AN ABSTRACT OF THE DISSERTATION OF
Rebecca L. Poulson for the degree of Doctor of Philosophy in Oceanography presented on
June 12, 2008.
Title: The Influence of Early Diagenesis on Trace Element and Molybdenum Isotope
Geochemistry
Abstract approved:
______________________________________
James McManus
This thesis investigates the influence of early diagenesis on trace metal and
molybdenum isotope behavior in marine and lacustrine environments. Chapter one is a
synthesis of previous research in all the marine environments investigated, providing an
essential geochemical context for interpreting the observed behavior of Mo in these settings.
Chapter two discusses Mo behavior in three sites from the anoxic Mexican continental
margin. The data from these sites suggest that a unique Mo isotopic signature exists for
authigenic Mo enrichments in anoxic sediments. Chapter three discusses Mo geochemical
and isotopic behavior from a variety of marine environments to further constrain Mo
behavior during early diagenesis. At sites representing end-member cases for oxic and
anoxic conditions, the observed sediment Mo isotope compositions agree with those
predicted from previously reported natural and laboratory fractionations. However, data
from surface sediments of several study sites suggest that Mo associated with organic matter
has an isotopic composition that is less fractionated (relative to modern seawater) than either
oxic or anoxic authigenic Mo phases, and that this biogenic Mo may dominate the bulk
sediment Mo pool in certain environments. In addition, redox cycling of Mn within the
sediment column appears to strongly influence Mo geochemical and isotopic behavior.
Chapter four investigates sediment geochemistry along a depth transect in Lake Tanganyika,
East Africa. Permanent stratification of waters in this lake has produced a strong chemocline,
with oxic conditions in the surface and sulfidic waters at depth. The sediment distributions of
trace metals (specifically Mo and U) along this transect are investigated herein to evaluate
changes in sediment geochemistry across this transition. Despite sulfur limitation in this
freshwater system, conditions are sufficiently reducing (particularly at depths below the
chemocline) to generate substantial authigenic metal enrichments.
© Copyright by Rebecca L. Poulson
June 12, 2008
All Rights Reserved
The Influence of Early Diagenesis on Trace Element and Molybdenum Isotope
Geochemistry
by
Rebecca L. Poulson
A DISSERTATION
submitted to
Oregon State University
in partial fulfillment of
the requirements for the
degree of
Doctor of Philosophy
Presented June 12, 2008
Commencement June 2009
Doctor of Philosophy dissertation of Rebecca L. Poulson presented on June 12, 2008.
APPROVED:
____________________________________________________
Major Professor, representing Oceanography
____________________________________________________
Dean of the College of Oceanic and Atmospheric Sciences
____________________________________________________
Dean of the Graduate School
I understand that my dissertation will become part of the permanent collection of Oregon
State University libraries. My signature below authorizes release of my dissertation to any
reader upon request.
____________________________________________________
Rebecca L. Poulson, Author
ACKNOWLEDGEMENTS
First and foremost I would like to thank my primary advisor and dear friend, Dr. Jim
McManus. Without his mentorship, guidance, support, and unfailing sense of humor, none of
this would have been possible. I would also like to extend my sincere gratitude to the faculty
members of my graduate committee - Dr. Clare Reimers, Dr. Gary Klinkhammer, and Dr.
Dave Graham - for donating their time and providing insightful comments throughout this
process. I thank my friend and colleague Dr. Silke Severmann, who convinced me all this
was possible, and provided invaluable assistance and support all along the way. I would also
like to thank Dr. William Berelson, for providing many of the sediment samples discussed in
this thesis, as well as thoughtful comments on the resulting manuscripts. I thank Dr.
Christopher Siebert for teaching me how to make the measurements in the first place, and
Angela Bice for doing all the dirty work. I would also like to thank Andy Ungerer, whose
tireless efforts in the Keck Collaboratory made all of the analyses possible. I thank Rob
Wheatcroft, Anna Pakenham, and Rhea Sanders for assisting with gamma detection. I thank
the IGERT Subsurface Biospheres program (NSF #0114427) for providing me with
invaluable research opportunities and financial assistance. In particular, I thank Julie Cope,
who always has the answers. This work was funded by several NSF grants (OCE-0219651,
EAR-0518322, OCE-0551605, OCE-0721102) and I am sincerely grateful for the
opportunities afforded me. I would also like to thank the Nyanza Project - Andy Cohen,
Kiram Lezzar, Catherine O'Reilly, Ellinor Michel, and Hudson Nkotagu - for allowing me to
participate in their fantastic program and providing me with essential infrastructure that
made all the African research possible. I thank my office mates Mark Nielsen and Jesse
Muratli for tolerating my endless banter and contributing countless essential tidbits along the
way. To all the Friends of the Plasma, I wish to extend my sincerest thanks - you have made
COAS a fun and interesting place to work, and have been true friends throughout this
process. I thank my family, who has always supported me in all my endeavors - your
unconditional love and encouragement has never waned, and I would never have been able
to accomplish all that I have without you. And finally, my Jake - for tolerating my stresses
and sharing in my successes throughout each and every step of this process, I am forever
grateful.
CONTRIBUTION OF AUTHORS
Dr. William Berelson provided the Mexico margin sediment samples discussed in Chapters 2
and 3 and provided thoughtful comments on both manuscripts. Dr. Christopher Siebert
provided analytical assistance on the Mo isotope measurements and contributed helpful
comments on Chapter 2. Dr. Silke Severmann assisted with the collection and preparation of
sediment samples discussed in Chapters 3 and 4, supplied additional sediment data discussed
in Chapter 4, and provided insightful contributions to both the Chapter 3 and 4 manuscripts.
TABLE OF CONTENTS
Page
General Introduction ...................................................................................
2
Abstract .........................................................................................
3
Introduction ...................................................................................
3
Mo in the Marine Environment .....................................................
5
Mo Behavior During Early Diagenesis .........................................
9
Study Sites for this Research ........................................................
12
MANOP Sites ..................................................................
12
Peru Margin ......................................................................
15
California Margin .............................................................
18
Mexico Margin .................................................................
29
Summary .......................................................................................
36
References ....................................................................................
37
Authigenic Molybdenum Isotope Signatures in Marine Sediments .......... .
54
Abstract .........................................................................................
55
Introduction ...................................................................................
55
Authigenic Molybdenum ...............................................................
57
TABLE OF CONTENTS (Continued)
Page
Results ...........................................................................................
59
Discussion .....................................................................................
60
Conclusions ...................................................................................
62
Acknowledgments .........................................................................
63
References ....................................................................................
64
Molybdenum Behavior During Early Diagenesis: Insights from Mo
Isotopes ......................................................................................................
71
Abstract .........................................................................................
72
Introduction ...................................................................................
73
Molybdenum in the Marine Environment .....................................
74
Lithogenic Mo ..................................................................
75
Biogenic Mo .....................................................................
75
Authigenic Mo ..................................................................
76
Site Descriptions ............................................................................
79
MANOP Sites ..................................................................
79
Peru Margin ..................................................................... .
79
California Margin .............................................................
80
Mexico Margin .................................................................
81
Methods .........................................................................................
83
Results and Discussion ..................................................................
85
The Lithogenic Mo Correction ..........................................
86
TABLE OF CONTENTS (Continued)
Page
The Authigenic Mo-Manganese Signature ........................
87
The Authigenic Mo-Sulfide Signature ...............................
89
The Biogenic Mo Signature ................................................
91
Authigenic Mo with Seawater δ98Mo (aq) Source .................
92
Authigenic Mo with Manganese δ98Mo (aq) Source .............
95
Conclusions ......................................................................................
98
References ........................................................................................
100
Sediment Geochemistry along a Chemocline Transect,
Lake Tanganyika, East Africa ........................................................................
117
Abstract .............................................................................................
118
Introduction .......................................................................................
119
Methods .............................................................................................
121
Results and Discussion ......................................................................
123
Age Model ............................................................................
123
Regional Sedimentation ........................................................
125
Sediment Distributions of Diagenetic Reactants ..................
126
Sediment Trace Metal Distributions .....................................
129
Authigenic Accumulation of U and Mo .................................
132
Conclusions .........................................................................................
135
References ...........................................................................................
136
Conclusion .....................................................................................................
159
TABLE OF CONTENTS (Continued)
Page
Bibliography ..............................................................................................
161
Appendix A: Additional Data Tables for Chapter 3 ..................................
170
LIST OF FIGURES
Figure
Page
1.1 Major Mo sources to modern marine sediments .................................
43
1.2 Published marine Mo isotope values and fractionation factors ...........
43
1.3 Schematic summarizing Mo behavior under various diagenetic
regimes ..................................................................................................
44
1.4 Map of study areas ..............................................................................
45
1.5 Map of Peru margin and MANOP sites ...............................................
45
1.6 Pore water profiles and generalized diagenetic regimes for MANOP
sites M and H .......................................................................................
46
1.7 Pore water profiles and generalized diagenetic regimes for Peru
margin .................................................................................................
47
1.8 Map of California margin Borderland Basin sites ..............................
47
1.9 Pore water profiles and generalized diagenetic regimes for the three
inner basins of the California margin ..................................................
48
1.10 Pore water profiles and generalized diagenetic regimes for the four
outer basins of the California margin ...............................................
49
1.11 Pore water profiles of sulfate, ammonia, and phosphate for all
Borderland basin sites investigated in this study ..............................
50
1.12 Map of Mexico margin sites .............................................................
51
1.13 Pore water profiles of ammonia and sulfate for two Mexican margin
sites, the Peru margin, and all Borderland basin sites investigated in
this study ...........................................................................................
51
2.1 Measured Mo isotope compositions of various marine sediments .....
66
2.2 All δ98Mo data (without lithogenic correction) from down-core
profiles .................................................................................................
67
LIST OF FIGURES (Continued)
Figure
Page
2.3 Schematic of the authigenic Mo isotope system in marine
sediments .............................................................................................
68
3.1 Map of study areas ..............................................................................
104
3.2 Major Mo sources to modern marine sediments .................................
104
3.3 Published marine Mo isotope values and fractionation factors ..........
105
3.4 Map of California margin study areas .................................................
105
3.5 Map of Mexico margin study areas ....................................................
106
3.6 MANOP Site H Profiles ......................................................................
107
3.7 Peru Margin Profiles ............................................................................
108
3.8 Sediment Mo enrich concentrations and isotope compositions from
Pescadero margin .................................................................................
109
3.9 Sediment Mo enrich concentrations and isotope compositions from
Magdalena margin ...............................................................................
110
3.10 Sediment Mo enrich concentrations and isotope compositions from
Alfonso and La Paz basins, Mexico margin ......................................
110
3.11 Pore water Fe and Mn profiles, sediment Mo enrich concentrations,
and Mo isotope compositions from Santa Catalina and San Nicolas
basins .................................................................................................
111
3.12 Whole core average Mo enrich concentrations versus δ98Mo enrich values
for all sites in this study .....................................................................
112
4.1 Map of study area: Lake Tanganyika, Tanzania ..................................
139
4.2 Water column profiles from Northern Lake Tanganyika,
Kigoma Basin .......................................................................................
140
4.3 Down-core profiles of 210Pb and 137Cs used to create age models
for all Luiche Platform sites .................................................................
141
4.4 Down-core profiles of sediment carbonate contents for all
Luiche Platform cores ..........................................................................
143
LIST OF FIGURES (Continued)
Figure
Page
4.5 Sediment Ti concentrations (%) versus estimated sedimentary
lithogenic fractions (X LITH ) for all study sites ..........................................
143
4.6 Average total sediment, carbonate, and Ti accumulation rates for
all study sites ............................................................................................
144
4.7 Whole core average organic carbon (C org ) and total reducible
iron sulfide (TRIS), C org and TRIS accumulation rates, and sediment
C/S (C org /TRIS) ratios for all study sites ..................................................
145
4.8 Individual sample (left panel) and whole core average (right panel)
sediment organic carbon (C org ) versus total reduced iron sulfide (TRIS)
for all study sites ........................................................................................
146
4.9 Whole core average sediment Mn:Ti and Fe:Ti ratios for all study sites ....
147
4.10 Pore water Mn and Fe profiles for all study sites .....................................
147
4.11 Whole core average predicted and measured trace metal
concentrations .........................................................................................
148
4.12 Whole core average "enriched" U and Mo accumulation rates ..................
149
4.13 Whole core average "enriched" U and Mo versus whole core average
organic carbon (%C org ) and total reduced inorganic sulfur (%TRIS) .........
149
LIST OF TABLES
Table
Page
1.1 Study Site Characteristics ....................................................................
52
2.1 Mo Isotope Compositions of Various Marine Depositional
Environments .......................................................................................
69
3.1 General site characteristics and average sediment Mo enrich data ..........
113
4.1 Water depths, locations, and sedimentation data for all study sites ....
150
4.2 Sediment data: major elements, carbon and sulfur contents, carbonate
fractions, estimated densities, and mass accumulation rates ................
151
4.3 Sediment trace metal data .................................................................... .
155
4.4 Average standard reference material compositions ..............................
158
LIST OF APPENDIX TABLES
Appendix Table
Page
1.1 Standard reference materials for Mo, Al, Ca, Fe, Mn, and Ti .............
171
1.2 Sediment Mo concentration data ..........................................................
172
1.3 Sediment major element compositions and lithogenic Mo fractions ...
181
1.4 Sediment Mo isotope compositions .......................................................
192
1.5 Average sediment Mo enrich concentrations and isotopic compositions ...
196
1.6 Detailed sediment data from previously published California and
Mexico margin sites ...............................................................................
201
The Influence of Early Diagenesis on Trace Element and Molybdenum Isotope
Geochemistry
Rebecca L. Poulson
2
GENERAL INTRODUCTION
3
ABSTRACT
This thesis focuses on molybdenum and molybdenum stable isotope geochemistry in
marine and lacustrine sedimentary settings. In addition, this work examines a number of
other redox-sensitive trace metals (e.g. U, Cd, and Re) whose behavior may also vary
through the range of environments captured in this study. The sites that are studied have
been chosen to represent a broad range in sedimentary oxidation-reduction (redox) potential.
Sediments have been analyzed from a variety of locations, including the California
continental margin, the Mexico margin, the Peru margin, and two Eastern Pacific hemipelagic sites. Decades of previous research at all the study sites provide the necessary
diagenetic framework essential for interpreting redox-sensitive trace metal behavior. The
purpose of this chapter is to review what is known about Mo geochemistry, to summarize the
available geochemical data from these locations, and to provide a framework for the trace
metal and Mo isotope data discussed in subsequent chapters.
INTRODUCTION
Information regarding the geochemical evolution of the global oceans is contained
within the marine sediment and rock records. Interpretation of these ancient records
leverages off our often imperfect interpretation of the chemical signatures that are
sequestered within remnant geologic materials. We therefore continue to refine our
understanding of proxy behaviors by investigating modern systems with characteristics that
we seek to uncover from the geologic record.
4
This thesis focuses on the modern geochemical behavior of molybdenum and its
stable isotopes along with a number of other elements (e.g. U, Cd, and Re). These trace
elements have solubilities that are sensitive to the oxidation-reduction (redox) potential of
the environment. Because of this sensitivity, these elements have been employed as proxies
for the geochemical conditions of ancient depositional environments (e.g. Tribovillard et al.,
2006, and references therein). Specific to this study, the solubility of Mo appears to be
sensitive to the availability of reduced sulfur species; sediment Mo enrichments have been
interpreted to indicate a lack of dissolved oxygen (e.g., Crusius et al., 1996; Dean et al.,
2006). However, Mo behavior is also impacted by manganese and iron cycling in more
oxygenated environments (e.g., Bertine and Turekian, 1973; Calvert and Pedersen, 1993;
Chappaz et al., 2008). Sediment Mo enrichments thus occur in both well-oxygenated and
reducing marine settings, and Mo concentrations alone are therefore difficult to interpret in
terms of depositional redox conditions.
Mo isotopes in conjunction with elemental ratios may provide a more robust
paleochemical proxy. Molybdenum has seven stable isotopes (92, 94, 95, 96, 97, 98, and
100) with relative natural abundances from 9 to 24%. From an analytical perspective, we
take advantage of the multiple isotopes of Mo by measuring Mo isotope ratios by the
published 100/97Mo double spike technique (Siebert et al., 2001; all values reported in δ98Mo
notation: δ98Mo = [(98/95MoSAMPLE/98/95MoSTANDARD -1) x 1000]). Recent work has identified
marine sediment Mo isotope signatures unique to the dominant mechanisms controlling Mo
speciation and enrichment (Barling et al., 2001; Siebert et al., 2003; 2006; Poulson et al.,
2006). Laboratory experiments and natural samples have quantified Mo isotope fractionation
in Mn-dominated systems (Barling et al., 2001, 2004; Siebert et al., 2003), but sulfidecontrolled Mo isotope fractionations remain poorly understood.
