlight rhombohedral crystals indicative of dolomite
mineralogy. Microfossils from each are shown in B, D, F and G (arrows), and are
dominated by foraminifera and diatoms. Inset in H demonstrates preservation of thin test
walls in MNC2.

98
to –7‰], 3) low CAS [100 to ~500ppm], near seawater to low δ
34
S
CAS
[14 to 20‰],
mildly positive δ
13
C
carb
[1 to 7‰] and 4) very low to negligible CAS [below 70ppm],
near seawater to very low δ
34
S
CAS
[9 to 21‰], highly positive δ
13
C
carb
[9 to 17‰].

Group 1—high CAS, near seawater δ
34
S
CAS
and very negative δ
13
C
carb
: One
concretion (MNC2) falls into this category. CAS concentrations of MNC2 range from
843 to 1443ppm, the highest values recognized in all samples (the 2059ppm MSC5
sample is an outlier). δ
34
S
CAS
values cluster near the value proposed for Miocene
seawater sulfate (~22‰ VCDT: Paytan et al., 1998) and range from 20.1 to 21.8‰.
δ
13
C
carb
values in these samples range from –12.7 to –9.6‰ (VPDB), are the lightest
compositions recognized and are consistently ~4‰ depleted compared to the dolomites
of Shell Beach, which exhibit the next lightest values. Within MNC2, CAS concentration
and δ
13
C decrease while δ
34
S
CAS
increases from the center outward (Figure 3.7).

Group 2—moderate to high CAS, near seawater to high δ
34
S
CAS
and negative
δ
13
C
carb
: This category consists of the samples from Shell Beach. CAS concentrations
are variable and range from 26 to 812ppm. The majority of these samples fall between
~400 and ~800ppm CAS. All but one sample yield elevated δ
34
S
CAS
compared to
Miocene seawater with values ranging from 19.0 to 36.9‰. The carbon isotopic
compositions of Shell Beach samples show a tight grouping and range from –6.6 to –
4.6‰. One concretionary layer (MSC1) was sampled multiple times revealing upward

99
Table 3.1: Monterey geochemical data listed according to site and sample.

100
trends of increasing δ
34
S
CAS
, decreasing δ
13
C and generally decreasing CAS
concentration (Figure 3.6).

Group 3—low CAS, near seawater to low δ
34
S
CAS
and mildly positive δ
13
C
carb
: The
samples from MNC1 fall into this category. CAS in this concretion ranges from 181 to
508ppm and δ
34
S
CAS
ranges from 14.3 to 20.0‰. δ
13
C
carb
values are positive and range
from 1.5 to 6.5‰.

Group 4—very low to negligible CAS, near seawater to very low δ
34
S
CAS
, highly
positive δ
13
C
carb
: This category consists of samples from the concretions of Montana de
Oro. CAS concentrations are very low and range from 0 to 66ppm with the majority of
samples (30/47) containing 0ppm. Samples with very little to no CAS are not amenable
to δ
34
S
CAS
analyses, thus most samples from Montana de Oro do not yield CAS sulfur
isotope data. The samples that contain appreciable CAS yield δ
34
S
CAS
values from 9.0 to
21.3‰, most of which fall between 9.0 and 17.4‰. All of the samples from Montana de
Oro exhibit enriched δ
13
C
carb
values, ranging from 9.5 to 16.6‰.

δ
18
O
carb

δ
18
O
carb
values range from –2.3 to 3.5‰ (VPDB) and show considerable
correlation with δ
13
C
carb
. However, the Naples Beach and Shell Beach samples show
negative correlations whereas the Montana de Oro samples display a strong positive
correlation (Figure 3.11).

101
Figure 3.9: Weight percent dolomite in Monterey samples.

102
Pyrite Concentration and δ
34
S
pyr
Concretion-hosted pyrite concentrations range from 0 to ~1 wt% (similar to non-
concretion hosted disulfide species reported in Zaback and Pratt (1992). Samples from
Montana de Oro and Naples Beach all contain less than 0.2 wt% pyrite while samples
from Shell Beach contain up to 0.94 wt% (Table 3.1). The sulfur isotopic composition of
pyrite dominantly falls between –1 and 7‰, however a small set of samples from
Montana de Oro exhibit values between –22 and –16‰. A solitary data point from
Montana de Oro has an isotopic composition of 20.9‰ (Table 3.1). Pyrite δ
34
S values
also agree well with those of host rock disulfide δ
34
S values of Zaback and Pratt (1992),
which range from –18.5 to +15.0‰.

DISCUSSION
Concretion Formation Environment
Aerobic Oxidation and Minor Sulfate Reduction: The high CAS concentration and
near-Miocene seawater isotopic composition of the Naples Beach concretion MNC2
(Group 1 above) are consistent with shallow precipitation near the sediment-water
interface. Shallow formation is also corroborated by high dolomite content and strong
laminar deflection around the concretion, indicating formation before compaction. The
retention of thin microfossil tests (Figure 3.8H) also supports shallow concretion
formation because significant pre-precipitation compaction would likely destroy such
intricate features.

103
Figure 3.10: Monterey geochemistry cross plots

A) δ
34
S
CAS
vs. δ
13
C; dashed line corresponds to the approximate δ
34
S value of Miocene
seawater sulfate (~22‰). B) CAS concentration vs. δ
13
C.

104
In shallow sediments, organic carbon is ideally oxidized first by dissolved O
2
,
until it is depleted, then oxidation proceeds via reactions with nitrate, and metal oxides
before the initiation of sulfate reduction (Claypool and Kaplan, 1974; Froelich et al.,
1979). Sulfate reduction is the primary sink of sulfate in pore waters, however in
oxygenated sediments sulfate reducers are typically limited (Campbell and Postgate,
1965) and thus sulfate concentrations and δ
34
S
sulfate
are similar to those of overlying
seawater. The δ
34
S
CAS
values of MNC2 range from ~20 to 21.8‰ and are consistent with
precipitation in sediments experiencing organic matter degradation primarily above the
zone of sulfate reduction. Modern, primary marine carbonates incorporate CAS at an
average, yet highly variable, concentration of ~2400 ppm (Lyons et al., 2004; Gellatly
and Lyons, 2005). While this concentration is dependent on variables such as
precipitation rate, mineralogy and the amount of biogenic carbonate, modern concretions
precipitated above the zone of sulfate reduction should contain ~2400 ppm CAS.
Miocene seawater is thought to contain significantly less sulfate (~10 mM: Lowenstein et
al., 2003) than the modern (~28 mM), thus the ~1000 ppm CAS exhibited by MNC2 is
consistent with a near-surface formation depth. The near-seawater δ
34
S
CAS
and high CAS
concentrations suggest formation in sediments dominated by non-sulfate-reduction
remineralization pathways, however the presence of some pyrite within the MNC2
concretionary body indicates at least some sulfate reduction prior to or contemporaneous
with concretion growth. In addition, the progressive increase (although small) in δ
34
S
CAS

and decrease in CAS concentration from center to rim (Figure 3.7B) suggest that the
outer portions of MNC2 formed in sediments with increased (yet still minor) influence by

105
sulfate reduction. The outward trends are consistent with a concentric growth pattern in
which the concretion began growing in sediments experiencing primarily non-sulfate-
reducing conditions then, through progressive burial, impinged on sediments
experiencing sulfate reduction.
It is important to note that oxidation by nitrate and/or metal oxides would produce
CAS trends identical to oxidation by O
2
. Organic matter remineralization by these
compounds is expressed in reactions 3-6 in Table 3.2 (Kuvila and Murray, 1984). In
Monterey sediments, iron oxides and manganese oxides are more probable electron
acceptors particularly because of 1) the sheer size of the MNC2 concretion and 2)
alkalinity considerations. Given that concretions form entirely within sediments, a depth
requirement of at least that equal to the thickness of the concretion must be assumed.
The thickness of MNC2 is ~30cm (see Figure 3.2C) and O
2
would have had to penetrate
to this depth in order to contribute to its formation. Modern, organic carbon-rich
sediments typically exhibit O
2
penetration depths of <<30cm, insufficient to account for a
30cm-thick concretion. Therefore, it seems unreasonable that MNC2 formed entirely by
organic matter oxidation by O
2
.
Whereas oxidation by O
2
increases TCO2, it effectively consumes HCO
3
–

reducing alkalinity and promoting carbonate dissolution (see Eq. 3, Table 3.2). Reactions
involving nitrate, iron oxides and manganese oxides tend to increase alkalinity (Table
3.2) and produce conditions much more favorable for carbonate (and concretion)

106
Figure 3.11: Cross plot of Monterey δ
18
O and δ
13
C

Data separated by sample site and individual concretions or cemented layers. Notice the
strong negative (Shell Beach and Naples Beach) and positive (Montana de Oro)
correlations.

107
Table 3.2: Precipitation, remineralization and other diagenetic reactions

Equations from Kuvilia and Murray (1984), Coleman (1993), Mazullo (2000), Orphan et al. (2004) and Han et al. (2004). Also
included are the change in alkalinity (ΔA), the change in alkalinity per mole of carbon remineralized (ΔA/C), the change in alkalinity
per mole of oxidant consumed (ΔA/Ox) and the change in alkalinity per mole of TCO
2
produced (ΔA/ΔTCO
2
) for each relevant
reaction. Where applicable, reactions involving organic carbon are written in accordance with the Redfield Ratio (Redfield et al.,
1963).

108
precipitation (Redfield et al., 1963; Kuivila and Murray, 1984). However, as equations 1
and 2 demonstrate the increase in alkalinity must be greater than the increase in TCO
2
(in
other words ΔA/ΔTCO2 > 1) in order to drive carbonate precipitation. Interestingly,
reactions involving iron and manganese oxides produce four to eight times the alkalinity
per mole of carbon consumed than any other organo-diagenetic reaction (these reactions
have comparably high ΔA/ΔTCO2 ratios) (see Table 3.2). Therefore, it is more likely that
MNC2 formed below sediments experiencing oxidation by O
2
but above sediments
experiencing extensive sulfate reduction, a determination not possible with previous
techniques.

Sulfate Reduction: Nearly all of the concretionary layers from the Shell Beach locality
(Group 2) exhibit reduced CAS concentrations and elevated δ
34
S
CAS
compared to MNC2
and δ
34
S
CAS
values are generally elevated compared to Miocene seawater. These data
suggest that the concretionary layers of Shell Beach precipitated in sediments that have
experienced considerable sulfate reduction. Relatively high pyrite concentrations (up to
nearly 1 wt%) and negative δ
13
C
carb
values support this interpretation. Carbonate
concentrations are reduced in these samples compared to MNC2 (Figure 3.9 and Table
3.1), consistent with a comparatively deeper horizon of precipitation. Since the zone of
sulfate reduction typically occurs below the zone of aerobic oxidation (Claypool and
Kaplan, 1974), this also bolsters our interpretation. It is important to note that each
location, representing a different time period, is not strictly comparable to other locations

109
in terms of absolute depth, but is compared in terms of relative position within the
diagenetic sequence of Figure 3.3.
Concretionary layer MSC1 exhibits increasing δ
34
S
CAS
, decreasing δ
13
C and
overall decreasing CAS concentration from bottom to top (Figure 3.6). These trends are
consistent with an upward formation direction with increasing depth in the sulfate
reduction zone.
It is evident from Eq. 7 in Table 3.2 that sulfate reduction produces alkalinity
(Kuvila and Murray, 1984) and can create conditions favorable for carbonate
precipitation. When coupled with iron reduction and pyrite formation (Eqs. 8-10), net
ΔA/ΔTCO2 values exceed 1 and yield conditions favorable for concretion precipitation.
Given the abundance of sulfate in seawater, large amounts of organic carbon can be
remineralized via sulfate reduction. In modern marine environments, approximately half
of the sedimentary organic matter is oxidized by sulfate (Thode-Anderson and Jorgensen,
1989). The stoichiometry in Eq. 7 (Table 3.2) indicates that 0.5 moles of alkalinity are
produced per mole of organic carbon remineralized and 1 mole of alkalinity is produced
per mole of sulfate reduced. Therefore, sulfate reduction can account for a tremendous
input of alkalinity to marine sediments and foster extensive carbonate precipitation, given
the correct sedimentary conditions.

Methanogenesis and Sulfide Oxidation: The concretions of Montana de Oro and
MNC1 (Groups 3 and 4) exhibit positive carbon isotopic compositions consistent with
formation in sediments experiencing methanogenesis (Claypool and Kaplan, 1974; Irwin

110
et al., 1977). The preferential incorporation of
12
C into dissolved CH
4
produces residual
pore waters with elevated δ
13
C
DIC
(DIC = dissolved inorganic carbon), thus concretions
forming in these sediments will display similarly enriched δ
13
C
carb
. While the δ
13
C
values clearly indicate concretion formation influenced by methanogenesis, examination
of the CAS and δ
34
S
CAS
provide additional information.
Interestingly, δ
34
S
CAS
values fall below those of Miocene seawater, implying a
source of isotopically depleted sulfate in or near the methanogenic zone. A source of
isotopically depleted sulfate at such a depth may seem counter-intuitive—as sulfate
diffusing from overlying seawater is removed via microbial sulfate reduction, pore water
δ
34
S
sulfate
values should experience enrichment due to Raleigh fractionation effects
(Berner, 1984; Canfield and Thamdrup, 1994; Canfield, 2001). Values significantly
below 22‰ could not have been generated unless another mechanism was operating to
produce dissolved sulfate.
Sulfide oxidation could produce the observed trends. In anoxic sediments, sulfide
produced via sulfate reduction can build to appreciable levels (Jorgensen, 1983) and
diffuse into the underlying methanogenic zone (Figure 3.3). Sulfide produced via sulfate
reduction typically exhibits depleted δ
34
S compositions, again due to the preferential
reduction of
32
SO
4
by microbes. Oxidation of dissolved sulfide does not impart a
significant fractionation such that the produced SO
4
has a similar (and likely depleted)
isotopic composition (Fry et al., 1988). The significant CAS concentrations exhibited by
MNC1 suggest that appreciable quantities of sulfate can be produced by sulfide

111
oxidation. To our knowledge, this is the first reported occurrence of the probable
products of microbial sulfide oxidation recognized in concretionary carbonates.

Who or what may be responsible for sulfide oxidation within or near the
methanogenic zone?: Interestingly, the methane-yielding reactions of Eqs. 12 and 13
(Table 3.2) do not indicate a contemporaneous production of alkalinity (Coleman, 1993;
Mazullo, 2000). This raises questions concerning the likelihood of carbonate
precipitation under strictly methanogenic conditions. However in modern marine
sediments, microbial communities have been found that oxidize methane anaerobically (a
process known as the anaerobic oxidation of methane or AOM) using sulfate as an
electron acceptor (Reeburgh, 1982). Eq. 11 demonstrates that this reaction generates
significant amounts of alkalinity and produces conditions favorable for carbonate
precipitation. Modern concretionary, bedded and irregular carbonate precipitates have
been discovered in association with methane seeps (Orphan et al., 2004), lending support
to AOM as a process capable of promoting carbonate formation.
In addition to increased alkalinity, AOM yields dissolved sulfide which could
potentially be reoxized (perhaps by sulfide oxidizing bacteria, i.e., Thioploca) to “second
generation” sulfate that is isotopically depleted compared to the diffusional, “first
generation” sulfate pool. In fact, putative microfossils resembling modern, filamentous
sulfide oxidizing bacteria have been recovered from the Monterey Formation (Williams
and Reimers, 1983). Under certain conditions, some sulfide oxidizing bacteria use nitrate

112
(Prokopenko et al., 2006) as an electron acceptor and this reaction (Eq. 14, Table 3.2) can
act to further increase pore water alkalinity (through removal of H
+
).
Carbonate precipitates formed via AOM typically exhibit depleted δ
13
C values,
seemingly at odds with the positive values of MNC1 and Montana de Oro concretions.
However, Han et al. (2004) report positive δ
13
C
DIC
in pore waters of Costa Rica margin
sediments experiencing AOM. Although all solid-phase carbonates analyzed in this
study are isotopically depleted, the authors propose that the rapid upward ascent of deep
fluids from the methanogenic zone could produce relatively enriched pore waters,
potentially at the site of carbonate precipitation. In any case, the input of DIC sourced
from the methanogenic zone must have been significantly high to overcome the input of
12
C derived from AOM if this process were contributing to concretion formation in
Monterey sediments. It is difficult to distinguish whether or not AOM played a
significant role in the generation of MNC1 and Montana de Oro concretions, however it
is clear that the majority of DIC was sourced from environments experiencing
methanogenesis and that sulfide oxidation was actively occurring.
Ultimately, AOM generally occurs close to the sediment-water interface where
oxygen and/or nitrate are available to the sulfide oxidizing bacteria. Given the overall
thickness and size of MNC1 and the concretions of Montana de Oro, it is difficult to
imagine precipitation in the potentially thin zone (interface) of AOM, therefore derivation
of an alternate hypothesis is warranted.
Another possibility is sulfide oxidation by buried sedimentary iron oxides.
Sulfide oxidation coupled with the reduction of iron oxide minerals can act to increase

113
pore water alkalinity by a factor of 2 per mole of oxidant consumed (see Eq. 15 in Table
3.1) and promote carbonate precipitation. Coupling iron reduction (Eq. 15) with sulfur
disproportionation (Eq. 16) yields a net equation demonstrating the production of
dissolved sulfate and sulfide. Given that the source of sulfur in these reactions is sulfide
(likely isotopically depleted, see above), the produced sulfide and sulfate will likely have
a depleted isotopic value compared to seawater. Thus this mechanism not only increases
alkalinity but also supplies isotopically depleted sulfate to pore waters. In addition, the
coupled iron reduction-sulfur disproportionation reactions represent a positive feedback
situation that could foster the generation large concretions, given sufficient iron oxides.
This hypothesis does not require negative δ
13
C values nor is it as limited with respect to
depth as the AOM mechanism.

Evidence Supporting the Presence of Isotopically Depleted Sulfur: Pyrite δ
34
S values
of MNC1 and Montana de Oro concretions (Table 3.1 and Figure 3.12A) are isotopically-
depleted compared to the Miocene seawater, indicating the presence of low pore water
sulfide δ
34
S, a requirement of both the AOM and iron reduction-sulfide oxidation
mechanisms. However, concretion-hosted pyrite represents pyrite formed at all sediment
depths shallower than that of concretion precipitation. Pore waters of Namibian shelf
sediments exhibit sulfide δ
34
S values ca. 16‰ (similar to MNC1 and Montana de Oro
samples), extending well below the zone of diffusional sulfate (Dale et al., 2009). These
data indicate that isotopically depleted (compared to seawater) sulfide can extend below

114
Figure 3.12: Monterey sulfur isotope cross plots

δ
34
S
CAS
vs. A) δ
34
S
pyr
and B) CAS concentration. Circled data point in B is possibly
erroneous, as discussed in text. Dashed lines correspond to the approximate δ
34
S value of
Miocene seawater sulfate (~22‰).

115
the depth of diffusional sulfate in modern, organic-rich sediments and, given an oxidant,
could produce similarly depleted sulfate, albeit probably in small quantities.
Figures 3.5A and 3.5B demonstrate that CAS concentration and δ
34
S
CAS
do not
yield noticeable trends within individual concretions. This may indicate heterogeneities
in the nature and degree of sulfide oxidation in localized environments, perhaps on scales
of < ~1m laterally (compare to distances between samples Figure 3.5) in sediments of the
Monterey. Perhaps this lateral variation is the result of the distributions of sulfide
oxidizing communities or the relative abundances of oxidants within the sediment
column. Our limited understanding of natural sulfide oxidation in modern environments
prevents further interpretation, however the as of yet undiscovered, well-defined sulfide
oxidation “zone” (similar to the well-defined sulfate reduction zone) is consistent with
the heterogeneous data presented here.
The relatively high cement contents in MNC1 and Montana de Oro samples
suggest that methanogenesis and sulfide oxidation (and possibly AOM) operated at
relatively shallow sediment depths (however, see below). The high organic matter
concentration in the Monterey Formation likely led to the development of shallow redox
transitions. Thus, although one of the deeper organo-diagenetic environments, the zone
of methanogenesis could have occurred near the sediment-water interface during
Monterey deposition. Interestingly, despite having comparable organic matter
concentrations (~5 wt%: Zaback and Pratt, 1992), the sulfate reduction-associated
concretions of Shell Beach contain significantly less dolomite compared to concretions of
Montana de Oro and MNC1—a seemingly counter-intuitive observation. The relatively

116
high dolomite contents of Montana de Oro and MNC1 (compared to Shell Beach) can be
partially explained by the remineralization of originally siliceous microfossils (diatoms)
and phosphate nodules (Figure 3.8B and 3.8F). Remineralization of originally non-
dolomite phases to dolomite will yield minus cement porosities that are skewed toward
high values. Thus, dolomite wt% is not a precise measure of original porosity in some
circumstances. The wide spread in dolomite wt% of the concretions of Montana de Oro
(Figure 3.9) is likely in part the result of varying degrees of diatom remineralization.
Ultimately, alternate factors will determine the depth of each organo-diagenetic zone,
such that the different lithologic members of the Monterey may have experienced
different degradation horizons at different depths.

Possible Contamination of CAS Signal
It is important to consider the possibility of contamination during the CAS
extraction procedure. Of particular importance is the oxidation of pyrite during carbonate
acidification (Marenco et al., 2008). While pyrite oxidation is problematic, it should be
relatively easy to identify given predictable geochemical consequences. CAS data
influenced by pyrite oxidation should show all of the following: 1) negative correlation
between CAS concentration and δ
34
S
CAS
, 2) positive correlation between pyrite and CAS
concentration and 3) negative correlation between pyrite concentration and δ
34
S
CAS
.
Aside from those of Shell Beach, none of the samples exhibit these correlations (Figure
3.12 and Figure 3.13). The Shell Beach samples exhibit only one of them—a negative
correlation between pyrite concentration and δ
34
S
CAS
(Figure 3.13B). However, if we

117
Figure 3.13: Assessing pyrite oxidation in Monterey samples

Cross plots comparing pyrite concentration (wt%) to A) CAS concentration and B)
δ
34
S
CAS
. Sample contamination via pyrite oxidation should produce a positive correlation
in A and a negative correlation in B.