5
It is the goal of this study to further constrain Mo isotopic behavior in the marine
environment by analyzing well-characterized modern marine sediments from a range of
depositional redox conditions. Decades of previous research at these study locations provide
a framework for interpreting Mo behavior, and this chapter is an attempt to summarize and
synthesize that work. This thesis combines published laboratory experimental results with
observations from modern environments to fully characterize Mo geochemistry and
associated isotopic fractionations in modern marine systems, providing a necessary context
for future proxy applications.
Mo IN THE MARINE ENVIRONMENT
Under the oxygenated conditions predominant in the modern ocean, Mo exists
primarily as the soluble molybdate ion (MoO42-; Figure 1) and is the most abundant
dissolved trace element in seawater (Broecker and Peng, 1982). Molybdenum behaves
conservatively in the open ocean water column, with a concentration of ~105 nM and a
residence time of ~800,000 years (Collier, 1985; Emerson and Huested, 1991). Although
there are currently a limited number of analyses (n = 6), modern seawater is thought to have
a homogenous Mo isotopic composition of δ98MoSW = 2.3±0.1‰ (Figure 2, Barling et al.,
2001; Siebert et al., 2003). This seawater value is thought to be dictated by the balance
between a near-zero input term (rivers) and the various sedimentary sink terms described
below (Arnold et al., 2004; McManus et al., 2006).
Despite the fact that seawater appears to be a uniform Mo reservoir, marine
sediments exhibit a wide range in Mo concentrations and isotope values because bulk
sediments reflect contributions from multiple sources or processes (Figure 1). With respect
6
to marine sediments, Mo deposits can be thought of as the sum of three dominant processes:
1) incorporation of lithogenic Mo into bulk sediment through continental weathering; 2)
association of Mo with biological material which is delivered directly to the seafloor; 3)
precipitation or adsorption as an authigenic solid phase (under both oxic and anoxic
conditions; Figure 1).
Lithogenic Mo
Continental margin sediments include some quantity of terrigenous (lithogenic)
material; the importance of the detrital contribution relative to the total bulk sediment is
dependent upon the depositional environment. Though continental material delivers only a
small quantity of Mo to marine sediments (typically ~1 ppm Mo in most igneous and
sedimentary rocks; Turekian and Wedepohl, 1961), the lithogenic Mo contribution can
constitute a significant fraction of the bulk sediment Mo inventory (Figure 1). Analyses of
various terrigenous materials (e.g. granites, clastic sediments; n=12) constrain a homogenous
isotopic composition of δ98Mo = 0.0±0.2‰ (Figure 2; Siebert et al., 2003), and this value is
taken to represent the lithogenic (continentally-derived) Mo component in bulk sediment. As
discussed in Chapter 2 (Poulson et al., 2006), to constrain the isotopic composition of
sedimentary authigenic Mo, measurements of bulk sediment Mo isotope compositions
require correction for dilution by the lithogenic contribution.
Biogenic Mo
Molybdenum is considered a biologically essential trace element, playing a key
enzymatic role in a variety of processes, notably nitrogen fixation and nitrate reduction (e.g.,
Mendel and Bittner, 2006 and references therein). The relationship between organic matter
7
and Mo is complex because Mo is not only incorporated into cells, but it can also be sorbed
to organic material in the water column (Figure 1; Tribovillard et al., 2004). There is limited
data available to constrain a single organic matter Mo:C ratio. Reported Mo:C ratios in the
nitrogen-fixing cyanobacteria Trichodesmium erythraeum show large differences in the
Mo:C ratios of natural and cultured samples (23 and 3 μmol/mol, respectively; Tuit et al.,
2004). Available sediment trap studies report Mo:C ratios of ~9 nmol/mmol (Mazatlan
margin; Nameroff et al., 1996) and ~4 nmol/mmol (Santa Barbara Basin; Zheng et al., 2000)
in sediment trap materials. It is quite likely that Mo:C ratios in organic matter are variable,
as they are dependent upon multiple environmental factors. In addition to this compositional
variability, it is also likely that the preservation of Mo associated with organic material will
vary.
Recent experimental work has reported a -0.5‰ δ98Mo isotope fractionation
associated with biological assimilation of Mo (Figure 2; Wasylenki et al., 2007; Liermann et
al., 2005). Because it is unlikely that there is a single Mo:C ratio for organic matter, it is not
possible to unequivocally quantify the organic sedimentary Mo component in marine
sediments – at least based on the currently-available data base. However, as discussed in
Chapter 3, biogenic Mo (Mo associated with organic matter) represents a distinct fraction of
the total sediment Mo pool and (like the lithogenic Mo component) its isotopic contribution
must be considered (Figure 1).
Authigenic Mo
Under well-oxygenated sedimentary conditions, Mo is most commonly found sorbed
to solid-phase Mn and Fe-oxides. (Figure 1; e.g., Bertine and Turekian, 1973; Calvert and
Pedersen, 1993; Chappaz et al., 2008). Experimental work by Barling and Anbar (2004)
8
revealed a large (2.7 ‰) fractionation between soluble molybdate (MoO42-) and Mo sorbed
to Mn-oxides in the laboratory; that is, Mn-associated Mo bears a light isotopic signature
relative to the dissolved phase. Those findings are consistent with results from modern FeMn crusts and Mn nodules (δ98Mo = -1.0 to -0.5 ‰) (Barling et al., 2001; Siebert et al.,
2003), which demonstrate similar fractionation between seawater molybdate and Mnassociated Mo in natural samples (Figure 2). Negative Mo isotope compositions reported in
Mn-rich sediments from San Clemente Basin on the California margin (δ98Mo = -0.8±0.4‰)
also reflect the large fractionation between seawater Mo and Mn-associated Mo in sediments
(Siebert et al., 2006). The specific mechanisms behind these observed isotope fractionations
are not entirely clear; however, Mn-controlled authigenic Mo enrichments have the most
negative sediment Mo isotope compositions measured to date.
Under anoxic sedimentary conditions where sulfate reduction is a dominant electron
transfer process, Mo is sequestered into sediments through complexation with ambient
sulfide, forming less soluble thiomolybdates (MoOxS4-x2-) that may be scavenged by
sulfidized organic matter or Fe-sulfide phases such as pyrite (Helz et al., 1996; Zheng et al.,
2000). Helz et al. (1996) proposed a sulfide-controlled geochemical “switch” for Mo at ~10
μM H2S[aq] (Figure 3), where the dominant dissolved Mo phase abruptly transitions from
molybdate (MoO42-) to tetrathiomolybdate (MoS42). Experimental work has shown that, in
the presence of both H2S and S0-electron donors, thiomolybdate Mo(VI) may be reduced to
Mo(V) or Mo(IV) polysulfide anions (Vorlicek et al., 2004); it remains unclear whether
reduction of the metal itself is necessary for authigenic Mo accumulation. The pore water
work of Zheng et al. (2000) proposed two thresholds for Mo-sulfide formation (Figure 3); at
H2S[aq] concentrations of ~0.01 μM these authors proposed that Mo is removed from solution
via coprecipitation of Fe-Mo-S phases, at higher H2S[aq] concentrations (~10 μM) they
9
postulate that Mo precipitates independent of iron. It may be that the sulfide thresholds
proposed by Zheng et al. (2000) reflect changes in aqueous Mo speciation that impact solidphase Mo behavior. At low sulfide concentrations, thiomolybdate intermediate species
(MoOxS4-x2-) may dominate the aqueous phase and be scavenged by solid-phase Fe-sulfides,
while at higher sulfide concentrations tetrathiomolybdate (MoS42-) is likely to dominate,
precipitating independently as a solid phase Mo-sulfide.
An investigation of pore waters from Santa Monica Basin predicted a fractionation
(-0.7‰) between pore fluids and sediment Mo deposits under reducing conditions
(McManus et al., 2002). Initial data from anoxic sites on the Mexican continental margin
reported in Chapter 2 suggest a unique Mo isotopic signature of δ98Mo = 1.6±0.1‰ for Mosulfide sediment enrichments that is consistent with the predicted fractionation (Figure 2;
Poulson et al., 2006). The geochemical mechanisms responsible for the observed isotopic
fractionation have not been identified. Reported Mo isotope compositions from “suboxic”
surface sediments of the California margin span the full range between Mn-dominated and
more reducing environments (δ98Mo = -0.8 to 1.6‰; Figure 2; Siebert et al., 2006),
demonstrating the need for further refinement of the Mo isotope system in marine sediments
before it may be successfully employed as a paleoproxy in geochemical reconstructions.
Mo BEHAVIOR DURING EARLY DIAGENESIS
In environments with high organic matter rain rates, the microbially-mediated
breakdown of organic matter during early diagenesis proceeds through a well-known
sequence of electron donors (e.g., Froelich et al., 1979). Sedimentary organic matter is
oxidized by sequential reduction of the available oxidant with the greatest free energy
10
change (O2, NO3-, MnO2, Fe2O3, and SO42-). Figure 3 is a simplified depiction of the
diagenetic sequence as reflected in pore water profiles.
The depths of these diagenetic reactions are elastic - they may be compressed in the
top few centimeters of sediment, stretched out over long distances within the sediment
column, or even take place within the water column - depending upon the balance between
electron donors and electron acceptors in any given setting. From a practical sedimentary
perspective, oxygen penetration depth, availability of various electron donors, the rate of
organic matter delivery to the seabed, as well as advective and/or diffusive properties of the
sediments all play a role in determining the rates of early diagenesis in sedimentary
environments (e.g. Froelich et al., 1979; Hartnett et al, 1998). The products and reactants of
these reactions are observed in pore water profiles, and can be used to predict the dominant
geochemistry of diagenetic environments (as generalized in Figure 3; e.g. Froelich et al.,
1979).
Several additional geochemical parameters not depicted in Figure 3 are often used to
characterize depositional conditions. For example, ammonia is a known product of Fe and
sulfate reduction reactions, and pore water ammonia enrichments are often used to suggest
reducing sediment conditions (e.g. Froelich et al., 1979). Phosphate enrichments (above
those predicted by organic matter decomposition alone) are also used to indicate Fe
reduction in sediments; phosphate initially adsorbed to Fe-oxyhydroxides may be released
into solution upon Fe reduction (e.g. McManus et al., 1997).
The utility of Mo as a potential tool for tracking diagenetic processes in marine
sediments is that Mo cycling is intimately linked to the cycling of several principal redox
elements; not only is Mo associated with organic carbon deposited at the seafloor, but its
authigenic cycling is primarily controlled by Mn, Fe, and S behavior. Essentially, Mo
11
cycling is tied not only to the geochemical reactions taking place within the sediments, but
also the fuel that drives early diagenetic processes (i.e., organic matter).
As previously discussed, Mo is most commonly found sorbed to solid-phase Mn and
Fe-oxides under well-oxygenated conditions (Figure 3; e.g., Bertine and Turekian, 1973;
Calvert and Pedersen, 1993; Chappaz et al., 2008). Upon depletion of available oxygen (and
nitrate), Mn and Fe reduction ensue, releasing sorbed Mo back into solution. If Mn is
reoxidized within the sedimentary column during diagenesis, some dissolved Mo may readsorb to reoxidized Mn and/or Fe in this oxidized layer (Figure 3). Alternatively, some
dissolved Mo may diffuse downward in the sediment column, eventually reaching a depth
where sulfate reduction is a dominant diagenetic process; here authigenic Mo can be
removed from solution by the formation of insoluble Mo-sulfide complexes (see below and
Helz et al., 1996; Figure 3).
This study focuses on Mo geochemical and isotopic behavior during early diagenesis
at several continental margin sites of the eastern tropical North Pacific (Figure 4, Table 1). In
continental margin environments, decomposition of settling organic matter depletes oxygen
in the water column, which in turn increases organic matter preservation (via reduced
oxygen exposure time; e.g., Hartnett et al., 1998). Low bottom water oxygen concentrations
can exclude benthic macrofaunal communities (thus minimizing sediment mixing via
bioturbation). Model results suggest that, at total O2 fluxes less than ~100μmol/cm2/yr,
diffusion is the dominant sediment transport mechanism (Meile and Van Cappellan, 2003).
In combination, these environmental factors allow for the study of suboxic to anoxic
diagenetic processes in shallow sediments underlying low-oxygen bottom waters (e.g.,
Froelich et al., 1979). Continental margin sediments are generally characterized by high
12
sedimentation rates that generate ideal conditions for investigating early diagenetic processes
and the cycling of key elements associated with these reactions.
STUDY SITES FOR THIS RESEARCH
This study aims to further refine our understanding of Mo geochemical and isotopic
behavior during early diagenesis through analyses of surface sediment cores from previously
studied sites on the California, Mexico, and Peru continental margins, as well as two sites
from the Manganese Nodule Program (MANOP; Emerson, 1984; Figure 4). To varying
degrees, previous research at these locations allows for characterization of the regional
dominant diagenetic processes, providing an important framework for interpreting Mo
behavior.
MANOP Sites
Sediments were analyzed from two abyssal cores collected as part of the Manganese
Nodule Program (MANOP; Emerson, 1984) in the eastern tropical Pacific (Figure 5). These
two sites represent the most oxic conditions analyzed in this study; sediments are Mn-rich
(1-7 wt.% Mn, Lyle et al., 1984) and bathed in well-oxygenated waters (110-150 μM O2;
Bender and Heggie, 1984). These sites were selected because the anticipated high sediment
authigenic Mo concentrations (e.g. Bertine and Turekian, 1973) should have Mo isotopic
compositions representing an end-member case for authigenic Mo enrichment associated
with Mn-rich, oxic sediments.
13
Site M
MANOP site M is located 25 km east of the East Pacific Rise, at ~3100 m water
depth (Figure 5; Klinkhammer, 1980). This site is influenced by hydrothermal sedimentation
from the mid-ocean ridge crest, and was selected for the MANOP study to investigate supply
of transition metals by hydrothermal precipitates during early diagenesis (Lyle et al., 1984).
In addition to the hydrothermally-derived sediment component, site M has a high biogenic
input flux (130 μg Corg/cm2yr; sediments 1-2 wt.% organic carbon) and the highest input of
continental material of all five MANOP sites (Lyle et al., 1984). Site M is 1000 km off the
coast of Central America, in the source region of the westward-flowing North Equatorial
Current which efficiently delivers lithogenic material from the continent; surface sediments
contain 9-13 wt.% continentally-derived feldspars (Lyle et al., 1984).
Sediments at site M are marked by an oxidized zone in the top ~5 cm where oxygen
reduction generates a nitrate maximum, though pore water profiles indicate that all available
nitrate is reduced within the top ~10 cm of the sediment column (Figure 6; Klinkhammer,
1980). Evidence for Mn(IV) reduction occurs at 5 to 8 cm depth (Figure 6; Klinkhammer,
1980; Lyle et al., 1984), and the shallow depth of Mn reduction at this site may lead to net
export of Mn from site M sediments into overlying waters (Heath and Lyle, 1982). Pore
water profiles indicate that Fe(III) reduction begins ~15-20 cm (Figure 6; Klinkhammer,
1980), in agreement with an observed brown-green sediment color transition at 20 cm
signifying the Fe(III)-Fe(II) redox boundary (Lyle, 1983). A marked increase in dissolved
phosphate is also observed around 15-20 cm, likely reflecting the release of phosphate
adsorbed to ferric hydroxides with the onset of Fe reduction (Emerson et al., 1980).
Measured hydrogen sulfide is <1 μmol/kg throughout the uppermost 30 cm indicating little
14
or no sulfate reduction; however, increased NH4+ at depths greater than ~25 cm suggests
some sulfate reduction may be occurring at depth (Emerson et al., 1980).
Site H
MANOP site H is located in the Guatemala Basin at a depth of about 3600 m
(Klinkhammer, 1980), and was selected as a typical hemipelagic site for the MANOP study
(Lyle et al., 1984). This site was found to have an abundance of manganese nodules;
coverage estimated at ~300 nodules/m2 (Finney et al., 1984). These nodules have unusually
high Mn/Fe ratios, suggesting a diagenetic source for the Mn precipitated at the nodule
surfaces (Dymond et al., 1984). Modeled rates of bioturbation at site H decrease
dramatically below 5 cm sediment depth, though aqueous Mn profiles suggest mixing may
impact sediment geochemistry to a depth of at least 14 cm (Klinkhammer 1980; Kadko and
Heath, 1984; Figure 6).
Site H lies near the northern boundary of the eastward-flowing North Equatorial
Countercurrent, which carries nutrient-poor waters into the eastern Pacific (Lyle et al.,
1984). The organic carbon flux at site H is similar to that at site M (110 μg Corg/cm2yr), but
surface sediment organic carbon contents are slightly lower at site H (0.6 to 1.4 wt.%; Lyle
et al., 1984). Concentrations of continentally-derived feldspars (4-11 wt.%) are lower at site
H than site M, suggesting a smaller contribution of lithogenic materials (Lyle et al., 1984).
Site H currently lies at or just below the carbonate compensation depth, though it appears
calcite preservation occurred there during the last glacial maximum; carbonate
concentrations increase from ~1 wt.% in the surface sediments to ~20 wt.% below 15 cm
(Lyle et al., 1984).
15
Pore water profiles reveal high nitrate concentrations (>30 μmol/kg) throughout the
upper ~40 cm of sediment (Figure 6; Klinkhammer, 1980). Though nitrate reduction does
not completely remove all available nitrate at this site, pore waters indicate Mn(IV)
reduction begins between 10 and 15 cm depth (compared to ~5 cm at site M; Figure 6;
Klinkhammer, 1980). Site H sediments are marked by a subsurface maximum in solid-phase
Mn at ~9 cm, indicating Mn(II) re-oxidation at this depth (Lyle et al., 1984). Manganese
nodules are also most prevalent in the uppermost 10 cm of sediment (Finney et al., 1984).
No Fe (III) reduction is detectable in pore water profiles from the top ~40 cm of site H
(Klinkhammer, 1980). The absence of Fe(III) reduction at this site is also suggested by the
lack of a visible brown-green sediment color transition (sediments are brown to at least 10 m
depth; Lyle, 1983).
Peru Margin
Sediments of the Peru margin investigated in this work were collected from a 264 m
shelf site near 13oS (within the oxygen minimum zone; Suess et al., 1986; Figure 5). This
site represents the most reducing open-ocean conditions of all study sites, due to the high
organic carbon content and sulfidic nature of the sediments in this region (e.g. Reimers and
Suess, 1983; Froelich et al., 1988). The Peru sediments were chosen with the expectation of
high authigenic Mo concentrations (e.g. Böning et al., 2004), and the Mo isotope
composition of these sediments is taken to represent an end-member Mo isotopic signature
for authigenic Mo deposited under open-ocean reducing conditions.
The Peru margin is a region of wind-driven perennial upwelling, resulting in high
productivity in the surface waters and an associated intense oxygen minimum zone (OMZ,
<5 μM O2) in the water column between 50 and 650 m (Suess et al., 1986). In one instance,
16
total depletion of nitrate (along with the presence of hydrogen sulfide) was observed within
the water column in this region (Dugdale et al., 1977). The dominant surface current in the
region is the equator-ward oxygen-rich Peru Chile Current, which turns away from the coast
near 15oS (Wyrtki, 1967; Brockmann et al, 1980). North of ~15oS, the system is dominated
by the high salinity, nutrient-rich, low oxygen southward-flowing Peru Undercurrent that
flows from 5oS to 15oS, impacting the seafloor between 150-400 m depth (Hill et al., 1998).
Organic carbon is preferentially accumulated on the outer shelf and upper slope (100-450 m
water depth) between 11oS to 16oS, with strong bottom currents preventing significant
accumulation on the mid-slope, resuspending and depositing sediments deeper on the
continental slope (>2000 m; Reimers and Suess, 1983).
An early study by Henrichs and Farrington (1984) near 15oS reported that
denitrification was essentially complete within the uppermost 3 cm of sediment, with sulfate
reduction being the dominant mechanism for organic matter remineralization. These authors
observed H2S[aq] >200 μM in the surface pore waters of all sites located within the OMZ.
Radiotracer incubations reported in Rowe and Howarth (1985) from a similar transect (also
near 15oS) also concluded that sulfate reduction was the most important diagenetic process at
work within the sediments off Peru. Fossing (1990) performed similar incubations on
sediments from a transect near 15oS, and concluded that an average sulfate reduction rate of
~20 mmol SO42-/m2day was a best estimate for sediments within the OMZ. More than 50%
of the total sulfate reduction occurred within the upper 20 cm of the sediment column at
most sites, being most intense between 1 and 4 cm depth (Fossing, 1990). Unlike Henrichs
and Farrington (1984), Fossing did not observe detectable sulfide in sediment pore waters
until a depth of ~10-15 cm. This investigator found that solid-phase Fe-S was present
between the sediment surface and the depth of free sulfide, and concluded that sulfate
17
reduced to sulfide in the top ~15 cm of sediment was either reoxidized or precipitated with
ferrous Fe. These findings are consistent with those of Froelich et al. (1988) from sediments
near 12oS, where low concentrations of dissolved Fe were found only to a depth of ~8 cm,
with detectable sulfide below (Figure 7).
Böning et al. (2004) analyzed shelf and slope sediments of the Peru margin between
9oS and 14oS, with sites covering depths above, within, and below the OMZ. These authors
observed that Mn was depleted in all sediments relative to the lithogenic background,
suggesting reduction of Mn-oxides within the suboxic water column, not within the
sediments. The highest sulfate reduction rates were measured within the upper ~5 cm at all
sites, with the highest total rates observed from sites at ~200 m water depth. Free sulfide was
detected below ~10 cm in the sediment column at sites in ~100 m water depth, and below
~20 cm at sites ~250 m deep (Böning et al., 2004).
Böning et al. (2004) also analyzed the sediments for Mo, reporting the highest
concentrations in sites < 300 m water depth (≤150 ppm Mo in uppermost ~20 cm at 255 m).
These authors suggested that the presence of H2S[aq] near the sediment-water interface (rather
than delivery via incorporation with organic matter) was the most likely cause of Mo
accumulation for sites within the OMZ. Because the Peru margin represents the most
reducing conditions of all sites analyzed, these sediments are expected to have the highest Scontrolled authigenic Mo concentrations measured herein. This site was selected to represent
an end-member for the authigenic sediment Mo isotope signature, where open-ocean anoxic
conditions (low to no O2 and no sulfide in the overlying waters) and associated authigenic
Mo deposits should dominate.
18
California Margin
Off the coast of California, the southward flowing California Current overlies the
more saline, oxygen-depleted, nutrient-rich California Undercurrent (Reid et al., 1958;
Murray et al., 1983). This region experiences northerly winds most of the year, leading to
offshore Ekman transport and associated upwelling. Upwelling is most intense in the spring;
southerly winds dominate in late summer/fall leading to relaxed (or no) upwelling (Lynn and
Simpson, 1987). The intense seasonal upwelling enhances biological productivity, and the
associated high productivity zone bordering the East Pacific (combined with circulation
patterns) results in a strong OMZ in the water column between depths of 200-1000 m (e.g.
Sverdrup and Allen, 1939). Today, the OMZ is deeper and less intense off central California
(~35oN) than off the Mexican margin (~23oN), reflecting ventilation in the north (van Geen
et al., 2003). This explains why laminated sediments (indicative of low bottom water
oxygen) are generally only preserved on the open margin within the OMZ at southern sites
(south of ~24oN; Keigwin, 1998).
The seven basins along the California margin investigated in this study are all part of
a larger complex known as the Borderland Basins region, described in detail by Emery
(1960). This region consists of approximately parallel belts of basins and banks (Figure 8).
All together, there are at least 18 known basins in this region, four of which are now
subareal. Basins with the flattest floors are generally those nearest shore, and the entire
region is characterized by a general deepening of the basins with distance offshore; basins
along each belt become deeper to the southeast. The topographic ridges surrounding the
basins limit the exchange of water between adjacent basins. Because circulation is restricted,
there are distinct inter-basinal differences in water chemistry that reflect both advection of
new water into the basin as well as reactions within the basin water and sediments.
19
Most lithogenic sediment is transported to the Borderland basins via rivers, eolian
deposition, and storm runoff. The outer basins are dominated by suspended particle
transport, while turbidites are only generally important for the innermost basins (Emery
1960). The nearshore sites have higher lithogenic inputs than those farther offshore; it has
been estimated that 40% of the total sediment input to the entire Borderland region can be
accounted for by deposition within just six nearshore basins (Schwalbach & Gorsline, 1985).
The three inner basins investigated in this study, Santa Barbara, Santa Monica, and
San Pedro, comprise a single “belt” ~30 km offshore (Figure 8). The Santa Barbara Basin is
580 m deep, with a western sill to the Pacific at ~450-475 m depth (Sholkovitz, 1973). The
Santa Monica Basin is 930 m deep, and is connected with the 890 m deep San Pedro Basin
by a sill at 740 m (Emery, 1960). All three inner basins have sill depths well within the
oxygen minimum zone, and these sites have the lowest measured bottom water oxygen
concentrations (<10 μM, Reimers, 1987; Berelson et al., 1987; Jahnke, 1990; Table 1).
Much of the deepest portions of these basins are characterized by laminated sediments
(Finney and Huh, 1989), suggesting that diffusion generally dominates solute exchange over
bioirrigation and bioturbation (Berelson et al., 1987). The inner basins are the most reducing
environments studied on the margin (Figure 9); sulfate reduction dominates organic matter
remineralization at these sites (e.g., Kaplan et al., 1963; Berelson et al., 1987; Jahnke, 1990).
The other four basins in this study (moving offshore) are San Clemente, Santa
Catalina, San Nicolas, and Tanner (Figure 8). The Santa Catalina and San Clemente basins
are part of the same “belt” ~80 km offshore (parallel to the coast), San Nicolas is one of
three basins in the next “belt” further west ~130 km offshore, and Tanner is one of four
basins ~160 km offshore comprising the belt furthest from land (Figure 8). Santa Catalina
(~1300 m deep), San Nicolas (~1800 m deep), and Tanner (~1500 m deep) basins all have
20
sill depths at or just below the base of the OMZ (1000-1200 m; Emery, 1960), and measured
bottom water concentrations are generally between 15-35 μM (Reimers, 1987; Berelson et
al., 1987). At these sites, organic matter is primarily oxidized by suboxic reactions (Mn and
Fe reduction; Figure 10; e.g., Berelson et al., 1987; Shaw, 1990). San Clemente is the
deepest Borderland basin site studied (~2100 m); its sill depth (~1800 m) is well below the
oxygen minimum and sediments are bathed by relatively oxygen rich waters (~60 μM O2;
Reimers, 1987). San Clemente represents the most oxic depositional environment studied in
this region; aerobic processes dominate organic matter decomposition (Bender et al., 1989)
and high concentrations of solid-phase Mn have been reported in surficial sediments
(McManus et al., 2006; Figure 10).
Inner Basins
Santa Barbara Basin has a higher marine productivity, sedimentation rate (0.4 cm/yr;
Reimers et al., 1990) and Corg rain rate compared to basins farther offshore (Crisp et al.,
1979). The sill depth (~475 m) is within the oxygen minimum zone, leading to nearly anoxic
bottom waters and the preservation of laminated sediments. These laminated sediments
represent annual varves, the product of periodic flushing of the basin as outside waters spill
over seasonally (Sholkovitz and Gieskes, 1971; Reimers et al., 1990).
Seasonal flushing of the basin generally accompanies spring upwelling, temporarily
delivering more oxygenated and nitrate-rich waters to the lower basin. This spillover is
usually marked by a lack of detectable sulfide in the uppermost surface sediments, though
conditions return to a more reducing state within three to four months with bottom water
oxygen < 2 μM and H2S[aq] detectable just below the sediment-water interface (Reimers et
al., 1990). Although the deep basin is periodically replenished with more oxygenated
21
seawater from outside the basin, several decades of research in Santa Barbara basin suggests
that it is the most reducing of the all the Borderland Basins studied herein (Figure 9).
High-resolution microelectrode work has shown the oxygen penetration depth in
deep basin sediments can range from 0-0.5 cm (Reimers et al., 1996). Pore water nitrate and
nitrite peaks occur within 0.25 cm of the sediment water interface, disappearing entirely by
~2 cm depth (Reimers et al., 1996; Figure 9). Small maxima in dissolved Mn at ~1 cm
suggest some Mn (IV) reduction in these sediments, but larger dissolved Fe peaks suggest Fe
reduction is a more prevalent mechanism for organic matter degradation in the most surficial
sediments (Reimers et al., 1996; Figure 9).
Sulfur isotopic data indicate that bacterial sulfate reduction is the dominant
mechanism of organic matter oxidation in Santa Barbara; one study suggests sulfate
reduction is responsible for nearly all of the observed carbon oxidation in deep basin
sediments (Kaplan et al., 1963). Maximum sulfate reduction occurs within the top 2-4 cm of
the sediment column, and decreases with depth (Reimers et al., 1996). Dissolved sulfide
concentrations up to 15 nM have been observed in deep basin waters (>400 m), confirming
that sulfate reduction is indeed an important process at work in the basin sediments
(Kuwabara et al., 1999). Iron reduction reactions and subsequent Fe-S formation appear to
regulate sulfide fluxes in the deep basin. Reimers et al. (1996) reported low ΣH2S[aq]
concentrations (<0.5 μM) until a depth of ~4 cm, the depth where reduced Fe was no longer
observed in pore water profiles.
Alkalinity and NH4+ enrichments (Figure 11) are higher in pore waters of the deep
basin than those of the slope (Sholkovitz, 1973). The higher alkalinity observed in deep
basin sediments (a product of sulfate reduction) leads to better preservation of carbonates in
the basin (10-12 wt.%) versus the slope (5-7 wt.%; Sholkovitz, 1973). Pore water sulfate
22
depletions (Figure 11) are larger in deep basin sediments than those of the slope, reflecting
the more reducing conditions at the basin floor (Sholkovitz, 1973). Pore water pH values
generally increase with depth in the sediments, though a minimum at ~3 cm was observed
that correlates with the observed depth of nitrate depletion and maximum sulfate reduction
(Reimers et al, 1996). Measured pore water pH values are higher than bottom water (7.5), a
feature attributed to the dominance of Fe reduction reactions consuming protons in the
uppermost surface sediments (Reimers et al., 1996). Iron reduction is also suggested by the
accumulation of excess phosphate in pore waters with depth (Reimers et al., 1996; Figure
11)
Santa Monica Basin is connected with the San Pedro Basin by a sill at 740 m
(Emery, 1960), and chemical profiles indicate similar diagenetic regimes in both basins
(Figure 9). A surface mixed layer is evident in Santa Monica sediments from the basin slope
underlying oxic bottom waters, but mixing is subdued in sediments of the deep basin below
the sill depth (Huh et al., 1987). Oxygen is <5 μM in the water column below the sill depth,
and the oxygen penetration depth is only ~0.2 cm in deep basin sediments (Berelson et al.,
1996); therefore oxygen does not appear to play a major role in the breakdown of organic
matter in the deep basin (Shaw et al., 1990). In fact, oxic mineralization of organic matter is
estimated to account for only ~10% of the total carbon decomposition in the Santa Monica
Basin (Jahnke, 1990).
Pore water profiles indicate that nitrate reduction is complete within the top ~1cm of
sediment, and Fe reduction begins just below this depth (Shaw et al., 1990; McManus et al.,
1997; Figure 9). Pore water phosphate profiles are also consistent with shallow depths of Fe
reduction (Figure 11); the observed phosphate flux from the sediments is in excess of that
predicated solely from organic matter decomposition (McManus et al., 1997). A maximum
23
in solid phase Fe (8-11 wt.%) in the uppermost ~1 cm of sediment is contained within a
yellow-brown surface oxidized layer (Finney and Huh, 1989). Below this depth, sediments
are green in color and have less solid-phase Fe (5-7 wt.%), suggesting Fe reduced below ~1
cm diffuses upwards and is reoxidized at the sediment-water interface (Finney and Huh,
1989; Shaw et al., 1990). Despite the Mn maxima observed in the Mn profile depicted in
Figure 9 (McManus et al., 1998), other published Mn profiles from Santa Monica Basin
show little evidence of Mn reduction at this site (Shaw et al., 1990; McManus et al., 1998)
suggesting that very little Mn is deposited or recycled in the deep basin, and that perhaps
Mn-oxides are recycled in the overlying suboxic water column (Shaw et al., 1990). Sulfate
reduction in the sediments of Santa Monica Basin is suggested by an observed decrease in
pore water sulfate with depth, as well as nearly linear increases in alkalinity and NH4+
beginning just below the sediment-water interface (Figure 11; Jahnke, 1990; McManus et al.,
1998).
San Pedro Basin is next to Los Angeles, and receives more detrital sediment than
basins farther offshore (Shwalbach and Gorsline, 1985). Bottom water oxygen contents in
the deep basin are typically between 3-5 μM, though periodic flushing of the basin by
colder, more oxygen- and nitrate-rich waters has been documented (~15 μM O2; Berelson,
1991). Flushing of the deep basin in San Pedro is more irregular than the annual flushing
events observed in the nearby Santa Barbara Basin (e.g., Reimers et al., 1990) and may be
related to El Niño climatic events (Berelson, 1991). Low bottom water oxygen
concentrations in the basin limit bioturbation, preserving laminated sediments in the deep
basin (Finney and Huh, 1989).
The oxygen budget for San Pedro Basin, as determined over an 11-year study period
(including times of flushing and stagnation), requires some oxygen consumption within the
24
water column (Berelson, 1991). Benthic flux studies indicate similar carbon oxidation rates
in San Pedro Basin sediments and in the adjacent Santa Monica Basin (1.8 and 1.7
mmol/m2d, respectively); though the estimated net sulfate reduction rate in San Pedro is
lower than Santa Monica (0.23 and 0.34 mmolS/m2d, respectively; Berelson et al., 1996).