118
assume that the elevated δ
34
S
CAS
values of SB concretions were influenced by pyrite
oxidation, the true values would be even higher, still consistent with a sulfate reduction
zone origin.
Special care must be taken when interpreting ca. ≤22‰ δ
34
S
CAS
values such as
those in samples of Naples Beach and Montana de Oro due to possible contamination by
pyrite oxidation, even in the absence of predictable geochemical correlations (see above).
The overall low concentration of pyrite in these samples (below 0.2 wt%) makes
significant influence by pyrite oxidation unlikely. In fact, experimental-based modeling
by Marenco et al. (2008) shows that at these low pyrite concentrations and Δ
34
S
CAS-pyr

values, pyrite oxidation could only reduce the true δ
34
S
CAS
value by <0.5‰ and <0.1‰
for Montana de Oro and Naples Beach samples, respectively. Thus the δ
34
S
CAS
of these
samples closely record initial precipitation conditions and not methodological artifacts.

δ
18
O and Relative Precipitation Depth
As Figure 3.11 demonstrates, all of the concretions analyzed display strong
correlation between δ
13
C and δ
18
O. Samples from Shell Beach and Naples Beach exhibit
negative correlations and samples from Montana de Oro exhibit a positive correlation.
Previous studies have used concretion δ
13
C and δ
18
O variation to determine relative
precipitation depth and temperature (Mertz, 1984; Kastner et al., 1984; Garrison et al.,
1984; Kablanow et al., 1984; Henderson et al., 1984; Kushnir and Kastner, 1984).
However, δ
18
O values in carbonates are complicated by the inability to directly determine
both the oxygen isotopic composition and temperature of precipitate-yielding fluids

119
(Urey, 1947; McCrea, 1950; Epstein et al., 1953; Shackleton, 1974; Walls et al., 1979;
Allan and Matthews, 1982; Erez & Luz, 1983; Meyers and Lohmann, 1985; Land, 1986;
Grossman & Ku, 1986; Kim & O’Neil, 1997; Bemis et al., 1998). Thus, without
additional constraints this approach is somewhat speculative.
An example of the potential problems associated with the traditional approach is
evident in MNC2. MNC2 exhibits a negative correlation between δ
13
C and δ
18
O,
however multiple lines of evidence (decreasing δ
13
C and CAS, increasing δ
34
S
CAS

outward) support a concentric growth pattern in this concretion. Thus a more
parsimonious interpretation is that MNC2 grew concentrically in sediments characterized
by increasing fluid δ
18
O values possibly reflecting influence by silica/clay diagenesis
(Behl and Garrison, 1994), the dissolution of clathrates (Perry et al., 1976) and/or the
influx of high salinity brines (Morton and Land, 1987). A similar situation is recognized
MMC3.
The traditional δ
13
C and δ
18
O relative depth/temperature approach can only be
utilized if concretion growth geometry is known, thus strong correlations in Figure 3.11,
although intriguing, must only be considered in light of additional evidence.

CONCLUSIONS
Dolomite concretions and cemented layers occur in all four members of the
Miocene Monterey Formation. Traditional stable isotopic analyses of δ
13
C
carb
and
outcrop textural characteristics indicate that concretions form in shallow sediments as a
result of the remineralization of sedimentary organic matter. However, δ
13
C
carb
alone

120
cannot distinguish among the different microbial and abiotic processes operating in the
diagenetic environment. Here we utilize carbonate-associated sulfate (CAS) and its
isotopic composition (δ
34
S
CAS
) in addition to δ
13
C
carb
in order to distinguish among the
primary organo-diagenetic horizons exhibited in marine sediments.
A solitary concretion from the Phosphatic Shale Member of Naples Beach
exhibits δ
34
S
CAS
values similar to Miocene seawater sulfate (~22‰), relatively high CAS
concentrations and negative δ
13
C
carb
. These data are consistent with shallow formation
near the sediment-water interface in a zone experiencing organic matter oxidation by
possibly nitrate but most likely metal oxides and only minor sulfate reduction. Nearly all
samples of the Phosphatic Shale Member at Shell Beach display significant
concentrations of CAS, elevated δ
34
S
CAS
(up to ~37‰) and moderately negative δ
13
C
carb
,
consistent with precipitation in the sulfate reduction zone. Concretions of the Siliceous
Member at Montana de Oro and a second concretion from the Naples Beach locality
exhibit reduced to negligible CAS concentrations, low δ
34
S
CAS
and positive δ
13
C
carb
. The
carbon isotope data and near negligible CAS concentrations are consistent with formation
in a zone experiencing methanogenesis. Interestingly, these concretions yield δ
34
S
CAS

<22‰, suggesting influence by sulfide oxidation (possibly fostered by nitrate or iron
reduction), perhaps the first time the influence of sulfide oxidation has been noted during
concretion formation.

121
CHAPTER 4: CARBON AND SULFUR ISOTOPIC COMPOSITIONS OF
CALCITIC CONCRETIONS OF THE UPPER CRETACEOUS HOLZ SHALE:
CATCHING SULFIDE OXIDATION IN THE ACT

CHAPTER 4 ABSTRACT
The Holz Shale, an organic-rich, late Cretaceous marine slope deposit, hosts
disseminated nodular concretions and locally continuous cemented layers. These
concretionary structures are composed of fine-grained calcitic cement and contain a
relatively high concentration of insoluble material (~50wt%). Carbon isotope values are
consistently negative and decrease outward from center to rim in all nodular concretions.
These trends in δ
13
C indicate that carbon was sourced from the remineralization of
organic matter and a concentric growth habit with increased burial. Concentric growth is
further corroborated by a decrease in δ
18
O from center to rim in the largest concretion
analyzed in this study.
Carbonate-associated sulfate (CAS) concentrations are significantly reduced
compared to those expected for late Cretaceous, primary marine carbonates, which is
consistent with a diagenetic origin. Sulfur isotope compositions of CAS are negative
with values ranging from –7.7 to –26.2‰ VCDT. Geochemical relationships among
pyrite concentration, CAS concentration and δ
34
S
CAS
, in addition to a stepwise CAS
acidification test, suggest that the signal is not a methodological artifact. Negative
δ
34
S
CAS
values are consistent with pore waters strongly influenced by sulfide oxidation,
possibly coupled with the reduction of sedimentary iron oxides. Modeling based on the

122
measured sulfur isotope values suggests that more than 39% (likely much more) of the
pore water sulfate in the zone of concretion formation was derived from the oxidation of
isotopically depleted sulfide. Such a high contribution of sulfate from sulfide oxidation
implies sediment depths below or near the bottom of the zone containing seawater-
derived sulfate. Sulfate generated in this low sulfate zone would be immediately
consumed by sulfate-reducing bacteria, therefore simultaneous carbonate precipitation
must have occurred in order to explain the CAS concentrations reported here. Sulfide
oxidation by iron oxides, coupled with sulfur disproportionation, provides the necessary
alkalinity and sulfate to account for the generation and geochemistry of Holz Shale
concretions. The results of this study demonstrate that concretionary carbonate can
preserve signals that indicate the presence of sulfide oxidation in past marine sediments.

INTRODUCTION
Calcitic concretions and concretionary layers have been recognized in
sedimentary units of all ages and depositional environments (Mozley and Burns, 1993),
including the marine slope deposit of the Cretaceous Holz Shale (Buck and Bottjer; 1985)
(Figures 4.1 and 4.2). Concretion precipitation occurs within sediments and in some
cases the carbon is derived from the remineralization of sedimentary organic compounds,
as revealed by δ
13
C analyses (i.e., Claypool and Kaplan, 1974). However in modern
sediments a multitude of biotic and abiotic processes remineralize organic matter
(Jorgnesen, 1983) and have the potential to attribute to carbonate precipitation. These

123
Figure 4.1: Map showing location of the Holz Shale outcrop examined in this study

124
processes include, but are not limited to, methanogenesis, thermal decarboxylation and
organic matter oxidation by free oxygen, nitrate, metal oxides (iron and manganese
oxides, dominantly) and sulfate (Claypool and Kaplan, 1974; Froelich et al., 1979).
Aside from environments dominated by methanogenesis, degradation environment
cannot be distinguished based solely on δ
13
C analyses (see Claypool and Kaplan, 1974).
Carbonate-associated sulfate (CAS), pyrite concentrations and their respective sulfur
isotopic compositions can be employed in order to further characterize degradation
environment.
CAS and pyrite data can be used not only to identify processes directly associated
with the transformation of organic carbon into dissolved inorganic carbon (such as sulfate
reduction), but also has the potential to distinguish contemporaneous diagenetic reactions,
including sulfide oxidation. Although recognized in modern sediments, sulfide oxidation
is not typically considered “dominant”, particularly in comparison to sulfate reduction
(Jorgensen, 1983; Berner, 1984; Canfield and Thamdrup, 1994; Canfield, 2001).
However, our ability to identify the presence of sulfide oxidation in modern
environments may be limited because of sampling resolution and the transient nature of
the sulfide oxidation signal.
Here we couple traditional δ
13
C with the above sulfur analyses in order to
determine which major degradation environment(s) led to the formation of the calcitic
concretions of the Holz Shale. These analyses reveal that carbonate carbon is at least
partially sourced from organic matter remineralization and that sulfide oxidation was
active in the zone of precipitation.

125
Figure 4.2: Photographs of concretionary structures and host rock (A-D) and stratigraphic column (E) of the Holz Shale.

126
GEOLOGIC CONTEXT
The Holz Shale is organic-rich, fossiliferous and crops out in the Santa Ana
Mountains of southern California (Schoellhamer et al., 1954). The Cretaceous strata of
the Santa Ana Mountains is dominated by siliciclastic rocks and major lithologies include
shales, sandstones and conglomerates. The Holz Shale comprises the upper member of
the Ladd Canyon Formation which is interpreted as having been deposited sometime
between the late Turonian and early Campanian (ca. ~85 Ma) (Ogg et al., 2006), based on
a biostratigraphically representative fossil assemblage. The Holz Shale is dominantly
composed of rubbly mudstone and siltstone and contains disseminated white calcitic
concretions and concretionary layers (Figure 4.2).
The Holz Shale has been most recently interpreted as a marine slope deposit
(Buck and Bottjer, 1985), based largely on the presence of fossils of the thin-shelled
bivalve Inoceramus, an assemblage of agglutinated foraminifera and the trace fossils
Thalassinoides and Ophiomorpha. This paleoenvironment is supported by the common
occurrence of chute and gully deposits, conglomeratic debris flows and turbidites (Buck
and Bottjer, 1985). Alternate depositional environments have been suggested by Wheeler
(1952), Orr (1964), Lang (1978), Sundberg and Cooper (1978) and Sundberg (1980,
1982), however the most recent interpretation by Buck and Bottjer (1985) seems most
parsimonious with the available data.

127
METHODS
Five concretions and concretionary layers were sampled from the upper ~50 m of
the Holz Shale, just west of Silverado, California (Figure 4.1). Individual concretionary
bodies were sampled multiple times in order to explore spatial variability in geochemical
trends. Geochemical measurements included carbonate-associated sulfate (CAS)
concentration, pyrite concentration, δ
34
S
CAS
, δ
34
S
pyrite
, δ
13
C and δ
18
O. Each sample was
thin-sectioned and examined under a petrographic microscope in order to characterize the
nature of its carbonate cements and its insoluble components. Select samples were
examined under scanning electron microscope (SEM).
Refer to chapter 1 for methods concerning pyrite concentration and isotopic
analyses. Refer to chapters 1 and 3 concerning CAS extraction procedure. Duplicate
measurements of CAS concentration were performed on 10 samples and precision of
replicates was within +/–100 ppm.