This difference is possibly due to higher rates of nitrate reduction in San Pedro, as reflected
by higher benthic nitrate fluxes into the sediments of San Pedro than Santa Monica (-1.13
and -0.91 mmol/m2d, respectively; Berelson et al., 1996)
Presley and Kaplan (1968) reported significant sulfate depletion in the pore waters
of San Pedro sediments (as well as increasing ammonia concentrations with depth). These
observations suggest sulfate reduction is a dominant process for organic matter degradation
in this basin (Figure 11; Presley and Kaplan, 1968; Berelson et al., 1987; Leslie et al., 1990).
Leslie et al. (1990) calculate sulfate reduction rates of ~8 μmole S/cm2yr in the uppermost
sediments, decreasing exponentially with depth.
High Fe concentrations in pore waters from the top 0-50 cm of sediment also
suggest reducing conditions in San Pedro Basin (Figure 9; Presley and Kaplan, 1968; Leslie
et al., 1990; McManus et al., 1997). Bruland et al. (1974) report a sharp decrease in solidphase Fe concentrations at ~2 cm depth, suggesting Fe reduction may be an important
diagenetic process in the most surficial sediments. Pore water phosphate profiles also
suggest Fe reduction (McManus et al., 1997; Figure 11). Near-surface pore water Fe and
phosphate concentrations are lower in San Pedro than nearby Santa Monica Basin
(McManus et al., 1997; Figures 9 and 11); however, as in Santa Monica, an excess
phosphate flux was also observed from San Pedro sediments (McManus et al., 1997). It
appears that pyrite formation removes available reduced Fe from pore waters in the sediment
column <200 cm; below this depth, pore water H2S[aq] concentrations increase rapidly (to ~2
25
mM at 350 cm; Leslie et al., 1990). The lack of an observed minimum in solid-phase Mn
concentrations in these sediments (Bruland et al., 1974) suggests that, as proposed for Santa
Monica (Shaw et al., 1990), reduction of Mn-oxides is primarily taking place in the
overlying water column in this basin.
Outer Basins
Santa Catalina has a sill depth of ~980 m (Emery, 1960), and measured bottom
water oxygen is ~15 μM (Reimers, 1987). High resolution microelectrode profiles
demonstrate the oxygen penetration depth is only 0.3-0.5 cm; however, cm-scale reversals in
oxygen profiles below this depth suggest irrigation takes place through biogenic sedimentary
structures (Reimers, 1987). Leslie et al. (1990) report smaller pore water TCO2 and ammonia
gradients in Catalina surface sediments (0-30 cm) than those observed in San Pedro Basin,
presumably reflecting the effects of bioturbation (Figure 11). Despite this irrigation, sulfate
reduction rates of 4-5 μmole S/cm2yr are reported for surface sediments in Catalina basin
(Leslie et al., 1990). Sulfate reduction is indicated by an observed decrease in porewater
sulfate below the sediment-water interface, and sediments contain ΣH2S[aq] >4 mM below a
depth of ~2 m (Figure 11; Presley and Kaplan, 1968; Leslie et al., 1990; McManus et al.,
1998). Increases in pore water Fe and Mn concentrations below the sediment surface are
evident in pore water profiles (McManus et al., 1998); though maxima are less than those
observed in the more reducing inner basins (Figures 9 and 10). Similarities in Fe and
phosphate profiles suggest that Fe and P cycling are linked in this basin (McManus et al.,
1997; Figures 10 and 11).
San Nicolas Basin is silled just below the OMZ (~1140 m), and bottom water
oxygen is generally 20-35 μM (Berelson et al., 1987; Shaw, 1990), though sporadic periods
26
of “flushing” with more oxygen- and nitrate-rich waters have been observed (Berelson,
1991). The oxygen budget for this basin between times of flushing and stagnation can be
entirely explained by oxygen uptake into the sediments, suggesting that oxygen uptake in the
water column is not an important process at this site (Berelson, 1991). Evidence of
bioturbation (e.g. worm tubes) is apparent in some cores from this site (Berelson et al.,
1987); however, oxygen penetration depths as shallow as ~0.2 cm have been reported (Shaw,
1990). Pore water profiles show that nitrate is completely consumed by ~0.75 cm in San
Nicolas sediments (Shaw, 1990; Figure 10). Pore water profiles also suggest that Mn(IV)
reduction begins below ~0.5 cm; Fe release upon Fe reduction is evident below a depth of ~2
cm (Shaw, 1990; McManus et al., 1997; Figure 10). As in Santa Catalina, coincident
maxima in San Nicolas Fe and phosphate profiles suggest that Fe and P cycling are linked in
this basin (Berelson et al., 1987; McManus et al., 1997; Figures 10 and 11). Leslie et al.
(1990) report TCO2 and ammonia profiles with slight increases over uppermost 10 cm of the
sediment column, but large gradients are only evident below ~2 m depth. Sulfate profiles in
surface sediments are generally vertical, suggesting little net sulfate reduction in the
uppermost ~40 cm (Berelson et al., 1987; Shaw, 1990; Figure 11). Unlike the nearby
Catalina Basin, ΣH2S[aq] in San Nicolas does not exceed ~1 mM over the top ~4 m of the
sediment column, and lower sulfate reduction rates (~3 μmol S/cm2yr) have been reported
for surface sediments in San Nicolas (Leslie et al., 1990).
Tanner Basin has a sill depth of 1165 m and the bottom water oxygen concentration
is ~25 μM (Berelson et al., 1996). The sediment oxygen penetration depth is modeled to be
~0.4 cm (Berelson et al., 1996). Small maxima in Mn, Fe, and phosphate profiles are
reported in the uppermost 10 cm of the sediment column (Figures 10 and 11), but benthic Fe
fluxes are among the lowest reported for the Borderland Basins (McManus et al., 1997;
27
1998). Sulfate concentrations are generally constant over the uppermost 25 cm, suggesting
little or no net sulfate reduction in the uppermost sediments (McManus et al., 1998; Figure
11).
San Clemente is the deepest and most oxic of all basin sites studied, with a sill depth
well below the OMZ (~1800 m; Emery, 1960). Measured bottom water oxygen in San
Clemente Basin is ~60 μM (Reimers, 1987). Bioturbation has been reported in basin
sediments (Bender et al., 1989), and surface sediments contain high concentrations of solidphase Mn (~3 wt.%; McManus et al., 2006). Despite high bottom water oxygen
concentrations, high resolution microelectrode profiles demonstrate the oxygen penetration
depth is only ~0.5 cm in this basin (Reimers, 1987). Nitrate begins decreasing just below the
sediment-water interface, and there appears to be a small zone of net nitrate production (due
to O2 reduction and nitrification of NH4+) between 2-5 mm depth, followed by nitrate
reduction below (Bender et al., 1989). Pore water profiles show nitrate is completely
consumed at a depth between ~1.75 cm (Shaw, 1990; Figure 10) and ~4 cm (Bender et al.,
1989).
There is no dissolved Mn detectable in the pore waters of San Clemente until a depth
of 0.5 to 1cm, below which Mn(IV) reduction ensues and Mn(II) is released into the
interstitial waters (Bender et al., 1989; Shaw, 1990; Figure 10). Pore water profiles suggest
Fe release upon Fe reduction below a depth of ~3 cm (Shaw, 1990; McManus et al., 1997).
Phosphate concentrations generally increase with depth, but enrichments are lower than
those reported from other basin locations (McManus et al., 1997; Figure 11). Sulfate and
ammonia concentrations are generally constant, indicating sulfate reduction is not a
dominant process in the top ~10 cm of sediment (Shaw, 1990; McManus et al., 1998; Figure
28
11). However, Bender et al. (1989) observed an increase in pore water NH4+ concentrations
below ~5 cm depth, which they suggested reflects the onset of sulfate reduction.
Bender et al. (1989) used pore water profiles, benthic lander fluxes, and assumed
stoichiometries for the reactions of interest to calculate the relative importance of various
oxidants for degrading organic matter in San Clemente sediments. Neglecting Fe reduction,
these authors estimated that ~70% of organic carbon was degraded via O2. They estimated
that nitrate and sulfate reduction were responsible for 16% and 13% of the total organic
carbon degradation, respectively, with Mn-reduction representing <1% of the total organic
carbon oxidation in San Clemente Basin.
The representative sediment profiles depicted in Figures 9, 10, and 11 illustrate the
range of geochemical environments found in the Borderland Basins region. In general, the
nearshore sites are the most reducing, with bottom water oxygen concentrations < 10 µM.
Bottom water oxygen contents increase in basins with distance offshore (e.g. Berelson,
1996). The largest sulfate depletions and ammonia and phosphate enrichments are observed
in inner basin cores, reflecting the more reducing character of the near shore basins (e.g.
McManus et al., 1997, 1998; Figure 11). Previously reported Mo isotopic compositions from
Borderland basin sediments span the full range of Mo isotopic behavior observed in marine
environments (Siebert et al., 2006). Additional sites analyzed in this work, notably those
from San Nicolas and Santa Catalina Basins, were selected to further constrain marine Mo
geochemical and isotopic behavior under suboxic conditions. Given the wealth of published
geochemical data available for the sites investigated in this study, observed sediment Mo
isotopic compositions can be confidently interpreted in light of the geochemical conditions
specific to each depositional environment.
29
Mexico Margin
The region off the Mexican margin investigated in this study experiences the same
dominant currents and hydrography as the region investigated off of southern California.
Though productivity off Southern California is generally higher than off of southern Baja
California (van Geen et al., 2003), the same oxygen-deficient North Pacific Intermediate
Water dominates subsurface currents and establishes an OMZ between depths of 500 and
1000 m (Thunell, 1998). This low oxygen core extends >1500 km off the coast of Mexico
(Sansone et al., 2004).
Within the Gulf of California, the movement of water masses is complex. There are
alternating cores of flow into and out of the Gulf, with no apparent consistent seasonal or
spatial patterns in flow (Castro et al., 2006). High salinity Gulf of California Water is present
year round (Castro et al., 2006), but the subsurface waters are dominated by low-oxygen
Pacific Intermediate Water that flows into the Gulf between depths of ~500 and 1000 m
(Thunell, 1998). The anoxic bottom waters throughout this region limit bioturbation and
allow for the preservation of laminated sediments underlying the OMZ (e.g. Calvert, 1966).
The sediment cores investigated in this study are all subsampled from those
previously reported in Sansone et al. (2004) and Berelson et al. (2005). There are two
stations sampled off the west coast of Southern Baja: the Soledad Basin, and the open
margin of the continental slope (Magdalena). Four sites were cored within the Gulf of
California; two stations lie within depositional basins (Alfonso and La Paz), and two are
from the open margin on the eastern side of the Gulf (Carmen and Pescadero). Two
additional sites were cored off the western margin of mainland Mexico (south of Baja
Peninsula); the open margin off of Mazatlan, and San Blas further to the south (Figure 12).
30
Because the sediments on this margin are bathed in low-oxygen waters (e.g. Thunell,
1998), we anticipate limited decomposition of organic carbon via aerobic processes.
Evidence of denitrification and Mn reduction has been reported within the water column off
mainland Mexico (e.g. Nameroff et al., 2002; Hartnett and Devol, 2003), suggesting
diagenesis within the sediments will likely be dominated by reactions associated with Fe and
S cycling. In addition, reported methane production at these sites also confirms reducing
diagenetic conditions (Sansone et al., 2004).
Sansone et al. (2004) focused on the processes controlling methane fluxes in the
water column and sediments of this area. Berelson et al. (2005) sought to better understand
anaerobic silica and carbon diagenesis throughout the region. While this study builds upon
their previous findings, far less data is available in the literature for these locations than, for
example, the Borderland Basin sites off of southern California previously discussed. As
such, a detailed discussion of the anticipated diagenetic regimes within the sediments is not
possible for many of these sites, and the sediment analyses of this study (both major element
and trace metal work) will be required to fully characterize the nature of these sedimentary
environments.
W. Coast of Southern Baja
Soledad Basin (sometimes called San Lázaro Basin) lies ~45 km west of the Baja
California coast (Figure 12). Unlike the near shore basins off of southern California, the site
is not measurably affected by human activity (Bruland et al., 1974). The basin has a very flat
bottom with a maximum depth of 545 m and an effective sill depth of ~290 m (van Geen et
al., 2003). Water column profiles within the basin show a reduction from ~8 to <5 μM O2
with depth, though a slight increase observed in the water column O2 profile near the bottom
31
may indicate spillover to the deep basin from waters outside (van Geen et al., 2003). A
multi-year sediment trap study by Silverberg et al. (2004) found that the deep waters in the
basin remained almost uniform in temperature and salinity, and this was taken to reflect the
constant presence of North Pacific Intermediate Water that floods the bottom of the basin.
Water column nitrate behaves conservatively with depth inside the basin (van Geen et al.,
2003). The sediments of the deep basin are laminated (where bottom water is <5 μM O2),
and the sedimentation rate appears to be steady at ~0.1 to 0.3 cm/yr (Bruland et al., 1974;
van Geen et al., 2003; Berelson et al., 2005). The core analyzed in this study was collected
from a water depth of 542 m in the deepest part of the basin, where bottom water oxygen
was 0 μM and sediment laminations were preserved (Berelson et al., 2005).
Previous work in this basin revealed organic carbon contents fairly constant with
depth (~6 wt.%), while carbonate decreased from 20 wt.% to 10 wt.% below ~20 cm
(Bruland et al., 1974). Sediment Mn and Fe concentrations were low and fairly constant with
depth, (~2 wt.% Fe, ~150 ppm Mn; Bruland et al., 1974) suggesting they do not play a
significant role in carbon oxidation at this site. Sansone et al. (2004) reported the largest
surface sulfate gradient (0.24 mmol/L/cm) in Soledad sediments of all Mexican margin sites
analyzed, with sulfate entirely depleted within the uppermost 80 cm of the sediment column
(Figure 13).
The Magdalena margin sediment core analyzed in this work was collected from a
water depth of 713 m on the open margin of the continental slope (Figure 12). Bottom water
oxygen at this depth was 1.3 μM, and sediments from this core were bioturbated in the top 23cm (Berelson et al., 2005). A steady sedimentation rate of ~0.03 cm/yr was estimated by
van Geen et al. (2003) for a similar site on this margin at ~700 m water depth. Sansone et al.
32
(2004) noted a high surface sulfate gradient (0.2 mmol/L/cm; comparable to Soledad) for
this site, with sulfate depleted by 100 cm depth in the sediment column.
Gulf of California
Alfonso Basin is located offshore the city of La Paz in La Paz Bay, on the
southeastern coast of Baja within the Gulf of California (Figure 12). The southern portion of
La Paz bay is shallow, but deepens towards the north and descends abruptly below 200 m to
form Alfonso Basin, with a maximum depth of 420 m (Cruz-Orosco et al., 1989, 1996 in
Silverberg et al., 2006) and a sill depth of ~325 m (Sansone et al., 2004). Bottom water
oxygen at the study site has been measured as 0.5 μM, and the sediments are laminated with
turbidite layers evident ~250 cm depth in the sediment column (Berelson et al., 2005).
Surface sediments from within the deep basin are estimated to contain ~15 wt.% carbonate
and ~30 wt.% organic matter (Silverberg et al., 2006). A sediment accumulation rate of 0.04
cm/yr estimated from sediment traps (Silverberg et al., 2006) agrees well with reported
210
Pb-determined sedimentation rates of 0.05, 0.04, and 0.06 cm/yr (Nava-Sanchez (1997);
Perez-Cruz (2000); Rodriguez-Castaneda (2001) in Silverberg et al., 2006) and 0.05 cm/yr
(Gonzalez-Yajimovich, 2004 in Berelson et al., 2005). Sansone et al. (2004) reported a
surface sulfate gradient of 0.14 mmol/L/cm at this site, and observed that pore water sulfate
was depleted by 170 cm depth.
La Paz Basin lies just east of Alfonso Basin, at the northeast edge of La Paz Bay
(Figure 12). This site is much deeper (~900 m), with a measured bottom water oxygen
concentration of 0 μM. (Berelson et al., 2005)
The open Carmen margin lies on the eastern side of the Gulf of California. The two
cores analyzed from the Carmen margin were collected at 575 and 800 m water depth on the
33
continental slope, where measured bottom water oxygen concentrations were 0.2 and 1 μM,
respectively (Berelson et al., 2005; Figure 12). Sediments from both cores were laminated,
with no signs of bioturbation (Berelson et al., 2005). Baba et al. (1991) estimated that 8090% of total sedimentation along the eastern margin of the central and southern Gulf was of
terrigenous origin. The site at 575 m depth was found to have a low surface sulfate gradient
(0.06 mmol/L/cm), and sulfate concentrations were greater than 3 mmol/L throughout the
upper 4 m of the sediment column (Sansone et al., 2004).