RESULTS
Nature of Concretions of the Holz Shale
Concretionary structures of the Holz Shale consist of light gray, disseminated
nodules and locally continuous cemented layers (Figure 4.2). Nodular concretions range
in size from ~5cm to 1m across and the host rock typically exhibits laminae deflection
around the concretions. Cemented layers range in thickness from ~10 to 30cm and are
relatively resistant compared to the surrounding rubbly host rock. In thin section (see
Figure 4.3), all concretionary structures exhibit fine-grained cements with crystal

128
Figure 4.3: Photomicrographs and SEM images of Holz Shale concretionary structures

Photomicrographs (A-D) and SEM images (E and F) of Holz Shale concretionary
structures. A) Photomicrograph dominated by angular quartz grains (arrows), a major
insoluble component in Holz concretionary bodies. B) Diatom frustule within a Holz
Shale concretion. Notice fine-grained carbonate matrix. C) Fine-grained carbonate
intruded by a sparry cement-containing septarian vein. D) Photomicrograph of a cluster
of pyrite (nearly all of the opaque grains) within a concretion of the Holz Shale. E) SEM
image of pyrite framboids in a fine-grained carbonate matrix. F) SEM image of a
rectangular diatom frustule partially infilled by cubic pyrite.

129
diameters <<5µm. The fine-crystalline texture makes it difficult to distinguish grain
boundaries and to define crystal habit (Figure 4.3A-3D). Biogenic material includes
variably recrystallized microfossils, including foraminifera and diatom frustules (Figure
4.3B). Septarian fissures occur in some concretionary structures and are infilled with a
relatively coarse-crystalline, sparry cement (Figure 4.3C). Insoluble materials include
angular quartz grains (Figure 4.3A), rare micas, pyrite (Figure 4.3D-3F), unidentified
opaque grains, fine-grained clays and non-reminerallized diatom tests. Pyrite is readily
identifiable in most thin sections and can occur in high-density groupings (Figure 4.3D).
SEM imaging reveals that most of the pyrite exhibits a framboidal morphology and the
framboids range up to ~40µm in diameter (Figure 4.3E). Pyrite is also found infilling
microfossil tests (Figure 4.3F). Insoluble material comprises between 38 and 82 wt %
which, assuming that all other material is calcium carbonate, yields a range of carbonate
concentrations from 18 to 72 wt % (Table 4.1).

Geochemistry
Three nodular concretions and two cemented layers were sampled for this study:
nodules HC2, HC3 and HC4 and layers HCL and HCL2 (Figure 4.4 and Table 4.1).
Photographs of HC2, HC3, HC4 and HCL are shown in Figure 4.4 along with sketches
containing CAS concentration, δ
34
S
CAS
and δ
13
C data for reference.

Carbon and Oxygen Isotopic Composition: δ
13
C values for the concretionary structures
analyzed are consistently negative and range from –10.2 to –3.3‰ VPDB (Table 4.1).

130
Table 4.1: Geochemical data from Holz Shale concretionary structures.

131
δ
18
O values are also consistently negative and range from –9.2 to –5.4‰ VPDB. All
three nodular concretions exhibit carbon isotopic depletion from center to rim (Figure
4.4). In the case of HC3, the largest nodule, the carbon isotopic composition of the
peripheral samples is ~4‰ depleted compared to the inner samples (Figure 4.4). HC3
exhibits a significant positive correlation between δ
13
C and δ
18
O, whereas HC2 and HCL
do not (Figure 4.5). HC4 and HCL2 were not sampled sufficiently enough to determine
the presence or lack of isotopic correlation between δ
13
C and δ
18
O.

CAS and Pyrite Concentration: The total range in CAS concentration among all
samples analyzed is from 0 to 665ppm (Figure 4.6 and Table 4.1). The majority (all but
three) of samples contain less than 300ppm CAS. CAS concentrations do not show
consistent spatial variation in any of the concretionary bodies analyzed (Figure 4.4).
Pyrite concentrations range from essentially 0 to ~0.9 wt %, and most samples (all but
one) contain between 0 and 0.54 wt %.

Sulfur isotopic composition of CAS and pyrite: δ
34
S
CAS
values are consistently
negative and range from –26.2 to –7.7‰ (Figure 4.4 and Table 4.1). As with CAS
concentration, δ
34
S
CAS
values do not exhibit consistent spatial trends in individual
concretions or cemented layers (Figure 4.4). δ
34
S
pyrite
values display a tremendous range
from +32.6 to –39.3‰ VCDT, however only HCL exhibits positive values; all other
structures yield values ranging from –3.4 to –39.3‰ (Table 4.1).

132
Figure 4.4: Photographs and associated sketches showing sampling pattern. Refer to
Table 4.1 for geochemical data. Notice the center to rim depletion in δ
13
C in concretions
HC2, HC3 and HC4 and the lack of noticeable trends in CAS and δ
34
S
CAS
. The outward
depletion of δ
13
C in HC2 is exhibited in all three dimensions. The front portion of HC2
is missing, likely due to subsequent erosion. Therefore, the central regions of panels A-B
and C-D represent the true core of the concretion. The back panel E-F and the outer
samples of A-B and C-D represent the peripheral samples of the nodule.

133
Figure 4.4 continued

134
DISCUSSION
Growth Habit and Carbon Source of Nodular Concretions
The nodules examined in this study (HC2, HC3 and HC4) all exhibit an outward
depletion in carbon isotope values such that inner or core samples yield relatively heavy
δ
13
C compared to peripheral samples (Figure 4.4). Notice that HC2, the only concretion
adequately sampled in all three dimensions, yields an outward depletion in all three (see
caption of Figure 4.4). HCL samples (1A and 1B) may appear nodular in Figure 4.4,
however these structures are portions of a continuous cemented layer. A center to rim
decrease in δ
13
C is consistent with concentric concretion growth (Raiswell and Fisher,
2000) with a progressive increase in the amount of carbon sourced from the
remineralization of organic matter or the oxidation of methane, both of which are
common in marine sediments. Regardless, the ultimate source of carbon is most likely
sedimentary organic matter, as methane is derived from the degradation of organic
compounds through biotic (Bernard et al., 1977) or abiotic processes (Tissot and Welte,
1984).
HC3, the largest nodular concretion, exhibits a contemporaneous outward
decrease in δ
18
O (in agreement with the correlation shown in Figure 4.5), whereas HC2
and HC4 do not. These characteristics suggest that HC3 grew in sediments experiencing
increasing depth and temperature in contrast to HC2 and HC4, which grew at constant
depths and temperatures. The relative sizes of these concretions is consistent with a
longer growth time for HC3 compared to those for the two smaller nodules. Decreasing

135
Figure 4.5: Cross plot of Holz Shale δ
13
C and δ
18
O

Notice strong positive correlation in HC3.

136
δ
18
O on the order of 2.7‰ (the difference between the heaviest and lightest δ
18
O values
of HC3, see Table 4.1) is consistent with an increase in temperature of ~18°C, based on
the most recent δ
18
O
calcite
/temperature relationship of Kim and O’Neil (1997). However,
this calculated temperature increase is only valid if the initial pore water δ
18
O value was
constant in the region of concretion precipitation. If correct, the temperature increase
suggests progressive concretion growth over a depth range of ~720m, given an average
geothermal gradient of 25°C/km. It is difficult to account for a sufficient supply of Ca
2+

by diffusion from seawater over this depth range. This discrepancy may arise because of
1) varying pore water δ
18
O with concretion growth, 2) a geothermal gradient >25°C/km
and/or 3) a source of calcium other than diffusion from seawater.
Overall, the δ
13
C data are consistent with concentric growth in HC2, HC3 and
HC4. In addition, the larger size of HC3 and variation in δ
18
O is consistent with a longer
growth time compared to HC2 and HC4.

Possible Contamination of the CAS Signal
Prior to the interpretation of the sulfur isotope and CAS data, it is important to
consider the possibility of contamination during the CAS extraction procedure, especially
given the extremely light δ
34
S
CAS
values exhibited by the concretions and cemented
layers of the Holz Shale. Of particular importance is the oxidation of pyrite during
carbonate acidification (Marenco et al., 2008). While pyrite oxidation is problematic, it
should be relatively easy to identify given predictable geochemical consequences. CAS
data influenced by pyrite oxidation should show all of the following: 1) positive

137
Figure 4.6: Assessing pyrite contamination in the Holz Shale

Cross plots of pyrite concentration versus A) CAS concentration and B) δ
34
S
CAS
. Neither
show a strong correlation, suggesting the absence of pyrite oxidation during the CAS
extraction procedure.

138
correlation between pyrite and CAS concentration, 2) negative correlation between pyrite
concentration and δ
34
S
CAS
and 3) consistently positive Δ
34
S (Δ
34
S = δ
34
S
CAS
– δ
34
S
pyrite
)
values. Figure 4.6 demonstrates that neither 1 nor 2 above are exhibited by concretionary
structures of the Holz Shale. In addition, while most Δ
34
S values are positive, a few are
negative (Table 4.1) and it is impossible to achieve δ
34
S
CAS
values lower than δ
34
Spyrite by
pyrite oxidation. If an isotopic fractionation accompanied the oxidation, negative Δ
34
S
values could potentially be generated, however as of yet no such fractionation has been
identified.
A two-step acidification test was conducted in order to further test contamination
by pyrite oxidation. Longer acidification should lead to more extensive pyrite oxidation
and thus an increase in CAS concentration (apparent CAS concentration as denoted by
Marenco et al., 2008). Five samples were acidified following the method for CAS
extraction outlined above. Replicate samples were exposed to the same HCl solution
over a 24-hour duration. As Figure 4.7 demonstrates, CAS concentrations did not
increase as a result of longer acidification.
Given the above geochemical characteristics it is unlikely that the samples
analyzed here experienced contamination by the oxidation of pyrite during CAS
extraction.

Characterization of Concretion Growth Environment
CAS and Pore Water Sulfate: In modern marine sedimentary environments, sulfate is
primarily supplied by diffusion from the overlying water column (Claypool and Kaplan,

139
Figure 4.7: CAS acidification test for five Holz Shale samples

Notice how increased acidification times do not produce an increase in CAS
concentration (within the ±100ppm error of the analysis). This is further support that
contamination by pyrite oxidation during CAS extraction was not significant.

140
1974). As the sulfate diffuses downward, it is progressively consumed by bacterial
sulfate reduction in organic matter-containing sediments (Jorgensen, 1983; Berner, 1984;
Canfield and Thamdrup, 1994; Canfield, 2001). This produces sedimentary pore waters
that are reduced in sulfate compared to the overlying water column. Therefore,
carbonates precipitated in marine sediments should contain less sulfate (CAS) that those
precipitated in the open ocean—a similar situation likely existed during the late
Cretaceous. Lowenstein et al. (2003) have shown that marine sulfate concentrations were
significantly lower during the late Cretaceous and suggest a marine sulfate concentration
of ~18mM (compared to the modern concentration of 28mM sulfate). Modern CAS
concentrations of open marine-precipitated carbonates average ~2400ppm (Lyons et al.,
2004; Gill et al., 2008). Assuming that ~2400ppm CAS is incorporated into carbonates
precipitating from a 28mM sulfate solution, it is reasonable to assume that ~1500ppm
CAS is incorporated into carbonates precipitating from an 18mM solution (the reported
sulfate concentration of late Cretaceous seawater). CAS concentrations of the
concretionary structures analyzed from the Holz Shale are significantly less than
1500ppm (Figure 4.8), consistent with precipitation in the sediment column.