The Pescadero slope, also on the eastern margin of the Gulf of California, was cored
at 506 and 600 m water depth (Figure 12). Measured bottom water oxygen was 0 μM for
both sites, and both cores had laminated sediments (Berelson et al., 2005). At the 600 m site,
a surface sulfate gradient of 0.17 mmol/L/cm was reported, and all sulfate was depleted by
180 cm depth (Sansone et al., 2004). Berelson et al. (2005) estimated that 40% of TCO2 flux
across the sediment-water interface was from a reaction zone much deeper within the
sediment column.
W. Coast of Mainland Mexico
The Mazatlan margin is perhaps the best studied of all the Mexican margin sites
investigated in this work, with several previous studies reporting water column and sediment
data from transects across the margin. Here, the Mexican margin has a narrow continental
shelf with a shelf break at ~200 m and a gradual deepening of the continental slope to ~3000
m (Hartnett and Devol, 2003). The oxygen minimum zone intersects the continental slope
off of Mazatlan between ~100 and 1000 m water depth; water column oxygen is <5 μM
between ~150 and 800 m water depth, and nitrate profiles indicate nitrate depletion in the
water column (Nameroff et al., 2002; Hartnett and Devol, 2003). A water column Mn
34
maximum (~8 nmol/kg at 400 m) suggests suboxic conditions in the water column
(Nameroff et al., 2002).
Sediments from below the core of the OMZ are not bioturbated, preserving sediment
laminations (Hartnett and Devol, 2003). There is an organic carbon concentration maximum
on the mid-slope off Mazatlan in sediments just deeper than the core of the OMZ. This
maximum is thought to be caused by winnowing by currents on the outer shelf, preferential
accumulation of organic matter in fine-grained sediments, and the offshore decrease in
primary productivity (and therefore less settling of organic matter) (Ganeshram et al., 1999).
At depths below 250 m, sediment organic carbon contents are less than those observed in
sinking particles, suggesting remineralization of organic matter at the sediment-water
interface (Nameroff et al., 2002).
The core analyzed in this study was collected from 442 m water depth (near site
NH15P of Ganeshram et al., 1999; Figure 12), containing laminated sediments and notable
turbidite layers between 250 and 400 cm (Berelson et al., 2005). Measured bottom water
oxygen at this site was 0.2 μM (Berelson et al., 2005). Similar to the sediments of the
Carmen margin, sediments of this site were found to have a low surface sulfate gradient
(0.06 mmol/L/cm; Sansone et al., 2004; Figure 13). A sedimentation rate of 0.015 cm/yr has
been reported for a similar site on this margin (Ganeshram et al., 1999).
Hartnett and Devol (2003) analyzed pore fluids and sediments from several sites
transecting the Mazatlan margin. At the 420 m site (nearest our core depth) bottom water
oxygen was measured at 0 μM and no oxygen was detectable in pore waters. These authors
report benthic nitrate fluxes of ~-1.1 mmol/m2d and sulfate reduction rates of ~0.9
mmol/m2d from chamber deployments at this depth. Hartnett and Devol (2003) estimate that
denitrification is regionally responsible for ~40% of the total carbon oxidation in Mazatlan
35
margin sediments, with sulfate reduction representing ~80% of carbon oxidation in shallow
sites and less than 20% for deeper sites (calculations ignore Mn and Fe reduction
contributions because they are assumed to be small). Sulfate reduction is also suggested by
ammonia enrichments observed in pore water profiles at this site (Hartnett and Devol, 2003;
Figure 13).
The San Blas station lies south and landward of Tres Marias Island chain, and is the
southernmost Mexican margin site investigated in this study. This site was cored within a
basin 430 m deep that is silled at ~300 m (Berelson et al., 2005; Figure 12). Reported bottom
water oxygen at this site is 0 μM, and sediment laminations are present. Sansone et al.
(2004) report a surface sulfate gradient of 0.10 mmol/L/cm for the San Blas site, similar to
that observed on other open margin sites (Carmen and Pescadero).
Sansone et al. (2004) observed the highest surface sulfate gradients and largest
methane fluxes on the Pacific side of Baja California (e.g. Soledad, Figure 13); inside the
Gulf of California and off the Mexican mainland, sulfate gradients and methane fluxes were
generally much lower and more variable (e.g. Mazatlan, Figure 13). Despite these
differences, all the Mexico margin sites are presumed to contain relatively anoxic sediments;
bottom water oxygen concentrations are low (<5 μM) and laminated sediments are present at
all but the Magdalena site. Published sulfate and ammonia gradients from sites on the
Mexican margin are similar to those from Peru and the most reducing inner basins of the
California margin (Figure 13), further suggesting anoxic sedimentary conditions at the
Mexico sites. As discussed in Chapter 2, initial sediment Mo isotope measurements from the
Soledad, Mazatlan, and San Blas sites suggest a unique Mo isotopic signature for anoxic Mo
enrichments (Poulson et al., 2006). Observations from several additional sites on the
Mexican margin will further constrain Mo behavior under reducing conditions.
36
SUMMARY
Although there are significant similarities among all the sites discussed here, there
are important differences in sediment geochemistry that allow us to evaluate Mo behavior
across a broad range of diagenetic regimes. The MANOP sites represent the most oxic
sedimentary conditions of this study; electron transport is dominated by oxygen and nitrate
consumption and manganese reduction, and these sites have a thick Mn-rich layer which
controls Mo behavior (e.g., Klinkhammer, 1980). In the sediments of the outer California
borderlands basins these diagenetic regimes become compressed spatially, and the onset of
subsequent diagenetic reactions of iron and sulfate reduction. Within the near-shore basins of
the California margin and throughout the Mexico margin, iron and sulfate reduction become
increasingly important; free sulfide is present in the uppermost sediment column in Santa
Barbara Basin (e.g. Reimers et al., 1990). It is likely that some of the Mexico margin sites
also have sulfide present in the upper sediment column; however, the limited data preclude a
more exhaustive characterization of the sulfur cycling in these settings. Along the Peru
margin, it is well know that the sediments are sulfidic and organic rich (e.g. Reimers and
Suess, 1983; Froelich et al., 1988); these sediments represent an end-member case for
sedimentary reducing conditions in this study. The sites selected for this study span a wide
range of geochemical conditions, and observations from these modern sedimentary
environments allow us to fully characterize Mo geochemistry and associated isotopic
fractionations in modern marine systems.
37
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43
Figure 1. Major Mo sources to modern marine sediments: 1) Lithogenic Mo terrigenous material incorporated into bulk sediment, 2) Biogenic Mo – sorbed to or
incorporated into organic material, 3) Authigenic Mo - directly precipitated as a solid phase
within the sediments (under both oxic and anoxic conditions).
Figure 2. Published marine Mo isotope values and fractionation factors from Barling et
al., 2001, McManus et al., 2002, Siebert et al., 2003 and 2006, Poulson et al., 2006, and
Wasylenki et al., 2007.
Figure 3. Schematic summarizing Mo behavior under various diagenetic regimes. At left, a general sequence of
diagenetic processes and associated pore water profiles. At right, schematic depicting Mo geochemical behavior under
various diagenetic conditions. Inset: Mo(aq) speciation diagram adapted from Helz et al. (2004).
44
45
Figure 4. Map of study areas showing approximate locations of all sites investigated.
Figure 5. Map of Peru margin and MANOP sites.
Figure 6. Pore water profiles and generalized diagenetic regimes for MANOP sites M and H. All data from
Klinkhammer (1980).
46
47
Figure 7. Pore water profiles and generalized diagenetic regimes for Peru margin. Data
from Froelich et al., 1988 (near 12oS, 183m water depth).
Figure 8. Map of California margin Borderland Basin sites.
Figure 9. Pore water profiles and generalized diagenetic regimes for the three inner basins of the California
margin. Iron and manganese data for Santa Barbara from McManus (personal comm.; core analyzed in this study, MC22);
nitrate data from Reimers et al. (1996) (core BC12). In Santa Barbara basin, depth of red “anoxic” zone (as in Figure 3)
estimated from Fe profile and depth of detectable sulfide (12cm) reported in Sholkovitz (1973). Data for Santa Monica and
San Pedro basins from cores analyzed in this study; Fe and Mn data from McManus et al. (1998); nitrate data from
McManus (personal comm.).
48
Figure 10. Pore water profiles and generalized diagenetic regimes for the four outer basins of the California margin.
Iron and manganese data for all sites from cores analyzed in this study (McManus et al., 1997; 1998; personal comm..); nitrate
data from Santa Catalina, Tanner, and San Clemente from McManus (personal comm.); San Nicolas nitrate data from Shaw et
al. (1990). Depth of blue “oxic” zone (as in Figure 3) estimated from reported oxygen penetration depths (Berelson et al.,
1996).
49
Figure 11. Pore water profiles of sulfate, ammonia, and phosphate for all Borderland basin sites investigated in this
study. Sulfate data for Santa Barbara basin from Sholkovitz (1973); sulfate data for San Nicolas from Berelson et al., (1987);
sulfate data for all other basins from McManus et al. (1998). Ammonia data for San Pedro and San Clemente basins reported
in McManus et al. (1997); data for all other sites from McManus (personal comm.). Phosphate data for Santa Barbara basin
from Reimers et al. (1996); phosphate data for San Nicolas from Berelson et al., (1987); phosphate data for all other basins
from McManus et al. (1997).
50
51
Figure 12. Map of Mexico margin sites.
Figure 13. Pore water profiles of ammonia and sulfate for two Mexican margin sites,
the Peru margin, and all Borderland basin sites investigated in this study. Sulfate data
for Soledad and Mazatlan sites from Berelson et al. (2005). Sulfate data for Peru margin
from McManus (personal comm.) Sulfate data for Santa Barbara basin from Sholkovitz
(1973); sulfate data for San Nicolas from Berelson et al., (1987); sulfate data for all other
California basins from McManus et al. (1998). Ammonia data for Mazatlan site reported in
Hartnett and Devol (2003); data for Peru margin from McManus (personal comm.).
Ammonia data for San Pedro and San Clemente basins reported in McManus et al. (1997);
data for all other sites from McManus (personal comm.).
52
Table 1. Study Site Characteristics. Depths of oxygen penetration, Mn reduction, and
detectable sulfide are estimated from pore water data, water column measurements, sediment
data, and geochemical modeling (references listed). OLW is overlying water; SWI is
sediment-water interface. Data from [1] Klinkhammer (1980); [2] Lyle et al. (1984); [3]
Kadko (1981); [4] Bender and Heggie (1984); [5] Suess et al. (1986); [6] Froelich et al.
(1988); [7] Fossing (1990); [8] McManus et al. (2006); [9] Emery (1960); [10] Reimers et al.
(1990); [11] Reimers et al. (1996); [12] Shaw et al. (1990); [13] McManus et al. (1998); [14]
Berelson et al. (1996); [15] Bruland et al. (1974); [16] Berelson et al. (1991); [17] Leslie et
al. (1990); [18] Reimers et al. (1987); [19] Berelson et al. (1987); [20] McManus et al.
(1997); [21] Bender et al. (1989); [22] van Geen et al. (2003); [23] Berelson et al. (2005);
[24] Cruz-Orosco et al. (1989); [25] Sansone et al. (2004); [26] Ganeshram et al. (1999);
[27] Nameroff et al. (2002); [28] Hartnett and Devol (2003).
53
54
AUTHIGENIC MOLYBDENUM ISOTOPE SIGNATURES IN MARINE
SEDIMENTS
Rebecca L. Poulson, Christopher Siebert, James McManus, and William M. Berelson
Geology
Geological Society of America
3300 Penrose Place
PO BOX 9140
Boulder, CO 80301-9140
Volume 34, No. 8, doi: 10.1130/G22485.1
55
ABSTRACT
We present new Mo isotope data from the Mexican continental margin that, in
conjunction with previous data, allow us to propose a mechanistic description of the Mo
isotope system in marine sediments. We hypothesize that there are unique environmentallydependent Mo isotope signatures recorded in marine sediments that reflect the mechanisms
responsible for authigenic Mo accumulation. Open-ocean anoxic sites, defined as having
dissolved oxygen and sulfide concentrations near zero in the overlying water, exhibit a
δ98/95Mo isotope signature of +1.6‰. We believe this value reflects Mo sulfide formation
via diagenetic processes within sediments. Quantitative formation of Mo sulfide within the
sulfidic water column of euxinic environments results in sediment isotope values similar to
the modern seawater value (+2.3‰), as typified by samples from the highly-sulfidic Black
Sea. In contrast to these reducing settings, manganese oxide-rich sediments have measured
Mo isotope values that are more negative (relative to seawater) than any other sediment
samples analyzed to date, similar to Fe-Mn crusts (~-0.7 ‰). Most measured Mo isotope
compositions of marine sediments from open ocean settings appear to reflect a mixture of
lithogenic Mo (0.0‰) and the Mo signature of a specific authigenic Mo accumulation
mechanism. We therefore suggest that Mo isotopes may record unique signatures that reflect
the dominant chemical mechanism for Mo sequestration into sediments.
INTRODUCTION
Molybdenum enrichments in reducing marine environments have been used in
paleochemical reconstructions (e.g., Crusius et al., 1996), and the Mo isotope system has
56
recently been utilized to the same end (Siebert et al., 2003; Arnold et al., 2004; Anbar,
2004). Studies have demonstrated significant natural variations in modern sediment Mo
isotope compositions (Barling et al., 2001; Siebert et al., 2003; Siebert et al., 2006), but
interpretation of these data is equivocal because not all mechanisms generating sediment
isotope signatures have been constrained. We present new Mo isotope data from several
marine anoxic sediments that, in conjunction with previous data, allow us to propose a
possible mechanistic description of the Mo isotope system in marine sediments.
There are essentially three major reservoirs of Mo in the marine environment:
seawater, lithogenic materials, and authigenically precipitated Mo. In modern seawater, Mo
exists primarily as the soluble molybdate complex (MoO42-). It has a conservative
distribution with a concentration of ~100 nM, and a residence time of ~800 kyr (Morford
and Emerson, 1999, and references therein). Although data remain limited (n = 6), isotopic
analyses of modern seawater indicate a uniform Mo isotope composition of δ98MoSW =
+2.3‰ (Fig. 1) (Siebert et al., 2003; Barling et al., 2001), which is consistent with seawater
being a well-mixed reservoir (Anbar, 2004). Data are presented in delta notation as δ98Mo =
((98/95MoSAMPLE - 98/95MoSTANDARD) – 1) × 1000).
Analyses of terrigenous materials (e.g., granites, clastic sediments) yield an average
Mo isotope composition of ~0.0‰ (n = 13) (Siebert et al., 2003), and we take this value to
be representative of the lithogenic background in marine sediments. Lithogenic Mo can be
an important component of the total Mo measured in a sediment sample, thus measured
sediment Mo isotope compositions require correction for the lithogenic Mo contribution.
One of the potential strengths of Mo isotopes lies in the observation that, to date, only
authigenic accumulations of Mo in marine sediments show significant isotopic variability,
and it is this strength that we wish to exploit.
57
AUTHIGENIC MOLYBDENUM
Molybdenum is a trace metal that forms authigenic deposits under both oxic and
reducing conditions. In oxic sediments, where aerobic respiration is the primary pathway for
organic material decomposition, Mo is sequestered through adsorption onto Mnoxyhydroxides (Bertine and Turekian, 1973; Calvert and Pedersen, 1993). Authigenic Mo
accumulated under oxic conditions has measured Mo isotope values more negative (relative
to seawater) than any other samples analyzed to date (Fig. 1). Analyses of Fe-Mn crust
surfaces produce an average (n = 6) value of δ98Mo = 0.7 ± 0.1‰ (Barling et al., 2001;
Siebert et al., 2003), which corresponds to a ~3.0‰ fractionation between modern seawater
and the adsorbed Mo species (Fig. 1). This fractionation is in good agreement with
experimental determinations of isotope fractionation during Mo-adsorption (2.7‰) (Barling
and Anbar, 2004). The exact speciation of adsorbed Mo is unknown, but recent quantum
mechanical calculations consistent with natural observed fractionations suggest MoO3 is a
potential candidate (Tossell, 2005).
Under reducing conditions, where anaerobic processes dominate electron transport,
Mo is sequestered into sediments via Mo sulfide formation. Pore fluid studies of Zheng et al.
(2000) argue for two separate thresholds of Mo sulfide formation; coprecipitation of Mo-FeS phases at dissolved sulfide concentrations of ~0.1 μM, and Mo-S precipitation without Fe
at sulfide concentrations of ~100 μM or more. Authigenic precipitation of Mo at ~0.1μM
sulfide likely corresponds to the initial formation of thiomolybdate intermediate complexes
(MoOxS4-x2-), which can be scavenged by sulfidized organic and Fe-S phases (Helz et al.,
1996).
58
At high dissolved sulfide concentrations a “geochemical switch” is proposed where
the dominant dissolved Mo phase transitions from soluble molybdate to less soluble
tetrathiomolybdate (MoS42-) (Helz et al., 1996). This presumably corresponds to the ~100
μM sulfide threshold proposed by Zheng et al. (2000), where Mo depletion in pore waters
was observed in the absence of Fe-S precipitation. It is assumed that in euxinic settings with
high dissolved sulfide concentrations, the “geochemical switch” threshold will be met, and
the dominant dissolved species may be MoS42- rather than MoO42- (Helz et al., 1996).