δ
34
S and the Source of Sulfate: As Table 4.1 and Figure 4.9 demonstrate, δ
34
S
CAS

values are consistently negative and as low as –26.2‰. In modern marine sediments,
sulfur systematics are dominated by bacterial sulfate reduction (BSR), which acts to
remove pore water sulfate (Jorgensen, 1983; Berner, 1984; Canfield and Thamdrup,
1994; Canfield, 2001). The preferential reduction of
32
SO
4
2-
by microbes acts to increase

141
Figure 4.8: CAS concentrations of concretionary structures of the Holz Shale

All samples exhibit CAS concentrations that are significantly reduced compared to that
expected for late Cretaceous, marine-precipitated carbonate minerals.

142
the δ
34
S of the residual pore water sulfate with depth (Figure 4.10). Enrichment of
34
S
sulfate
requires a mechanism by which
32
S is removed and sequestered from the
dissolved sulfate pool. In general, pyrite represents the dominant sink for dissolved S in
marine sediments, such that in a typical sedimentary system the isotopic mass balance
must follow the relationship,
δ
34
S
sw
= f
pws
(δ
34
S
pws
) + f
ds
(δ
34
S
ds
) + f
bp
(δ
34
S
bp
) (Eq. 1)
where δ
34
S
sw
, δ
34
S
pws
, δ
34
S
ds
and

δ
34
S
bp
represent the sulfur isotopic composition of
seawater sulfate, pore water sulfate, pore water dissolved sulfide and buried pyrite,
respectively. The fractions of sulfur present as pore water sulfate, sulfide and buried as
pyrite are represented by f
pws
, f
ds
and f
bp
, respectively. This relationship dictates that all
sedimentary sulfur be present as pore water sulfate, sulfide and buried pyrite, all sulfur be
supplied to the system by diffusion from seawater and f
pws
+ f
ds
+ f
bp
= f
diffusional
= 1. As
stated above, δ
34
S
pws
values are typically enriched compared contemporaneous seawater,
however δ
34
S
CAS
values in samples from the Holz Shale are significantly depleted
compared to the late Cretaceous δ
34
S
sw
value of ~18‰ (Paytan et al., 2004a), suggesting
an alternate source of sulfate.
The oxidation of dissolved sulfides could represent a source of isotopically
depleted sulfate. Isotopically depleted sulfides are generated during BSR as shown in the
following reaction,
C
org
+ SO
4
2–
+ 2H
2
O  2HCO
3
–
+ H
2
S (Eq. 2)
where C
org
is simplified organic carbon. The product H
2
S can then undergo additional
reactions in order to form pyrite (Wilkin and Barnes, 1997). Alternatively, the H
2
S can

143
Figure 4.9: Cross-plot of Holz Shale δ
13
C and δ
34
S
CAS

All data points exhibit moderately negative δ
13
C and extremely negative δ
34
S
CAS
.

144
be reoxidized and reintroduced into the dissolved sulfate pool. The oxidation of
dissolved sulfide to sulfate is not accompanied by a significant isotopic fractionation (Fry
et al., 1988), therefore the produced “second generation” sulfate should retain an
isotopically depleted signature and lower pore water sulfate isotope values (Figure 4.10).
Sulfide oxidation could therefore account for the isotopically light CAS signature of the
concretionary structures of the Holz Shale. It is important to note that the generation of
isotopically light sulfate can only be achieved by this mechanism if sulfide oxidation is
incomplete since Eq. 1 above must still hold. In other words, all of the sulfur entering the
system through diffusion cannot equal the amount of sulfide oxidized—such a scenario
would yield pore water sulfate with an isotopic value identical to the input value (~18‰
in the late Cretaceous).

Quantifying the Amount of Sulfate Sourced from Sulfide Oxidation
The recurrence of isotopic enrichment with depth in modern marine sediments
suggests that sulfide oxidation is a relatively minor process. A simple box model
constructed using the sulfur isotopic values of the Holz Shale concretions demonstrates
the role of sulfide oxidation as it relates to the production of pore water sulfate (Figure
4.11A). This model assumes that pore water sulfate is a mixture of diffusional,
isotopically enriched sulfate derived from seawater and isotopically depleted sulfate
derived from the oxidation of dissolved sulfide. The mass balance follows,
δ
34
S
pws
= f
D
(δ
34
S
D
) + f
SOX
(δ
34
S
SOX
) (Eq. 3)

145
Figure 4.10: Diagram depicting the relative concentration of dissolved pore water sulfate
and sulfide and the evolution of δ
34
S
pws
with increasing depth

If influenced strictly by bacterial sulfate reduction accompanied by pyrite burial, δ
34
S
pws

will evolve along the curve labeled BSR. Sulfide oxidation will act to drive δ
34
S
pws

lower, as shown by the arrows. The gray envelope represents the range of sulfide δ
34
S
generated by sulfate reduction and/or sulfur disproportionation given the value of initial
δ
34
S
pws
value at that depth. The envelope extends to 70‰ depleted compared to the
initial δ
34
S
pws
curve and represents the maximum isotopic fractionation generated during
sulfur disproportionation (Canfield and Teske, 1996). The dashed line represents an
isotopic fractionation of 46‰, the maximum offset generated by BSR (Canfield and
Teske, 1996). The red region highlights the range in δ
34
S
CAS
values of the Holz Shale
concretionary carbonates. The stippled region represents a reasonable range of sulfide
δ
34
S based on measured H
2
S δ
34
S values. These values are plotted based on sulfate-H
2
S
δ
34
S pairs compared to the BSR curve. The yellow box highlights the proposed depth of
Holz concretion precipitation.

146
where δ
34
S
pws
, δ
34
S
D
and δ
34
S
SOX
represent the isotopic compositions of pore water
sulfate, diffusional sulfate and sulfide oxidation-produced sulfate, respectively. The
fraction of sulfate derived from diffusing seawater and sulfide oxidation sources are
denoted by f
D
and f
SOX
, respectively.
The value of δ
34
S
pws
is set at the uppermost (–7.7‰, Figure 4.11B) and lowermost
(–26.2‰, Figure 4.11C) values of δ
34
S
CAS
exhibited by the concretionary structures of the
Holz Shale in order to determine the contribution of sulfate from sulfide oxidation. The
value of δ
34
S
D
is varied in 10‰ intervals, from +20 to +80‰. The value of δ
34
S
SOX
is
varied from –50‰ to the δ
34
S
CAS
value and spans the range of values within the gray
envelope of Figure 4.10.
The results of this box model are depicted in Figures 4.11B and 4.11C. The
model indicates that at least 39.5% of the sulfate in the Holz Shale pore waters at
concretion precipitation sites was derived from sulfide oxidation. The extremely negative
δ
34
S
CAS
value of HC3-IRA requires pore water receiving >60% of its sulfate from sulfide
oxidation. In both scenarios, the maximum amount of sulfate that could be derived from
the oxidation of sulfide is 100%.
This study demonstrates that even at the lowermost limits, the concretionary
structures of the Holz Shale received a significant amount of sulfate from sulfide
oxidation. These findings raise questions concerning the importance of sulfide oxidation
in not only past, but modern marine sedimentary environments.

147
Degree of Sulfide Oxidation and Environment
As stated above, the mass balance in Eq. 1 must hold true in all marine
sedimentary systems. It requires that any S with δ
34
S below that of the input value must
be balanced by S with a δ
34
S above the input value. This could be taken as being at odds
with the data presented here, which suggest that both reduced and oxidized S phases
exhibited δ
34
S values below that of the input. In order to maintain mass balance, an
additional sulfur pool must have been present with δ
34
S above the input value of ~18‰.
This
34
S-enriched pool must have been removed or isolated from the region of concretion
precipitation. Isolation between dissolved sulfate and sulfide is achieved in modern
sediments by diffusion coupled with consumption of sulfate by BSR (see concentration
profiles in Figure 4.10). Typically, dissolved sulfide reaches maximum concentrations
near or below the horizon at which dissolved sulfate is exhausted (Jorgensen, 1983)
(Figure 4.10). This scenario yields a relatively deep sedimentary environment that is
dominated by dissolved sulfide that may exhibit δ
34
S values below that of the input. The
more shallow sediments are dominated by dissolved sulfate with δ
34
S values above that
of the input. As a whole, the sediments maintain isotopic mass balance as long as the
shallow sulfate pool is large and isotopically heavy enough to account for both the deeper
dissolved sulfide pool and buried pyrite sulfur. Therefore the concretionary structures of
the Holz Shale likely formed at depths near or below that of residual seawater-sourced
sulfate in a zone dominated by dissolved sulfide (see Figure 4.10).

148
Figure 4.11: Holz Shale box model parameters and results

A) Diagram defining the inputs to the pore water system and the associated isotopic
values used in the model. B) Model run using the highest δ
34
S
pws
input value of –7.7‰.
C) Model run using the lowest δ
34
S
pws
input value of –26.2‰. Notice how at least 39.5%
of sulfate must have been sourced from sulfide oxidation to account for the δ
34
S
CAS

values of Holz Shale concretionary carbonates.

149
Carbonate precipitation must have been contemporaneous with sulfide oxidation
in order to account for the CAS concentrations in the Holz Shale concretions.
Contemporaneous precipitation is required because any sulfate produced at this depth
would be immediately consumed by sulfate-reducing bacteria. Therefore, a mechanism
of sulfide oxidation that can be directly coupled with increased alkalinity is most
reasonable.

Mechanisms of Sulfide Oxidation
Under certain conditions, some sulfide-oxidizing bacteria use nitrate (Prokopenko
et al., 2006) as an electron acceptor. This bacterially mediated reaction can act to further
increase pore water alkalinity (through removal of H
+
;

Kuvilia and Murray, 1984),
promoting the precipitation of concretionary carbonate. However, it is difficult to
account for nitrate penetration down to the depths calculated above for HC3, particularly
given the tendency for nitrate to react with organic carbon and be consumed at relatively
shallow sediment depths (Froelich et al., 1979). In addition, the presence of abundant
framboidal pyrite and the dominance of low δ
34
S
pyrite
values suggest that significant BSR
occurred prior to cementation (Wilkin and Barnes, 1997). In modern marine sediments,
nitrate is typically exhausted prior to the onset of sulfate reduction (Froelich et al., 1979),
adding additional complications.
Alternatively, sedimentary iron oxides could be responsible for sulfide oxidation.
Sulfide oxidation coupled with the reduction of iron oxide minerals can act to increase
pore water alkalinity, generate ΔA/ΔTCO2 values greater than 1 and promote carbonate

150
precipitation (see reaction 15 in Table 3.2). Coupling iron reduction (Kuvilia and
Murray, 1984) with sulfur disproportionation (Canfield and Teske, 1996) yields the net
production of dissolved sulfate and sulfide. Given that the source of sulfur in these
reactions is sulfide (likely isotopically depleted, see above), the produced sulfide and
sulfate will likely have depleted isotopic values compared to seawater. Thus this
mechanism not only increases alkalinity but also supplies isotopically depleted sulfate to
pore waters. In addition, the coupled iron reduction-sulfur disproportionation reactions
represent a positive feedback situation that could foster the generation of large
concretions such as HC3, given sufficient iron oxides. The buried iron oxide mechanism
is not as limited with respect to depth as the nitrate reduction mechanism above and is
perhaps more reasonable.