Previous work has proposed that because the conversion of MoO42- to MoS42- in euxinic
waters is quantitative, any transient fractionation between species will not be recorded in the
underlying sediments; as evidenced by sediments from the euxinic Black Sea with Mo
isotope compositions analytically indistinguishable from the modern seawater value (δ98Mo
= +2.4‰) (Barling et al., 2001; Arnold et al., 2004) (Fig. 1). Though restricted euxinic
basins are rare in the modern ocean, they were more widespread on the ancient Earth, and
represent an important Mo isotopic end-member composition as the most reducing
authigenic Mo deposits in the marine environment (e.g., Arnold et al., 2004).
A fractionation between MoO42- and Mo sulfide is suggested by the findings of
McManus et al. (2002), who calculate a 0.7‰ fractionation as dissolved Mo is removed into
sediments in Santa Monica Basin. We propose that this isotopic signature would most likely
be preserved in open-ocean anoxic sediments (defined here as those areas having dissolved
oxygen and sulfide concentrations near zero in the overlying bottom water) where
thiomolybdates (MoOxS4-x2-) are formed through diagenetic processes within the sediments.
In this study we have analyzed Mo isotope compositions from a suite of modern anoxic
sediments to determine if such a fractionation exists.
59
RESULTS
We selected three sites on the Mexican continental margin for Mo isotope analysis.
Each site has low bottom water oxygen concentrations (Table 1), and presumably no
dissolved sulfide present. The Soledad and San Blas sites are located within depositional
basins, while the Mazatlan site is from the open continental margin. The Soledad basin is
545 m deep with an approximate sill depth of 250 m (van Geen et al., 2003; Silverberg et al.
2004). The San Blas basin is 430 m deep, with an approximate sill depth of 300 m (Berelson
et al., 2005). The Mazatlan station is at a depth of ~440 m, and is located near the sites
previously discussed in Ganeshram et al. (1999) and Hartnett and Devol (2003).
The two basin sites have higher sedimentation rates than the open margin Mazatlan
site, yielding authigenic Mo accumulation rates of 2.19 and 5.29 nmol/cm2/yr compared to
1.28 nmol/cm2/yr on the open margin (Table 1). At the Soledad site, it has been shown that a
significant portion of organic carbon degradation can be attributed to sulfate reduction
(Berelson et al. 2005), and it is reasonable to assume that sulfate reduction and
methanogenesis are dominant processes regulating electron transport at all three sites in this
study.
Mo isotope analyses of sediments from all three anoxic sites demonstrate a
strikingly invariant average Mo isotope composition of δ98Mo = +1.6 ± 0.1‰ (1SD, n = 29)
(Table 1, Fig. 2). A mathematical correction was applied to all Mo isotope compositions post
analysis, to account for lithogenic Mo (Table 1). This correction was applied based on the
mass balance assumption that the total Mo isotope composition measured is a mixture of the
lithogenic Mo (0.0‰) present and the authigenic Mo signature: δ98MoSEDSXSEDS =
δ98MoLITHXLITH + δ98MoAUTHXAUTH. For all sites in this study, the lithogenic Mo component
60
was ≤8% of the total Mo concentration, resulting in a very small isotopic correction (Table
1). The data from these anoxic sites indicate Mo isotopes record a fractionation of 0.7‰
between seawater and Mo sulfides diagenetically produced within the sediments (Fig. 1),
which is consistent with the Mo isotope fractionation calculated by McManus et al. (2002).
DISCUSSION
The anoxic Mo isotope signature determined in this study, when combined with
existing sediment Mo isotope data, allows us to construct a model for the behavior of Mo
isotopes in authigenic marine deposits (Fig. 3). This model proposes that there are unique
environmentally-dependent Mo isotope signatures recorded in marine sediments, each
reflecting the primary mechanisms responsible for authigenic Mo accumulation. Adsorption
of an oxic Mo phase (presumably MoO3) in Mn-rich sediments produces an adsorbed Mo
isotope signature more negative than any other sediments measured to date, being as much
as 3‰ lighter than the modern seawater value (e.g., Barling et al., 2001):
δ98Mo(MoO42-aq) - δ98Mo(MoO3s) ≈ 3.0‰ (1)
We further propose that transformation of pore water MoO42- to MoO3S2- and other
subsequent thiomolybdates (such as MoS42-) within sediments via diagenetically-produced
sulfide results in an anoxic authigenic Mo sediment isotope composition that is consistently
0.7‰ lighter than overlying water:
δ98Mo(MoO42-aq) - δ98Mo(MoOxS4-x2-) = 0.7‰ (2)
In euxinic environments, syngenetic formation of dissolved MoS42- within the
sulfidic water column is presumably quantitative, thus the fractionation between species
(equation 2) is not observed in the underlying sediments.
61
Recognition of distinct environmentally-dependent Mo isotopic signatures and the
importance of a lithogenic Mo correction allows for a refined interpretation of previously
measured Mo isotope data from sediments along the California margin (Siebert et al., 2006).
After application of the lithogenic correction, all but one value can be resolved to either the
oxic or anoxic signatures (Table 1). The exception is data from Tanner Basin, where the
corrected Mo isotope value is δ98Mo = +0.5‰. We suggest that this value indicates either a
mix of the oxic and anoxic Mo sources at this site, as proposed in Siebert et al. (2006), or
another unidentified process that is dominating Mo accumulation. The corrected Mo isotope
compositions of sediments from Santa Monica and San Pedro basins are consistent with the
anoxic signature observed in the Mexican margin sediments (Table 1), indicating that
sulfidization of Mo is the dominant mechanism of authigenic Mo accumulation at these sites.
In contrast, the corrected Mo isotope compositions of Mn-rich sediments from San Clemente
Basin are consistent with the oxic Mo isotope signature as measured in Fe-Mn crusts
(Barling et al., 2001; Siebert et al., 2003) (Table 1), indicating that Mo adsorption onto Mnoxyhydroxides is the dominant source for the accumulating Mo (Siebert et al., 2006). These
observations lead us to conclude that for most marine sediments measured to date, the
average recorded Mo isotope signature is indicative of the dominant authigenic Mo
accumulation mechanism.
Molybdenum isotope data from Cariaco Basin sediments (δ98Mo = +1.8‰ (Arnold
et al., 2004)) and shallow margin sediments of the Black Sea (δ98Mo = +1.7‰ (Nägler et al.,
2005)) (Table 1) suggest that sediment Mo isotope compositions in restricted basin
environments are heavier than the open-ocean anoxic value. Although these reported values
are consistent with data from our anoxic sites, they have not been corrected for lithogenic
Mo and their authigenic Mo isotopic compositions are likely to be heavier. We propose that
62
heavy Mo isotope compositions for sediments in restricted basin environments may reflect
Mo limitation (Algeo and Lyons, in press), such that inadequate resupply of dissolved
MoO42- drives aqueous Mo isotope compositions to heavier values as Mo sulfides are
precipitated (dashed line in Figure 3). We believe that the relative 0.7‰ fractionation
between MoO42- and authigenic Mo sulfide remains constant, and suggest that sediment Mo
isotope values heavier than those observed at open-ocean sites reflect a shift in the isotopic
composition of the dissolved MoO42- pool. The fractionation between these Mo species is no
longer observed in sediments when full euxinia is reached (as in the deep Black Sea) because
the dissolved Mo pool has been quantitatively converted from MoO42- to MoS42-.
Previous application of the Mo isotope system in paleochemical reconstruction has
relied on a simple mass balance between the oxic and reducing Mo sinks (Arnold et al.,
2004). The recognition of a unique Mo isotope signature in open-ocean anoxic sediments, in
conjunction with the potential impact of Mo limitation on Mo isotopic behavior in restricted
environments, obfuscates the use of Mo isotopes to constrain paleoredox conditions. Further
investigations to elucidate the affect of Mo limitation on sediment Mo isotope compositions
are warranted before the Mo isotope system can be effectively employed as a paleochemical
proxy.
CONCLUSIONS
We suggest that there are two primary Mo isotope fractionations recorded in marine
sediments: an oxic fractionation (~-3.0‰) reflecting Mo adsorption to Mn-oxides, and an
anoxic fractionation (~-0.7‰) indicative of thiomolybdate (MoOxS4-x2-) formation. As
previously proposed, in environments where sulfide concentrations exceed 100 μM there
63
appears to be quantitative conversion of aqueous MoO42- to the MoS42- phase, such that no
observable fractionation is recorded in the sediments. We further note that authigenic
signatures can be obscured by the lithogenic Mo contribution, and measured sediment Mo
isotope compositions require correction for this affect.
ACKNOWLEDGEMENTS
Ariel Anbar, Tim Lyons, John Crusius, and Jane Barling provided constructive
criticism on an early version of this manuscript. This research was supported by NSF grant
OCE-0219651 to JM and NSF grants OCE-0002250 and OCE-0129555 to WMB.
64
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Siebert, C., Nägler, T. F., and Kramers J.D. (2001) Determination of molybdenum isotope
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Figure 1. Measured Mo isotope compositions of various marine sediments and the
presumed dominant electron transport processes for each environment. Oxic sediment Mo
isotope data are from Fe-Mn crust surfaces (Barling et al. (2001); Siebert et al., 2003).
Suboxic sediment Mo isotope data are average site compositions from California margin
sites of Siebert et al. (2006). Anoxic sediment Mo isotope data are average site compositions
from three Mexican margin sites of this study. Euxinic sediment Mo isotope data are deep
Black Sea sediments from Barling et al. (2001) and Arnold et al. (2004). All data are shown
without lithogenic Mo correction.
67
Figure 2. All δ98Mo data (without lithogenic correction) from down-core profiles of the
three anoxic Mexican margin sites in this study; errors shown are 2-SD. The average Mo
isotope composition of all measured samples is δ98Mo = +1.6 ± 0.1‰ (1-SD, n = 29) (Table
1).
68
Figure 3. Schematic of the authigenic Mo isotope system in marine sediments depicting
the measured oxic (δ98Mo = 0.7‰), anoxic (δ98Mo = +1.6‰), and euxinic (δ98Mo = +2.3‰)
Mo isotope signatures and proposed Mo chemical behavior. Measured Mo isotope
compositions of marine sediments from open ocean settings appear to reflect a mixture of
lithogenic Mo (δ98Mo = 0.0‰) and the authigenic Mo signature of either the oxic or anoxic
Mo accumulation mechanisms. Measured Mo isotope values for restricted basin sediments
suggest a change in the isotopic composition of dissolved Mo (dashed line at top) to heavier
values as Mo is removed in the presence of dissolved sulfide.
69
Table 1. Mo Isotope Compositions of Various Marine Depositional Environments. All
samples processed and analyzed for Mo and Al concentrations and Mo isotope compositions
as described in Siebert et al. (2006). δ98Mo values are δ98Mo = ((98/95MoSAMPLE 98/95
MoSTANDARD) - 1) × 1000), standard used is JMC-ICP Mo standard solution (lot
602332B). Errors on reported δ98Mo values are 1-SD for all samples averaged; number of
samples averaged for suboxic and anoxic data is the number of samples analyzed from a
detailed down core profile averaged to determine the total site δ98Mo value reported (Fig. 2).
For all other data, number of samples averaged is the total number of separate samples
reported in previous studies. Lithogenic Mo correction applied as: δ98MoSEDSXSEDS =
δ98MoLITHXLITH + δ98MoAUTHXAUTH. Authigenic Mo in sediments was estimated from the
relationship: MoAUTH = MoMEASURED - [(Mo/AlLITHOGENIC) x AlMEASURED]. The background
lithogenic Mo:Al ratio used in this calculation was 11 × 10-6 (g g-1); an average of
background minimum Mo:Al values previously measured from sites along the California and
Chile margins, ranging from 8 to 14 x 10-6 (g g-1) (Siebert et al., 2006). Authigenic Mo
accumulation rate (MoAUTH) is calculated using the authigenic Mo concentration and the
local sedimentation rate as determined by 210Pb analyses. Carbon oxidation rates and bottom
water oxygen concentrations previously reported in Berelson et al. (2005). For the Soledad
and San Blas sites, total CO2 profiles showed curvature indicative of some bioirrigation,
such that the values for carbon oxidation are likely to be minimum values.
70
71
MOLYBDENUM BEHAVIOR DURING EARLY DIAGENESIS: INSIGHTS FROM
Mo ISOTOPES
Rebecca L. Poulson, James McManus, Silke Severmann, and William M. Berelson
72
ABSTRACT
This study presents molybdenum concentration and isotope data from surface
sediments of the central eastern tropical Pacific and the coastal Peru, Mexico, and Southern
California margins spanning a wide range of sedimentary diagenetic conditions. The
environments studied have been chosen to represent a broad range in oxidation-reduction
(redox) potential, which provide a framework for the behavior of this redox-sensitive
element. Molybdenum concentrations in marine sediments reflect a combination of three
primary sources: lithogenic Mo associated with continental weathering, biogenic Mo
associated with organic matter, and authigenic Mo deposited via either oxic (sorption to Mnoxides) or anoxic (precipitation of Fe-Mo-S solids) mechanisms. Our data suggest these
components have distinct Mo isotope compositions, and all modern marine sediments appear
to reflect some mixture of these sources. Molybdenum associated with organic matter
represents an isotopically discrete source of Mo to the bulk sediment inventory with a
composition of δ98/95Mo = ~ 2.0‰, and this biogenic Mo dominates the sedimentary Mo
pool in surface sediments at some locations. Authigenic Fe-Mo-S deposit with a seawater
aqueous Mo source have a unique Mo isotopic signature (δ98/95Mo = 1.63 ± 0.02 ‰, SDOM,
n=136). Manganese-rich hemipelagic sediments from the eastern tropical Pacific also have a
unique Mo isotope signature (δ98/95Mo = -0.5±0.1‰, n=14) that reflects fractionation
between ocean water and authigenic Mo associated with Mn oxides. In addition, redox
cycling of Mn within the sediment column appears to strongly influence Mo geochemical
and isotopic behavior.
73
INTRODUCTION
Information regarding the geochemical evolution of the global oceans is contained
within the marine sediment and rock records. Interpretation of these ancient records
leverages off our often imperfect interpretation of the chemical signatures that are
sequestered within remnant geologic materials. Specific to this work, sediment distributions
of molybdenum appear to be sensitive to the availability of reduced sulfur species, which is
often assumed to imply a lack of dissolved oxygen. This observation has led to the use of
sediment Mo abundance as a proxy for past reducing conditions (e.g., Crusius et al., 1996;
Dean et al., 2006). However, Mo is also enriched within metal oxides and these deposits are
generally formed in the presence of dissolved oxygen (e.g., Bertine and Turekian, 1973;
Calvert and Pedersen, 1993; Chappaz et al., 2008). Because authigenic Mo enrichments
occur in both oxidized and reducing environments, sediment Mo concentrations alone are
inadequate proxies for depositional redox conditions. Recent work has identified sediment
Mo isotope signatures that are unique to the specific mechanisms that control Mo speciation
and sedimentary enrichment (Barling et al., 2001; Siebert et al., 2003; Poulson et al., 2006).
Therefore, Mo isotopes, in conjunction with elemental abundances may provide a more
robust paleochemical proxy than either Mo concentrations or isotope compositions alone.
Laboratory experiments and natural samples have quantified Mo isotope
fractionation in Mn-dominated systems (Barling et al., 2001; Siebert et al., 2003; Barling and
Anbar, 2004); however, the possible range in Mo isotope fractionation under more reducing
conditions remains poorly defined. In this study, Mo geochemistry is investigated in
reducing sediments from three continental margins of the eastern Pacific, as well as two
open-ocean Mn-rich sites (Figure 1). Defining the possible range of isotope signatures and
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developing an understanding of the mechanisms responsible for the observed sediment Mo
isotope signatures is an important prerequisite for employing Mo as an effective proxy for
local or even global geochemical fluctuations.
MOLYBDENUM IN THE MARINE ENVIRONMENT
Under the oxygenated conditions predominant in the modern ocean, Mo exists
primarily as the soluble molybdate ion (MoO42-; Figure 2; e.g. Emerson and Huested, 1991),
and is the most abundant dissolved trace element in modern oxic seawater (Broecker and
Peng, 1982). Though considered an essential micronutrient (e.g., Mendel and Bittner, 2006
and references therein), Mo behaves conservatively in the open ocean water column with a
concentration of ~105 nM and a residence time of ~800,000 years (Collier, 1985; Emerson
and Huested, 1991). In addition, although there are currently only a limited number of
analyses (n = 6), modern seawater is thought to have a homogenous Mo isotopic
composition of δ98MoSW = 2.3 ± 0.1‰ (Barling et al., 2001; Siebert et al., 2003; Figure 3;
δ98Mo = ([98/95MoSAMPLE/98/95MoSTANDARD -1] x 1000)), as would be expected given its long
oceanic residence time.
Although modern seawater appears to be a uniform Mo reservoir both in terms of its
concentration and isotopic composition, analyses of marine sediments reveal that Mo
geochemical behavior can be quite dynamic (e.g. Siebert et al., 2006). When interpreting the
sediment record, it is important to recognize that bulk sediment Mo concentrations and
isotope compositions reflect multiple sources and processes that contribute to the total solidphase Mo (Figure 2). Delivery of Mo to marine sediments can be thought of as the sum of
three dominant processes: 1) incorporation of lithogenic Mo into bulk sediment through
75
continental weathering; 2) association of Mo with biological material which is delivered
directly to the seafloor; 3) precipitation or adsorption as an authigenic solid phase (under
both oxic and anoxic conditions).
Lithogenic Mo
Some fraction of all marine sediments contains a component of continental material,
and the relative importance of this terrigenous detrital component depends on a number of
processes. Though continental material delivers only a small quantity of Mo to marine
sediments (e.g., Turekian and Wedepohl, 1961; Taylor and McLennan, 1985), this
contribution represents a discrete fraction of the total measured bulk sediment Mo (Figure
2). Analyses of various terrigenous materials (e.g., granites, clastic sediments; n=12) suggest
a homogenous isotopic composition of δ98Mo = 0.0 ± 0.2‰ (Figure 3; Siebert et al., 2003),
and this value is taken to represent the lithogenic Mo component in bulk sediment
(δ98MoLITH). To constrain the isotopic composition of sedimentary authigenic Mo,
measurements of bulk sediment Mo isotope compositions require correction for dilution by
the lithogenic contribution (Poulson et al., 2006).
Biogenic Mo
Molybdenum is considered a biologically essential trace element, playing a key
enzymatic role in a variety of processes, notably nitrogen fixation and nitrate reduction (e.g.,
Mendel and Bittner, 2006 and references therein). The relationship between organic matter
and Mo is complex because Mo is not only incorporated into cells, but it can also be sorbed
to organic material in the water column (Tribovillard et al., 2004). There are limited data
available to constrain the organic matter Mo:C ratio. Reported Mo:C ratios in the nitrogen-
76
fixing bacteria Trichodesmium erythraeum show differences in the Mo:C ratios of natural
and cultured samples (23 and 3 μmol/mol, respectively; Tuit et al., 2004). Available
sediment trap studies report Mo:C ratios of ~9 nmol/mmol (Mazatlan margin; Nameroff et
al., 1996) and ~4 nmol/mmol (Santa Barbara Basin; Zheng et al., 2000) in sediment trap
materials. It is quite likely that Mo:C ratios in organic matter are variable, as they are
dependent upon multiple environmental factors. In addition, it is also likely that the
preservation of Mo associated with organic material will vary. Recent experimental work has
reported a -0.5‰ δ98Mo isotope fractionation associated with biological assimilation of Mo
(Figure 3; Wasylenki et al., 2007; Liermann et al., 2005). Biogenic Mo (Mo associated with
organic matter) is generally only a small fraction of the total sediment Mo pool, but its
isotopic contribution must be considered.
Authigenic Mo
Authigenic enrichment of Mo occurs through different mechanisms under both oxic
and anoxic conditions. In the presence of oxygen, Mo has been shown to associate with
solid-phase Mn and Fe-oxides (e.g., Bertine and Turekian, 1973; Calvert and Pedersen,
1993; Chappaz et al., 2008), and adsorption to Mn-oxides results in both sediment Mo
enrichment and Mo isotopic fractionation (e.g., Barling et al., 2001; Figures 2 and 3).
Experimental work by Barling and Anbar (2004) revealed a large (2.7‰) fractionation
between soluble molybdate (MoO42-) and Mo sorbed to Mn-oxides in the laboratory; that is,
Mn-associated Mo had a light isotopic signature relative to the dissolved molybdate phase
(Figure 3). These findings are consistent with field results (Barling et al., 2001; Siebert et al.,
2003), which exhibit a similar fractionation between seawater molybdate and Mn-associated
Mo in ferro-manganese crusts or nodules (Figure 3). The specific mechanism responsible for
77
the observed isotope fractionations is not well constrained, though quantum mechanical
calculations suggest the fractionation may reflect adsorption of a minor aqueous species
(MoO3) to the Mn-oxide surface (Δ98MoMoO4-MoO3 = 2.4‰; Tossell, 2005). Though the
mechanisms remain unclear, Mn-controlled authigenic Mo enrichments have the most
negative sediment Mo isotope compositions measured to date.
Molybdenum association with metal oxides is often dynamic in marine sediments.
During organic matter diagenesis, once oxygen is consumed Mn and Fe reduction may
ensue, releasing sorbed Mo back into solution. If Mn is reoxidized within the sedimentary
column during diagenesis, some dissolved Mo may re-adsorb to this reoxidized metal.
Alternatively, some dissolved Mo may also diffuse downward in the sediment column if
there is a deeper sedimentary Mo sink.
Under anoxic sedimentary conditions where sulfate reduction is the dominant
microbially mediated organic matter degradation process, Mo is sequestered into sediments
through complexation with sulfide, forming less soluble thiomolybdates (MoOxS4-x2-) that
may be scavenged by sulfidized organic matter or Fe-sulfide phases such as pyrite (Helz et
al., 1996, 2004; Zheng et al., 2000). Helz et al. (1996) proposed a sulfide-controlled
geochemical “switch” for Mo at ~10 μM H2S(aq), where the dominant dissolved Mo phase
abruptly transitions from molybdate (MoO42-) to tetrathiomolybdate (MoS42). The pore water
work of Zheng et al. (2000) proposed two thresholds for Mo-sulfide formation; at H2S(aq)
concentrations of ~0.01 μM these authors proposed that Mo is removed from solution via
coprecipitation of Fe-Mo-S phases, whereas at higher H2S(aq) concentrations (~10 μM) they
postulate that Mo precipitates independent of iron. It may be that the sulfide thresholds
proposed by Zheng et al. (2000) reflect changes in aqueous Mo speciation that impact solidphase Mo behavior. At low sulfide concentrations, thiomolybdate intermediate species
78
(MoOxS4-x2-) may dominate the aqueous phase and be scavenged by solid-phase Fe-sulfides,
while at higher sulfide concentrations tetrathiomolybdate (MoS42-) is likely to dominate,
precipitating independently as a solid phase Mo-sulfide.
An investigation of pore waters from Santa Monica Basin predicted a fractionation
of -0.7‰ between pore fluids and sediment Mo deposits under reducing conditions
(McManus et al., 2002). Anoxic sediments from three sites on the Mexican continental
margin suggest a unique Mo isotopic signature of δ98Mo = 1.6±0.1‰ for Mo-sulfide
sediment enrichments (Poulson et al., 2006), which is consistent with the predicted
fractionation from a dissolved seawater Mo source (Figure 3). Reported Mo isotope
compositions from “suboxic” surface sediments of the California margin span the full range
between Mn-dominated and more reducing environments (Figure 3; Siebert et al., 2006;
Poulson et al, 2006). Sediments from the most anoxic basins (Santa Monica and San Pedro)
have Mo isotope signatures consistent with those observed on the Mexican margin (core
average δ98Mo values of 1.4‰ and 1.6‰; respectively), while sediments from the welloxygenated San Clemente basin have reported Mo isotope values consistent with Mnassociated Mo (core average δ98Mo = -0.8‰; Figure 3; Siebert et al., 2006; Poulson et al,
2006). In Tanner basin, a site where environmental conditions are between these two
extremes, sediment Mo isotope compositions are intermediate and more variable than those
reported in other settings (core average 0.5‰; Figure 3; Siebert et al., 2006; Poulson et al,
2006).
The range of Mo isotope compositions measured on the California margin, and our
incomplete understanding of the mechanisms responsible for this variability, demonstrate the
need for further refinement of the Mo isotope system in marine sediments. This study aims
to further constrain Mo distributions and isotopic fractionation in the marine environment.
79
As described in detail below, we have selected sites that represent end-member cases for
authigenic Mo enrichment (under both oxic and anoxic conditions) as well as additional sites
from the California and Mexico margins where intermediate (“suboxic”) conditions prevail.
SITE DESCRIPTIONS
MANOP Sites
Two sediment cores from the eastern tropical Pacific (cores collected as part of the
Manganese Nodule Program, MANOP sites M and H) represent the most oxic conditions of
all sites analyzed in this Mo study; sediments are Mn-rich (1-7 wt.% Mn, Lyle et al., 1984)
and bathed in well-oxygenated waters (110-150 μM, e.g. Bender and Heggie, 1984). Both
sites are from water depths greater than 3000 m; site M is located ~25 km east of the East
Pacific Rise, and site H is located in the Guatemala Basin (Figure 1, Table 1). Site M was
selected for the original MANOP study to investigate hydrothermal sedimentation (Lyle et
al., 1984). Site H, originally selected as a representative hemipelagic site, is marked by the
presence of Mn-rich ferromanganese nodules (Finney et al., 1984).
Peru Margin
The Peru margin is a region of wind-driven perennial upwelling, resulting in high
productivity in the surface waters and an associated intense water column oxygen minimum
zone (OMZ, <5 μM O2) (Suess et al., 1986). Sediments of the Peru margin investigated in
this work were collected from a shelf site near 13oS cored at 264 m water depth (Figure 1,
Table 1). This site represents the most reducing open-ocean conditions of all study sites, due
to the high organic carbon content and sulfidic nature of the sediments in this region (e.g.,
80
Reimers and Suess, 1983; Froelich et al., 1988). Several authors have noted the presence of
free sulfide in pore waters from surface sediments (<15 cm) on this margin, but sulfide is
absent from the bottom water (e.g., Froelich et al., 1988; Fossing 1990; Böning et al., 2004).
California Margin
Off the coast of California, intense seasonal upwelling enhances biological
productivity, and the associated high productivity zone bordering the East Pacific (combined
with circulation patterns) results in an OMZ in the water column between depths of 200 and
1000 m (e.g., Sverdrup and Allen, 1939). The four basins investigated along the California
margin are all part of a larger complex known as the Borderland Basins region, described in
detail by Emery (1960) (Figure 4, Table 1). This entire region is characterized by a general
deepening of the basins with distance offshore; basins along each parallel belt become
deeper to the southeast. In general, the nearshore basin sites are the most reducing; bottom
water oxygen contents increase in basins with distance offshore (e.g. Berelson et al., 1996).
The nearshore basins investigated here, Santa Barbara and Santa Monica, are part of
a single “belt” of basins ~30 km offshore (Figure 4). The Santa Barbara Basin is 580 m deep,
with a western sill to the Pacific at ~450 - 475 m depth (Sholkovitz, 1973). The Santa
Monica Basin is 930 m deep, and is connected with the San Pedro Basin by a sill at 740 m
(Emery, 1960). These inner basins have sill depths well within the OMZ, and these sites
have the lowest measured bottom water oxygen concentrations (<10 μM) (Reimers, 1987;
Berelson et al., 1987; Jahnke, 1990). The deepest portions of these basins are characterized
by laminated sediments (Finney and Huh, 1989), suggesting that diffusion generally
dominates solute exchange over bioirrigation and bioturbation (Berelson et al., 1987). The
inner basins are the most reducing environments studied on this margin; sulfate reduction
81
dominates organic matter remineralization at these sites (e.g., Kaplan et al., 1963; Berelson
et al., 1987; Jahnke, 1990).
The other two basins investigated in this study, Santa Catalina and San Nicolas, are
located further offshore (Figure 4); Santa Catalina is located ~80 km from the coast, and San
Nicolas is one of three basins in the next “belt” further west (~130 km offshore). Santa
Catalina (~1300 m deep) and San Nicolas (~1800 m deep) basins both have sill depths at or
just below the base of the OMZ (1000-1200 m) (Emery, 1960), and measured bottom water
concentrations are generally between 15-35 μM (Reimers, 1987; Berelson et al., 1987). At
these sites, organic matter is primarily oxidized by suboxic reactions (Mn and Fe reduction)
(e.g., Berelson et al., 1987; Shaw et al., 1990).
Mexico Margin
The Mexican margin sites investigated in this study experiences the same dominant
currents and hydrography as the region investigated off southern California. Though
productivity off Southern California is generally higher than off of southern Baja California
(van Geen et al., 2003), the same oxygen-deficient North Pacific Intermediate Water
dominates subsurface currents and establishes an OMZ between depths of 500 and 1000 m
(Thunell, 1998). This low oxygen core extends >1500 km off the coast of Mexico (Sansone
et al., 2004). Today, the OMZ is deeper and less intense off central California (~35oN) than
off the Mexican margin (~23oN), reflecting ventilation in the north (van Geen et al., 2003).
This difference explains why laminated sediments (indicative of low bottom water oxygen)
are generally only found on the open margin within the OMZ at southern sites (south of
~24oN; Keigwin, 1998). Anoxic or very low oxygen (<1 μM) bottom waters throughout this
82
region limit bioturbation and allow for the preservation of laminated sediments underlying
the OMZ (e.g. Calvert, 1966).
Because the sediments on this margin are bathed in low-oxygen waters,
decomposition of organic carbon via aerobic processes is presumably limited. Evidence of
denitrification and Mn reduction has been reported within the water column off mainland
Mexico (e.g., Nameroff et al., 2002; Hartnett and Devol, 2003), suggesting diagenesis within
the sediments is dominated by reactions associated with Fe and S cycling. The diagenetic
production of methane at these sites also confirms reducing conditions within the sediments
(Sansone et al., 2004). All the Mexican margin sites are presumed to contain relatively
anoxic sediments; bottom water oxygen concentrations are low (<5 μM) and laminated
sediments are present at all but the Magdalena site (Berelson et al., 2005).
The sediment samples investigated in this study were all subsamples of cores
previously studied by Sansone et al. (2004) and Berelson et al. (2005) (Figure 5, Table 1).
One station is located off the west coast of Southern Baja on the open margin (Magdalena);
the remaining four sites are located within the Gulf of California. The Magdalena sediment
core analyzed in this work was collected from a water depth of 713 m on the continental
slope, and sediments from this core were bioturbated in the top 2-3cm (Berelson et al.,
2005). Of the four sites located within the Gulf of California (Figure 5), two lie within
depositional basins (Alfonso and La Paz), and two are from the open margin on the eastern
side of the Gulf (Carmen and Pescadero). Alfonso Basin is located in La Paz Bay, and has a
sill depth of ~325 m (Sansone et al., 2004). The sediment core analyzed in this study was
collected from a water depth of 542 m in the deepest part of the basin (Berelson et al., 2005).
La Paz Basin lies just east of Alfonso Basin at the northeast edge of the bay; this site was
cored at ~900 m water depth (Berelson et al., 2005). Two cores were analyzed from the
83
Carmen margin continental slope (collected at 575 and 800 m water depth) on the eastern
side of the Gulf of California (Berelson et al., 2005). The Pescadero slope, also on the
eastern margin of the Gulf of California, was cored at 506 and 600 m water depth (Berelson
et al., 2005).
METHODS
All sediment cores from the California, Mexico, and Peru margins were collected
using a multi-corer (Barnett et al., 1984). Organic carbon was measured using an elemental
analyzer, with samples first acidified to remove inorganic carbon prior to analysis (after
Verardo et al., 1990). Solid-phase metal analyses were performed on 50-100 mg of dry
ground bulk sediment samples digested using a series of HCl, HNO3, and HF digestion steps
(either on a hot plate or by microwave digestion (CEM, MARS 5000)). These two methods
are generally analytically indistinguishable (Appendix Tables 1, 2, and 3). Major element
compositions (Al, Ca, Fe, Mn, and Ti) were measured on total sample digestions by
inductively-coupled plasma optical emission spectrometry (ICP-OES, Teledyne Leeman
Prodigy; Appendix Tables 1 and 3). For the same bulk sediment sample digestions, trace
element concentrations were determined by inductively-coupled plasma mass spectrometry
(ICPMS, Thermo PQ ExCell; Appendix Tables 1 and 2).
The reproducibility of analytical techniques was evaluated by performing replicate
analyses of multiple standard reference materials (Appendix Table 1). Major element
concentrations (Al, Fe, Mn, and Ti) analyzed by ICP-OES for all standard reference
materials are typically reproducible within 5% (1-SD), and agree reasonably well with
previously reported values (Appendix Table 1). For Mo concentrations determined by
84
ICPMS, as well as those produced during isotopic analyses (Nu Instruments high resolution
multi-collector inductively-coupled plasma mass spectrometer, MC-ICPMS), standard
reference materials were typically reproducible to <13% (1-SD) and agree with published
values, with the exception of the standard reference material NBS-1645 (Appendix Table 1).
This standard is a river sediment standard, and it has been the most difficult matrix for our
group to reproduce analytically (as noted by its relatively high uncertainty). Our Mo
concentration data for NBS-1645 does not agree with the published value (34 ppm, Potts et
al., 1992), but we have confidence in our value (18 ± 2 ppm), as it represents replicate
digestions, multiple analytical techniques, and analyses of 38 separate sample aliquots
(Appendix Table 1).