Evidence for Sulfur Disproportionation
Given an initial seawater δ
34
S value of ~18‰ and the tendency for BSR to
increase δ
34
S
pws
with depth, it seems necessary to invoke sulfur disproportionation to
produce the low δ
34
S
pyrite
values in some of the Holz Shale concretionary structures (see
reaction 16 in Table 3.2). BSR is thought to produce a maximum fractionation of ~46‰,
whereas sulfur disproportionation can create fractionations up to ~70‰ (Canfield and
Teske, 1996, but see also Canfield et al., 2010). δ
34
S
pyrite
values below –28‰ (implying a
fractionation of 46‰ or greater) likely indicate bacterial sulfur disproportionation (see
Table 4.1). δ
34
S
pyrite
values above –32‰ do not necessarily imply a lack of
disproportionation as initial pore water sulfate δ
34
S could have been >18‰. Therefore, it

151
is likely that bacterial sulfur disproportionation was active in the sediments of the Holz
Shale.

CONCLUSIONS
Calcitic nodular concretions and cemented layers of the Holz Shale exhibit
multiple geochemical trends consistent with a diagenetic origin. These include negative
δ
13
C values that decrease from center to rim in nodular structures. In addition, the largest
nodule exhibits a decrease in δ
18
O from center to rim, consistent with progressive growth
with increasing temperature and depth in the sediment column. CAS concentrations are
significantly reduced compared to those expected for late Cretaceous, marine-precipitated
carbonate, corroborating a diagenetic origin.
δ
34
S
CAS
values are consistently negative and do not show geochemical
relationships consistent with contamination by pyrite oxidation during the CAS extraction
procedure. These δ
34
S
CAS
values suggest that sulfide oxidation was a significant source
of dissolved sulfate in pore waters of the Holz Shale. Simple modeling indicates that
more than 39% and likely much more of the sulfate was derived from sulfide oxidation.
Such a high contribution of sulfate from sulfide oxidation suggests sediments dominated
by sulfide, at depths below or near the bottom of the zone containing significant
seawater-sourced sulfate. Sulfate at these depths is quickly consumed by sulfate-
reducing bacteria, therefore in order to explain the CAS concentrations measured here,
sulfide oxidation must have occurred contemporaneously with carbonate precipitation.
Oxidation most likely occurred by the reduction of buried, sedimentary iron oxides,

152
which, along with sulfur disproportionation, provided the necessary alkalinity and sulfate
to account for the geochemistry of the Holz Shale concretionary structures. This study
demonstrates that concretionary carbonate can “trap” an otherwise transient isotopic
signal of sulfide oxidation in marine sediments.

153
CHAPTER 5: THE ORIGIN OF THE MILLIMETER-SCALE LAMINATION IN
THE NEOPROTEROZOIC LOWER BECK SPRING DOLOMITE:
IMPLICATIONS FOR WIDESPREAD, FINE-SCALE, LAYER-PARALLEL
DIAGENESIS IN PRECAMBRIAN CARBONATES

CHAPTER 5 ABSTRACT
The Neoproterozoic lower Beck Spring Dolomite exhibits a prominent,
millimeter-scale, light/dark, planar lamination commonly interpreted as a primary
depositional fabric. Microscopic examination reveals that the dark laminae are composed
of micrite and contain a micrometer-scale, wavy fabric, whereas the light laminae are
composed of relatively coarse-crystalline pseudospar that engulfs, surrounds, and
crosscuts the dark laminae. Nearly all of the light laminae surround, or occur in
proximity to, elongate (long axis fabric-parallel), spar-filled cavity structures. The
average δ
18
O values for dark and light laminae are –1.8‰ (std. dev. 1.6) and –3.0‰ (std.
dev. 1.3) VPDB, respectively. Cavity-filling cements are even more depleted in
18
O, with
an average isotopic value of –6.2‰ (std. dev. 1.1). These characteristics suggest that the
coarse-crystalline, light laminae are the result of the preferential, aggrading neomorphism
of a thinly laminated, micritic host rock and that the (now cemented) cavity structures
acted as conduits for the fluid responsible for recrystallization. The oxygen isotope
distributions are consistent with progressive recrystallization by an
18
O-depleted fluid,
likely of meteoric origin. Therefore, we interpret the prominent millimeter-scale,
light/dark layering in the lower Beck Spring Dolomite to result from widespread,

154
laminae-parallel, diagenetic processes—perhaps controlled by fine-scale layer-parallel
differences in initial porosity and permeability—and not from primary sedimentary
processes.
Extensive, layer-parallel diagenesis on such a fine scale is rare in Phanerozoic
shallow marine deposits, due to disruption by bioturbation. However, before the advent
of metazoan burrowing, fine-scale, layer-parallel fluid flow would have been more
common. Thus, the diagenetic fabric of the lower Beck Spring Dolomite may represent a
mode of carbonate diagenesis restricted to the Precambrian. Similar diagenesis may also
occur in Phanerozoic environments devoid of vertical bioturbation.

INTRODUCTION
Extensive planar lamination in marine environments can be produced by a
multitude of primary sedimentologic (i.e., current-driven deposition and settling),
biologic (i.e., microbial trapping and binding or precipitation, so called “cryptalgal
lamination”), and geochemical processes (i.e., abiotic precipitation such as hardgrounds).
As a result, carbonate rocks exhibiting planar fabrics in outcrop are commonly assumed
to result from such primary processes. The lower member (informal) of the Beck Spring
Dolomite (BSD) is a Neoproterozoic carbonate platform deposit that crops out in multiple
mountain ranges in the Death Valley region, including the Alexander Hills (Licari, 1971,
1978; Marian, 1979; Marian and Osborne, 1980; Tucker, 1982, 1983) and is
characterized by an extensive millimeter-scale, planar lamination. Previous workers
suggested that the lamination was primary in origin, and specifically “cryptalgal” (e.g.,

155
Gustadt, 1968; Licari, 1971; Tucker, 1983). Here we critically examine this
interpretation using petrographic and isotopic (δ
13
C and δ
18
O) techniques in order to
address the possibility of a diagenetic origin for the lamination.
There have been previous accounts of planar fabric and bedform generation by
diagenetic processes in Phanerozoic carbonate sequences. These studies include the
development and enhancement of shale-limestone rhythmites (Hallam, 1964; Arthur,
1976; Arthur and Fischer, 1977), burrow-mottled and shaly-chalk interbeds (Hattin,
1971), chalk-hosted stylolitization (Garrison and Kennedy, 1977; Alvarez et al., 1985;
Andrews and Railsback, 1997), isolated concretionary layers (Irwin, 1980), marl-
limestone interbeds (Ricken and Hemleben, 1982; Arthur et al., 1984), and stylolite-free,
limestone-hosted pressure solution horizons (Bathurst, 1987). However, if diagenetic in
origin, the lamination exhibited by the lower BSD is as of yet undiscovered and
unreported in that it (1) occurs in a Precambrian dolomite (the above examples are
Phanerozoic in age), (2) is of millimeter scale and both laterally and stratigraphically
extensive, occurring at multiple outcrops spanning at least 40 km laterally (and perhaps
much farther), (3) is not derived from pressure solution, and (4) appears convincingly
primary when examined at outcrop scale.

AGE CONSTRAINT AND STRATIGRAPHY
The BSD is the middle member of the late Mesoproterozoic to Neoproterozoic
Pahrump Group, which crops out in eastern California and western Nevada (Figures 5.1
and 5.2). The Pahrump Group is preserved in rift-related, downdropped blocks (Stewart,

156
Figure 5.1: Map of the Alexander Hills and other outcrops of Proterozoic to early
Cambrian miogeoclinal units

Shaded regions delineate mountain ranges containing Proterozoic to Cambrian units. The
star marks the location of outcrops of the Beck Spring Dolomite that were sampled for
this study. Modified from Stewart (1970).

157
1970; Heaman and Grotzinger, 1992) and was likely more regionally extensive before
rifting. Deposition of the BSD postdates 1.087 Ga, the age of diabase sills that intrude
the underlying Crystal Spring Formation (Heaman and Grotzinger, 1992). An upper-
bounding age for the BSD is more difficult to constrain because no radiometrically dated
units occur between it and the Precambrian-Cambrian boundary (~ 542 Ma). The
presence of vase-shaped microfossils (VSM) in the upper BSD have been correlated to
the Grand Canyon Chuar Group, where similar microfossils occur beneath an ash dated
742 ± 6 Ma (Porter and Knoll, 2000; Karlstrom et al., 2000; Dehler et al., 2001).
However, the temporal distribution of VSM is poorly understood, thus their use as global
correlation tools remains somewhat speculative. The BSD lies below the Kingston Peak
Formation, noted for its glacial diamictites and correlated by others to have been
deposited between ~ 720 and 635 Ma (e.g., Prave, 1999; Hoffmann et al., 2004; Condon
et al., 2005). The age of the Kingston Peak glacial units is supported via correlation to
the Neoproterozoic succession in Idaho and Utah, where several radiometric dates are
available (e.g., Christie-Blick and Levy, 1989; Fanning and Link, 2004; Corsetti et al.,
2007). Thus, precise age constraints are nonexistent (age range between 1.08 Ga and 542
Ma), but the BSD was likely deposited ca. 750 Ma.
The BSD is composed of three primary (informal) units (Gustadt, 1968; Marian
and Osborne, 1980; Tucker, 1983): a lower laminated member, a middle oolitic-
packstone member, and an upper stromatolitic member (Figure 5.2). The lower member
consists of ~ 200 meters of alternating light and dark, grayish-blue laminated dolomite
with subordinate breccias, rollup structures and soft-sediment deformation features

158
Figure 5.2: Stratigraphic column of the Pahrump Group and overlying units of the Death
Valley succession

Stratigraphic column after Corsetti and Kaufman (2003) with inset of the Beck Spring
Dolomite. Dated horizons are discussed in detail in text. Scale corresponds to left-hand
stratigraphic column.

159
(Corsetti and Kaufman, 2003). The laminae are of millimeter scale and appear
continuous in outcrop (Figure 5.3A). The rare occurrence of calcified algal filaments
within some lower BSD laminae led Tucker (1983) to conclude that the texture is likely
the result of sediment trapping, binding, and precipitation of carbonate minerals by “blue-
green algae” (“cryptalgal laminite”). Similarity in laminae thickness and crystal size to
those of the upper stromatolitic member also led Licari (1971) to infer influence by algae
in the lower BSD. Many other authors have also interpreted the planar lamination as the
result of microbial influences (Gustadt, 1968; Cloud and Semikhatov, 1969; Licari, 1978;
Marian, 1979). We present new information that questions the primary nature of the
millimeter-scale lamination in the lower BSD, and suggest that the diagenetic processes
responsible for the light/dark lamination may be more widespread in Precambrian
carbonates than previously thought. In fact this characteristic may further distinguish
Precambrian carbonates from their Phanerozoic counterparts.

METHODS
Samples of the lower BSD were collected near the Western Talc Mine in the
Alexander Hills of San Bernardino County, California (Figure 5.1). Representative hand
samples were thin-sectioned and microdrilled for carbon and oxygen isotopic analysis
(Figure 5.4 indicates the region of sample drilled for isotopic analysis). Fifty-seven total
drilled samples of light laminae, dark laminae, and cavity cements were analyzed. Refer
to chapter 1 for isotopic methods.

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Figure 5.3: The lamination of the lower Beck Spring Dolomite viewed at increasing magnification

A) Outcrop photo showing apparent continuity of alternating light/dark lamination. B) Hand sample showing a more patchy and
discontinuous lamination with individual laminae exhibiting varying thicknesses. C) Photomicrograph of millimeter-scale lamination.
Dark laminae appear as relatively fine-grained micrite while light laminae are composed of more coarse-crystalline pseudospar.
Arrows highlight a micrometer-scale lamination that occurs solely within the dark, micritic laminae.

161
OBSERVATIONS
The millimeter-scale lamination of the lower BSD is extraordinarily striking in
both outcrop and hand sample and exhibits increasing complexity at higher magnification
(Figure 5.3). Many characteristics that provide information about the origin of the
laminae can be seen only under the microscope. Isotopic analyses reveal that discrete
phases (identified petrographically) retain unique isotopic signatures with respect to δ
18
O,
similar to other textural phases as explored by Tucker (1982, 1983).

Petrology
At outcrop scale, the planar lamination of the lower BSD appears continuous,
with individual lamina apparently extending many meters laterally (Figure 5.3A). The
laminae consist of alternating light-gray and dark-gray layers of dolomite, approximately
1 mm thick. In hand sample (Figure 5.3B) the lamination of the lower BSD ranges from
highly continuous, to discontinuous, to hardly recognizable (Figure 5.4). Hand-sample
examination also reveals more variable laminae thickness than is apparent at outcrop
scale, with sizes ranging from ~ 1 to 4 mm.

Petrography
The millimeter-scale lamination exhibited by the lower BSD is composed of three
major dolomitic components: dark laminae, light laminae (laminae shown in Figures
5.3C, 5.5 and 5.6), and cavity infill (Figures 5.6 and 5.7). Of these major components

162
Figure 5.4: Cut and polished slabs of the lower Beck Spring Dolomite

Labels indicate regions where samples were microdrilled for carbon and oxygen isotope
analysis. Sample BS8 was drilled only for cavity infill and is not shown. D = dark
laminae, L = light laminae, C = Cavity infill.

163
there are two types of light laminae, one type of dark laminae, and two types of cavity
infill.

Dark Laminae: The dark laminae of the lower BSD are composed of micrite to
microspar (crystallographic size definitions follow Bathurst, 1975) with grains/crystals up
to ~ 20 µm in diameter. Grain and crystal boundaries in the dark laminae are difficult to
delineate. The majority of the dark laminae studied were simply micrite and/or microspar
(Figure 5.5B, D). However, a 10-30-µm-thick, wavy lamination subparallel to the
millimeter-scale lamination can be distinguished within some dark laminae (Figures
5.3C, 5.5A, C). The black, wavy lamination is highly discontinuous in some areas and
appears as loosely connected, stringer-like structures. The micrometer-scale wavy
lamination never occurs in either the light laminae or cavity infill.

Light Laminae: Two distinct types of light laminae are recognized. The most common
type consists of a mosaic of relatively coarse-crystalline (~ 20-100 µm), anhedral
microspar to pseudospar (Figure 5.5). In thin section, the light laminae can appear highly
discontinuous, exhibit a patchy texture (Figure 5.6), and completely surround small areas
of finer-grained, dark micrite (Figure 5.7). In some regions, this type of light laminae
crosscuts and grades into dark laminae (Figures 5.6, 5.7). The second type of light
laminae is much more rare, occurring in only one of the hand samples examined (sample
BS7, discussed below). This type consists of light-gray micrite and crystal silt with
minor brecciated portions (Figure 5.8).

164
Figure 5.5: Morphologic variations in light and dark laminae of the Beck Spring
Dolomite.

A) Region of extensive neomorphism surrounding a dark lamina exhibiting well
pronounced, micrometer-scale, wavy lamination (arrow). B) Region of clear and
extensive lamination. C) Photomicrograph highlighting discontinuity of light laminae
(top and left), also shown are micrite-hosted, micrometer-scale laminae (arrow). D)
Region of poorly defined lamination.

165
Cavity Structures: Cavity structures commonly exhibiting a consistent cement sequence
are abundant in thin section (Figures 5.6, 5.7) and are readily recognizable in outcrop.
Some of the cavity structures are conspicuous (tens of centimeters in diameter) and
display flat bottoms and irregular tops (“stromatactis”-like structures), while others are
less conspicuous and highly elongate (length-to-width aspect ratios up to 10:1) with long
axes parallel to the orientation of the millimeter-scale, light/dark fabric (Figure 5.7). The
cavities, whether large or elongate, were first lined with fibrous and/or bladed,
isopachous cements (these appear dark in thin section) followed by coarse, anhedral spar
(up to more than 500 µm across) in the cavity center (these appear light in thin section
and hand sample), as shown by Tucker (1983). The contact between the host rock and
the cavity structures is sharp whereas the contact between light and dark laminae is more
diffuse. The cavity structures can be quite numerous and comprise up to ~ 10% of the
total thin-section area. The equant, anhedral, coarse-crystalline, cavity-filling spar gives
them a light-colored appearance, which, when elongate, could be mistaken for light
laminae in hand sample. However, close examination of such infill cements with a hand
lens at outcrop allows them to be distinguished from light laminae by their sugary texture
and conspicuous dark rims (see Fig. 3 in Tucker, 1983). Another less-common type of
cavity infill is composed of fine-grained, “internal” sediment (also recognized by Tucker,
1983); where present, the internal sediment appears to be the final cavity-filling phase.

General Relationships Among Textural Components: Microscopically, both light and
dark laminae show significant thickness variation (Figure 5.5, 5.6). Typical thicknesses

166
Figure 5.6: Image of merged photomicrographs showing lateral variability and continuity of light/dark lamination

Notice how laminae in the uppermost portion are well defined compared to the bottom portion. Cavity structures (C) shown in the
central portion of light laminae and in association with more pronounced neomorphism. Also highlighted are zones where
neomorphism crosscuts micritic laminae (X) and exhibits a patchy appearance (P).

167
for both laminae range from 50 to 1500 µm, with the vast majority falling between 400 to
1000 µm. In nearly all instances the contacts between light and dark laminae are diffuse
(Figures 5.5, 5.6 and 5.7), and a precise boundary between the two is often difficult to
pinpoint; sharp, dissolution fabrics were not observed. Cavity structures always occur
entirely within the light laminae, are usually situated in the central portions of the light
laminae, and are associated with regions exhibiting more intense neomorphism than
surrounding areas (Figures 5.6 and 5.7). In many cases, multiple cavity structures occur
within a single light lamina and can be connected by thin seams now occluded by
isopachous cements (Figure 5.7).

Isotopic Analysis
δ
18
O: Oxygen isotope values of the lower BSD range from –8.1 to +2.2‰ (VPDB).
Analysis of individual fabrics reveals a progressive decrease in
18
O from dark laminae to
light laminae and finally to cavity cements (only spar-filled cavities were drilled for
isotopic analysis; internal sediment was avoided). Average values for all three groups are
–1.8‰ for dark laminae, –3.0‰ for light laminae, and –6.1‰ for spar-filled cavity
structures (respective standard deviations are 1.6, 1.3, and 1.1). The central portions of
cavities were targeted for analysis, but drilling of fibrous cavity rims may have occurred
in some samples. The isotopic values of individual hand-samples are listed in Table 5.1
and shown graphically in Figure 5.9. Except for sample BS7 (discussed below), all light
laminae are isotopically depleted versus their nearest dark laminae neighbor.

168
Figure 5.7: Merged photomicrographs showing relationship of cavity structures to light
laminae

Cavity structures (C) always occur within lighter, neomorphosed regions and are
associated with the zones of highest neomorphism. Multiple cavity structures can occur
in a single light lamina and can be connected by thin seams (S). Also shown are regions
where neomorphism crosscuts micritic laminae (X). Inset shows close-up view of typical
cavity structure cements. The inward crystallographic transition is as follows: host rock
(H), isopachous fringe (I), equant sparry fill (F).

169
δ
13
C: Carbon isotope values of the lower BSD mostly range from +2 to +6‰ (VPDB)
and show no systematic variation among individual hand samples for the different
sedimentary features. The total carbon isotope range exhibited by the lower BSD is from
–5.3 to +5.8‰. The average values and their standard deviations are +2.9 ± 2.5‰ for
dark laminae, +3.6 ± 1.4‰ for light laminae, and +2.8 ± 1.6‰ for cavity infill; that is,
their ranges overlap.

Sample BS7: Sample BS7 shows the opposite isotopic relationship between dark and
light laminae compared to the other samples analyzed. Petrographic examination reveals
that the lamination of this sample is distinctly different from samples BS1-6. Regions
appearing light gray in hand sample are not composed of coarse-crystalline, micro-
pseudospar, but instead of sedimentary micrite or crystal silt. In addition, lamination is
poorly defined in BS7, and finer-grained portions do not contain the micrometer-scale
lamination typical of the other samples (Figure 5.8).

INTERPRETATION
Nature of the Light/Dark Lamination
The intricate textural features and relatively heavy oxygen isotope composition of
the micritic dark laminae in samples BS1-6 suggest that these laminae are the least
altered portions of the lower BSD. The wavy nature of the micrometer-scale lamination
contained within the millimeter-scale dark laminae may indicate a microbial-mat origin,
but like most ancient putative microbial structures (e.g., Grotzinger and Knoll, 1999), no

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Table 5.1: Carbon and oxygen isotope values of the Beck Spring Dolomite separated by
sample

Data are shown graphically in Figure 5.9. See Figure 5.4 for precise sample locations.

171
Table 5.1 continued

172
microfossils or other microbial features were noted during our investigation. The light
laminations, on the other hand, are composed of coarse microspar and pseudospar that are
reasonably interpreted as neomorphosed from the micritic host rock. The neomorphism
crosscuts the dark lamination and occurs in direct association with cavity structures.
Thus, the striking light/dark lamination exhibited by the lower BSD was generated
diagenetically, and not by primary sedimentary processes.
The parallel relationship of cavity structures, the micrometer-scale lamination and
millimeter-scale, light/dark fabric suggests that original sedimentological features
controlled neomorphism in the lower BSD. Elongate, bedding-parallel cavities have been
identified in Phanerozoic carbonates (so called laminoid fenestrae, stromatactis, or
birdseye structures), and these structures have been attributed to a number of
explanations—some controversial—including desiccation (Fischer, 1964), organic-matter
decay (Bathurst, 1959; Shinn, 1968; Flajs and Hüssner, 1993), and as many as twenty
other hypotheses (see review in Flügel, 2004). However, a common theme among the
myriad of explanations is that the cavity structures were once open voids through which
fluids could flow, given interconnectivity (Ham, 1954; Bathurst, 1959; Wolf, 1965;
Fischer, 1964; Deelman, 1972; Pratt, 1982). The direct association between highly
neomorphosed regions and cavity structures in the lower BSD suggests that
recrystallizing fluids flowed through former cavities and preferentially neomorphosed
surrounding regions. The high density and interconnectivity of cavity structures yielded
the necessary permeability to allow pervasive neomorphism.

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Figure 5.8: Photomicrographs of the abnormal, light laminae of BS7

Crystal silt (cs) and breccia-containing internal sediments are clearly recognizable. Regions appearing dark in thin section appear
light in hand sample.

174
Why would this style of neomorphism generate a planar fabric of such an
extensive nature and not simply appear as isolated patches surrounding cavity structures?
Given the small grain size of the host rock, intergranular permeability would have been
low (Lucia, 1995), restricting neomorphism to cavity-proximal regions. However,
carbonate sediments generally exhibit higher horizontal permeability compared to
vertical permeability (Lucia, 2007), such that fluid migration is favored in a fabric- or
bedding-parallel direction. Diagenetic recrystallization would follow the same general
trend and could produce a fabric-parallel neomorphic lamination. Thus, interconnected
cavities (so called “touching-vugs”, Lucia, 1995) likely provided the major source of
recrystallizing fluids. Upon infiltration into the host rock, these fluids flowed in a fabric-
parallel direction, yielding the extensive lamination exhibited by the lower BSD.

Diagenetic Progression
The petrographic characteristics and isotopic values exhibited by the lower BSD
are consistent with progressive diagenesis involving aggrading neomorphic
recrystallization (Folk, 1965; Bathurst, 1975) by an
18
O-depleted fluid. Depleted δ
18
O
values in carbonates are likely reflective of precipitation or recrystallization from a fluid
of meteoric origin, in as much as these waters are typically enriched in
16
O (Gross, 1964;
Allan and Mathews, 1982; Lohmann, 1988) due to Raleigh fractionation processes
involved in the evolution of rain clouds (Dansgaard, 1964). Previous authors (Gustadt,
1968; Marian, 1979; Marian and Osborne, 1980) have concluded that deposition of the
BSD took place on a passive-margin, nearshore carbonate platform, suggesting that

175
influence from continentally derived waters was reasonable during diagenesis.
Zempolich (1989) noted extensive meteoric diagenesis in the upper part of the BSD, and
Kenny and Knauth (2001) reported a karst surface at the top of the BSD, indicative of
subaerial exposure and meteoric weathering. Furthermore, if the micrite was initially
aragonite (as suggested by Zempolich, 1989), contact with meteoric fluids would
certainly promote neomorphism to calcite (Folk, 1965) and perhaps dissolution
(formation of cavities).
An alternate hypothesis for decreasing δ
18
O is carbonate precipitation from fluids
experiencing elevated temperatures, likely during burial. Increasing temperatures reduce
the isotopic fractionation between precipitating carbonate minerals and the host fluid,
producing cements with decreasing δ
18
O values (McCrea, 1950; Epstein et al., 1953;
O’Neil et al., 1969; Kim and O’Neil, 1997).
The decreasing oxygen isotope progression from dark laminae to light laminae to
sparry cavity infill suggests that these carbonates experienced at least two, isotopically
distinct stages of recrystallization. The relatively heavy isotopic composition and
retention of an intricate, micrometer-scale, wavy lamination in the dark laminae is
consistent with it being the least-altered sedimentary fabric. The intermediate oxygen
isotope composition and coarser-crystalline, neomorphic texture of the light laminae and
its crosscutting tendencies place it as the second stage of diagenetic recrystallization. The
most
18
O-depleted values of the cavity-filling spar suggest that these are the latest-stage
diagenetic crystal phases exhibited by the lower BSD. We suggest that (1) fluid flow
began parallel to the sedimentary fabric, initially resulting in neomorphosed layers that in

176
Figure 5.9: Carbon and oxygen isotope values for each Beck Spring sample

177
Figure 5.9 continued

178
places “matured” into cavity structures, or (2) cavities acted as conduits for fluids
involved in initial fabric-parallel neomorphism and were occluded by a subsequent
crystallization event.
Both of the above hypotheses can account for the large difference between the
δ
18
O of light laminae (avg. –3.0‰) and cavity-infilling cements (avg. –6.1‰) given the
following additional circumstances. If 1 is correct, then the “maturation” of cavity
structures must have taken sufficient time in order for pore-water δ
18
O to decrease ~ 3‰.
In 2 above, the conduits must have remained open long after initial deposition, which is
possible in conditions where early marine precipitates (i.e., isopachous rims) can stabilize
open vugs.
It is important to note that δ
18
O values depend on carbonate mineralogy.
Typically, dolomites exhibit δ
18
O ~ 2.6-9‰ higher compared to coeval calcites (O’Neil
and Epstein, 1966; Northrop and Clayton, 1966; Vasconcelos et al., 2005). Dolomites of
the BSD are close to stochiometric (Ca/Mg molar ratio ~ 1) and show no appreciable
deviation in Mg concentration (Tucker, 1982); therefore the deviation in δ
18
O does not
arise from differences in Mg content.
A lack of consistent deviation in δ
13
C among the different textures in the BSD is
not inconsistent with recrystallization from a meteoric or high-temperature fluid. While
these fluids are typically enriched in both
16
O and
12
C (Allan and Mathews, 1982), the
concentration of carbon in recrystallizing fluids is much less than that of oxygen. In the
case of carbonates, where ~ 10-12% of the solid phase is carbon, the rock will act as a
buffer against alteration in δ
13
C by recrystallizing fluids (Banner and Hanson, 1990).

179
Thus, the large spread in δ
13
C (and lack of covariation with δ
18
O) in individual samples
and among multiple samples likely reflects the original carbon isotope signal (that is, the
signal yields no information about diagenetic modification). Additionally, diagenetic
recrystallization would be expected to homogenize an isotopic signal and dampen high
variability, a feature exhibited by δ
18
O and not by δ
13
C.

Sample BS7: The abnormal textural and isotopic characteristics of BS7 indicate that the
lamination of the entire lower BSD is complex and does not always follow the general
trend recognized here. A sedimentary origin for the light-gray laminae is consistent with
a relatively enriched δ
18
O signal. This less altered sedimentary micrite would be
expected to retain an isotopically heavy, more marine-like signal. The absence of a
micrometer-scale, wavy lamination in the relatively dark laminae of BS7 suggests that
these regions have undergone obliterative neomorphism, just as the light laminae of
samples BS1-6. Relatively light δ
18
O values in the dark laminae are also consistent with
a higher degree of recrystallization in sample BS7.

Further Implications: Precambrian Versus Phanerozoic Layer-Parallel Diagenesis
As noted above, layer-parallel diagenesis is recognized in carbonates of all ages,
but what sets the BSD apart is laterally pervasive, layer-parallel diagenesis on such a
fine-scale—the micrometer to millimeter scale over tens of kilometers. In Phanerozoic
deposits, extensive horizontal fabrics with intricate features such as cavities and
micrometer-scale laminae are typically disrupted and destroyed by burrowing metazoa in

180
shallow subtidal settings. However, during nearly all of the Precambrian, bioturbation
was absent and extensive laminated fabrics could form in most marine environments
(Cloud, 1968) and persist long enough to experience diagenetic modification. Thus, the
diagenetic morphology of the lower BSD may represent an as of yet unreported mode of
carbonate diagenesis restricted to the Precambrian or to Phanerozoic environments
devoid of vertical bioturbation.
Furthermore, the extensive colonization of the seafloor by microbial mats on
Precambrian carbonate platforms would have provided a widespread process for forming
extremely thinly laminated sediments, which in turn would provide the interconnected
fabric to allow extensive, but very fine-scale, fluid flow to occur. The transition from
extensive Precambrian “matgrounds” (mat-capped, well-laminated sediments) to the
typical Phanerozoic-style sediment-water interface (bioturbated) has been called the
“agronomic revolution” (e.g., Seilacher and Pflüger, 1994), and evolutionary changes that
followed the substrate modification have been called the “Cambrian substrate revolution”
(e.g., Bottjer et al., 2000). The impacts of burrowing macrofauna on diagenesis have
been considered for siliciclastic settings (McIlroy and Logan, 1999). Here we suggest
that a carbonate “diagenetic revolution” also took place between the Neoproterozoic and
the Cambrian, and that the style of diagenesis revealed in the lower BSD may be typical
of pre-bioturbated Precambrian carbonate platforms. Thus, previously unrecognized,
extensive, layer-parallel diagenetic processes may be a hallmark of Precambrian shallow
marine carbonates, and may be much wider spread than previously thought.

181
CONCLUSIONS
The prominent light/dark planar lamination of the lower Beck Spring Dolomite
has been routinely interpreted to be a primary sedimentary fabric. Upon more intensive
petrographic examination, it seems more likely that the easily recognized, millimeter-
scale lamination is in fact a diagenetic feature. Light laminae of the lower BSD are
composed of coarse-crystalline micro-pseudospar that can be seen crosscutting finer-
grained, micritic, dark laminae in thin section. These light laminae are likely the result of
aggrading neomorphic, fabric-destructive recrystallization of the darker host rock, an
interpretation supported by the depleted δ
18
O values in the light layers compared to those
in the dark layers. The relationship between the wavy lamination, the elongate cavity
structures, and the millimeter-scale lamination suggests a primary sedimentological
influence on preferential recrystallization, producing a diagenetic fabric orientation that
mimics the primary depositional fabric orientation. These findings suggest that
diagenesis can cause the development of extensive, millimeter-scale, planar lamination in
Precambrian carbonates and that lamination origin cannot always be discerned solely by
outcrop examination. Finally, such extensive, fine-scale, layer-parallel diagenesis may be
common in Precambrian carbonate strata, deposited before the onset of bioturbation.

.

182
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Asset Metadata

Creator Loyd, Sean J. (author)

Core Title Carbonate geochemistry in primary, diagenetic and biological systems

Contributor Electronically uploaded by the Library (provenance)

School College of Letters, Arts and Sciences

Degree Doctor of Philosophy

Degree Program Geological Sciences

Publication Date 08/12/2010

Defense Date 05/01/2010

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

Tag Carbon,carbonate,concretions,isotopes,Neoproterozoic,OAI-PMH Harvest,Sulfur

Language English

Advisor Corsetti, Frank A. (committee chair), Berelson, William M. (committee member), Bottjer, David J. (committee member), Hammond, Douglas E. (committee member), Ziebis, Wiebke (committee member)

Creator Email jenni.bjelland@gmail.com,loyd@usc.edu

Permanent Link (DOI) https://doi.org/10.25549/usctheses-m3345

Unique identifier UC151694

Identifier etd-Loyd-3795 (filename),usctheses-m40 (legacy collection record id),usctheses-c127-378823 (legacy record id),usctheses-m3345 (legacy record id)

Legacy Identifier etd-Loyd-3795.pdf

Dmrecord 378823

Document Type Dissertation

Rights Loyd, Sean J.

Type texts

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

Repository Name Libraries, University of Southern California

Repository Location Los Angeles, California

Repository Email uscdl@usc.edu

Abstract (if available)

Abstract The carbonate minerals calcite, aragonite and dolomite (and their rock-counterparts) precipitate directly from fluids. The mineral-yielding fluids must contain the necessary chemical constituents calcium, magnesium, carbon and oxygen. As the carbonates precipitate they inherit and incorporate chemical signatures that are ultimately governed by the nature of formation fluids. Therefore, carbonate rocks and minerals can be treated as geologic reservoirs for information concerning past fluid chemistries and very powerful geochemical databases.