Separate bulk sediment samples (~100 mg) were digested for Mo isotopic analyses
(Appendix Tables 1 and 4). Samples were spiked with a 97Mo and 100Mo double isotope
tracer and Mo was separated from the sediment matrix using a previously published column
separation technique (Siebert et al., 2001, 2006). Mo isotope compositions were analyzed on
a Nu Instruments MC-ICPMS. All Mo isotope measurements are reported relative to a
Johnson Matthey ICP Mo standard solution (lot #602332B). At present there is not an
accepted standard for interlaboratory comparison; normalizing measured values to different
standards could potentially generate offsets between reported Mo isotope values from
different lab groups.
Four of the standard reference materials have also been run several times to evaluate
the reproducibility of Mo isotope analyses (PACS-2, n=11, SDO-1, n=10, SX-12280, n=25
and NBS-1645, n=38); reproducibility for these standards is ≤ 0.3‰ (Appendix Table 1). I
present these different reference materials in part because there is no internationally accepted
standard reference material for Mo, and wish to establish a baseline of comparison for our
85
data. In general, analyses of individual samples tended to reproduce better than the standard
reference materials with an average of 0.1‰ (e.g., Appendix Table 4).
Replicate digestions and ICP-OES, ICPMS, and MC-ICPMS analyses were
performed on ~20% of all natural sediment samples in this study (Appendix Tables 2, 3, and
4). The average reproducibility of these data for Mo concentration is better than 10% (1SD), which is consistent with other measures of reproducibility. There are a number of
samples within the data set for which analyses reproduced poorly, but are nonetheless part of
the average. In general, these analytical anomalies have little impact on end-members that
emerge from the overall data set, given its large size and breadth of environmental coverage.
RESULTS & DISCUSSION
The total Mo reservoir in marine sediments reflects multiple Mo sources with unique
Mo isotopic compositions (Figure 2). The fractions of each Mo component relative to the
total sediment Mo concentration (Xn) and their isotopic signatures (δ98Mon) can be expressed
by the mass balance equation:
1)
δ98MoMEAS = δ98MoLITH(XLITH) + δ98MoBIO(XBIO) + δ98MoMn-AUTH(XMn-AUTH) +
δ98MoS-AUTH(XS-AUTH)
In this model X is the fraction of each Mo pool, MoLITH represents the lithogenic Mo
associated with continental weathering, MoBIO represents the biogenic Mo fraction
associated with organic matter deposition, and the authigenic Mo enrichments are
represented as MoMn-AUTH (Mo associated with Mn-oxides) and MoS-AUTH (Mo-sulfides
86
formed in reducing sediments). It is often possible to simplify this equation; for example,
Mn-controlled authigenic Mo deposits occur under oxygenated conditions where Mo-sulfide
formation is negligible. The relative importance of each source to the total bulk Mo reservoir
is dependent upon geochemical conditions specific to the sedimentary environment.
The Lithogenic Mo Correction
As previously discussed, the fraction of lithogenic Mo in bulk sediment samples
(XLITH) can be estimated from a regional background Mo:Al ratio. For all sites in this study,
a lithogenic Mo:Al ratio of 11 x 10-6 was used, the median value from a range of background
Mo:Al values (8 to 14 x 10-6) previously observed in sediments from the Californian and
Chilean margins (McManus et al., 2006). This value is also consistent with typical reported
values for igneous rocks and sandstones (Mo:Al = 6 to19 x 10-6 in Turekian and Wedepohl
(1961); Mo:Al = ~19 x 10-6 in Taylor and McLennan (1985)). The fraction of lithogenic Mo
was calculated as:
2) MoLITH = AlMEAS x (MoLITH:AlLITH); MoLITH = AlMEAS x (11x10-6)
For all sites in this study, the estimated lithogenic Mo contribution was ≤ 1 ppm Mo
(Appendix Table 5).
The δ98Mo value of 0.0‰ measured in terrigenous materials (Siebert et al., 2003)
has been suggested as the δ98MoLITH isotopic signature, such that measured Mo isotope
compositions of bulk sediment samples can be corrected for “dilution” by the lithogenic
fraction (Poulson et al., 2006). For all sediments in this study, the mass balance (Equation 1)
was simplified to:
87
3) δ98MoMEAS = δ98MoLITH(XLITH) + δ98MoENRICH(XENRICH)
where MoENRICH represents the sediment Mo enrichment over any lithogenic contribution.
For all sedimentary environments, MoENRICH represents some combination of authigenic
enrichment (either Mn- or S-controlled) and organic matter deposition (Figure 2), such that:
4) δ98MoENRICH = δ98MoBIO(XBIO) + δ98MoMn-AUTH(XMn-AUTH) + δ98MoS-AUTH(XS-AUTH)
This lithogenic Mo correction allows for more accurate interpretation of Mo isotopic
variability in sediment Mo enrichments, reported herein as δ98MoENRICH values (Table 1,
Appendix Table 5). It is important to recognize that this choice for the lithogenic Mo
signature is simply our "best guess" given the currently available data set. Later refinements
may deliver a different estimate for the lithogenic component, both in terms of concentration
and isotopic composition. It is also important to recognize that for many sites, the potential
inaccuracies in this estimate are small when compared to the bulk signature.
The Authigenic Mo-Manganese Signature
Surface sediments from the Mn-rich MANOP sites H and M in the eastern
Equatorial Pacific have significant MoENRICH concentrations associated with Mo sorbed to
Mn-oxides (Figure 6, Table 1). Site H, in particular, may be considered “proto-typical” Mnrich sediment, with Mn concentrations of ~5 wt.% (Lyle et al., 1984; Appendix Table 3).
Likewise, pore water Mn data at site H indicate that sediments are oxygenated to ~12 cm,
with Mn reduction below this depth (Klinkhammer, 1980; Figure 6). Sediment Mo
88
concentrations reflect the process of Mn reduction at depth; the upper ~10 cm are highly
enriched in Mo (>50 ppm) with MoENRICH concentrations decreasing below this depth
(Figure 6, Table 1). This solid phase Mo decrease suggests Mo is released back into pore
fluids as host Mn-oxide is reduced. We take Site H to represent an end-member case for
open-ocean authigenic Mo enrichment associated with Mn-oxides (MoMn-AUTH), such that:
5) δ98MoENRICH = δ98MoMn-AUTH(XMn-AUTH)
Consistent with the dominance of Mn cycling, site H sediments have the most
negative Mo isotopic compositions measured in this study (Table 1) with an average
δ98MoENRICH value for all site H samples of -0.5±0.1‰ (n=14; Table 1; Figure 6). This value
suggests a fractionation between a seawater aqueous Mo source and the authigenic Mo pool
of Δ98MoSW-Mn-AUTH = 2.8‰; consistent with previously reported Mo isotope data from FeMn crusts (Figure 3; Siebert et al., 2003; Barling et al., 2001), previously reported data from
Mn-rich sediments (Siebert et al., 2006; Poulson et al., 2006), and the experimental work of
Barling and Anbar (2004).
Site M sediments are less fractionated (relative to seawater) than those measured at
site H; nevertheless, δ98MoENRICH values from site M are generally negative, suggesting Mn
cycling is a primary control on Mo behavior at this site (Table 1). Pore water data suggests
Mn reduction at a depth of only ~5 cm at site M (Klinkhammer, 1980) and solid phase
MoENRICH concentrations also decrease below this depth (Table 1). It is worth noting that the
δ98MoENRICH value calculated for the deepest sample at this site is suspect (-0.8‰; Table 1,
Appendix Table 5); given the uncertainties associated with the lithogenic Mo estimate, the
89
magnitude of the lithogenic correction for this sample (XLITH = 0.87; Table 1) may have
produced a spurious result.
The Authigenic Mo-Sulfide Signature
Surface sediments from the Peru continental margin have the highest authigenic Mo
concentrations of all sites analyzed in this study (>80 ppm; Figure 7; Table 1). The Peru site
is the most enriched in organic carbon (>14% Corg; Appendix Table 5), due to the influence
of the Peru coastal upwelling system (Suess et al., 1986). This high Corg, low O2 environment
leads to anoxic diagenesis in surface sediments; pore water data from a similar site on the
Peru margin reveal H2S concentrations >1mM within the uppermost ~20 cm, with the
highest sulfate reduction rates observed within a few cm of the sediment surface (Fossing,
1990). Authigenic Mo dominates the bulk sediment Mo pool throughout this core (XLITH ≤
0.02; Table 1); we therefore take this site to represent the end-member case for open-ocean
authigenic Mo enrichment associated with Fe-Mo-S and/or Mo-S precipitation (MoS-AUTH),
such that:
6) δ98MoENRICH = δ98MoS-AUTH(XS-AUTH)
The average δ98MoENRICH value for all Peru samples is 1.5±0.1‰ (n=10; Figure 7;
Table 1). This value suggests a fractionation between a seawater aqueous Mo source and the
authigenic Mo pool of Δ98MoSW-AUTH-S = 0.8‰, a value consistent with pore water
predictions (Δ98Mo = 0.7‰; McManus et al., 2002) and the previously reported anoxic
sediment Mo isotope signature (δ98Mo = 1.6‰; Poulson et al., 2006; Figure 3). Authigenic
Mo enrichment in anoxic environments is controlled by the formation and deposition of Mo-
90
sulfides (e.g. Helz et al., 1996); however, it is unknown whether the observed Mo isotope
fractionation occurs during aqueous Mo species transformations or results from processes
controlling solid phase Mo enrichment.
We suggest that sorption of Mo-sulfides to pyrite may be the mechanism responsible
for the observed isotopic signature. Pyrite (FeS2) is thought to be the most important hostphase for Mo in anoxic sediments (e.g., Huerta-Diaz and Morse, 1992). In fact, recent work
has suggested that pyrite formation in reducing sedimentary microenvironments may capture
Mo more efficiently than previously believed (Tribovillard et al., 2008). Laboratory
experiments have demonstrated that sorption of Mo-sulfides to pyrite results in the formation
of a Mo-Fe-S cubane structure that, once formed, is highly resistant to desorption (Helz et
al., 1996). The formation of Mo-Fe-S cubanes significantly alters the bonding environment
around the Mo atom (Bostick et al., 2003; Vorlicek et al., 2004), and this restructuring is a
plausible mechanism for fractionating Mo isotopes. Experimental work has shown the
degree of Mo-sulfide sorption to pyrite varies with dissolved sulfide concentration; at high
dissolved sulfide concentrations, sorption of Mo-sulfides is suppressed, implying that sulfide
competes with Mo for surface sites (Bostick et al., 2003).
It may be that the primary mechanism responsible for both solid phase Mo
enrichment and isotopic fractionation in reducing environments is sorption to pyrite, and that
this mechanism will be sulfide-dependent. At low dissolved sulfide concentrations, most
sediment Mo-sulfides should therefore be sorbed to pyrite, potentially resulting in highly
fractionated sediments (relative to the seawater aqueous Mo source). This process is one
explanation for the sediment Mo isotope compositions observed in this study. Further
experimental work is warranted to constrain the role of sorption as a mechanism for
fractionating Mo isotopes.
91
At present, however, we cannot dismiss the possibility that Mo isotopes are
fractionated by processes associated with aqueous phase transformations. It is possible that
the formation of thiomolybdate species (MoOxS4-x2-) fractionates Mo isotopes in the aqueous
phase. In fact, quantum mechanical calculations predict a large (~7‰) fractionation between
MoO42- and MoS42- species (Tossell, 2005). Experimental work has shown that, in the
presence of both H2S and S0-electron donors, thiomolybdate Mo(VI) may be reduced to
Mo(V) or Mo(IV) polysulfide anions (Vorlicek et al., 2004). Changes in bonding around the
Mo atom, whether associated with S and O substitutions or with reduction of Mo, could
result in isotopic fractionation between dissolved Mo species. Subsequent scavenging and
deposition of these fractionated Mo species may be responsible for the observed authigenic
signature.
The Biogenic Mo Signature
Data from sites on the Mexico margin suggest that biogenic Mo (MoBIO) may be a
dominant sedimentary component in some marine environments. In particular, sediments
from both cores taken on the Pescadero slope have the lowest concentrations of MoENRICH
(2.0±0.2 ppm; n=19; Figure 8) and the lowest observed sediment MoENRICH:Corg ratios (~0.5)
of all sites analyzed on the margin (Table 1). Though it is not possible to accurately quantify
the relative contributions of biogenic and authigenic Mo in these sediments, the low
MoENRICH:Corg ratios indicate these sediments are likely the least impacted by authigenic Mo
enrichment. We therefore suggest the following mass balance for these sediments:
7) δ98MoMEAS = δ98MoLITH(XLITH) + δ98MoBIO(XBIO); or δ98MoENRICH = δ98MoBIO(XBIO)
92
and we expect biogenic Mo (MoBIO) to dominate the δ98MoENRICH values measured at these
sites.
Sediments from both Pescadero sites have the heaviest δ98MoENRICH values measured
on the margin, specifically in the uppermost ~4cm, averaging 2.1±0.2‰ for all samples over
this depth range (n=11; Figure 8; Table 1). Assuming that this value represents the Mo
isotopic composition of the organic sedimentary Mo component (δ98MoBIO), the data suggest
a small fractionation between seawater Mo and biogenic Mo (~0.2±0.2‰; Figure 8); similar
to the fractionation of 0.5‰ reported for biological uptake of Mo from solution in laboratory
experiments (Figure 3; Wasylenki et al., 2007; Liermann et al., 2005). Further analyses of
organic matter-associated Mo are warranted to discern if a unique isotopic signature for
biogenic Mo exists, but it does appear that biogenic Mo can be an isotopically relevant
fraction of the bulk Mo in certain environments. However, the ultimate fate of this
component remains unclear; biogenic Mo may not survive early diagenetic processes and
thus may have little impact in the rock record. Alternatively, some of this Mo may end up as
a source for deeper Mo precipitation upon organic matter decomposition (and subsequent
Mo release).
Authigenic Mo with Seawater δ98Mo(aq) Source
All the Mexican margin sites are presumed to contain relatively anoxic sediments;
bottom water oxygen concentrations are low (<5 μM) and laminated sediments are present at
all but the Magdalena site (Berelson et al., 2005). These conditions suggest Fe/S-controlled
processes are likely to dominate authigenic Mo enrichment at most if not all of these sites.
Data from the Pescadero sites, however, suggest that Mo associated with organic matter may
93
be a substantial sedimentary component in some locations, and we therefore propose a more
complete sediment mass balance for Mexican margin sediments:
8) δ98MoMEAS = δ98MoLITH(XLITH) + δ98MoBIO(XBIO) + δ98MoS-AUTH(XS-AUTH)
The sediment δ98MoENRICH values are taken to reflect a combination of both sedimentary
biogenic and authigenic Mo phases (MoBIO and MoAUTH-S).
Downcore data from the Magdalena margin suggest the presence of two isotopically
discrete sedimentary Mo sources (Figure 9). The sediments are bioturbated in the uppermost
2-3 cm (Berelson et al., 2005), and mixing likely inhibits the formation of Fe/S-controlled
authigenic Mo deposits in the most surficial sediments. Sediment MoENRICH concentrations
remain low throughout the uppermost ~3 cm, increasing below this depth (Figure 9, Table
1). Within the mixed layer, sediment δ98MoENRICH values are consistent with a biogenic Mo
component (1.9±0.1‰; n=5; Figure 9; Table 1). Below the mixed layer sediments are more
fractionated (relative to a seawater source), suggesting that an authigenic Mo phase
dominates the sediment Mo pool in the deepest portions of the core (average 1.3±0.1‰; n=4;
Figure 9; Table 1). It is worth noting that this site also has the highest sediment Ca
concentrations of all sites analyzed (~13 wt.% Ca; Appendix Table 3), and carbonate mineral
phases may play an as yet undefined role in the sediment Mo isotope compositions observed.
Sediment data from Alfonso basin suggests a gradual transition from organic matterdominated sedimentary Mo to Fe/S-controlled authigenic enrichment with depth at this site.
Measured δ98MoENRICH values are heaviest near the sediment-water interface (~2.0‰),
suggesting MoBIO dominates the sediment Mo pool (Figure 10; Table 1). Sediment
δ98MoENRICH values are increasingly more fractionated (relative to seawater) with depth,
94
approaching values consistent with the previously reported authigenic Mo signature (1.7±
0.1‰ below ~30cm, n=4; Figure 10; Table 1). Similar Mo isotopic behavior is also observed
in sediments from the La Paz basin; however, the sediment Mo enrichment is less
pronounced over the shorter depth range analyzed (MoENRICH < 9 ppm; 0-7 cm; Figure 10;
Table 1). Data from the Carmen sites suggest little or no increase in MoENRICH with depth
(average for both cores is 5.4± 0.8 ppm; Table 1); intermediate sediment δ98MoENRICH values
suggest Carmen sediments are a mixture of sedimentary biogenic and authigenic Mo phases
(core averages 1.8±0.1‰ and 1.7±0.2‰; Table 1).
Sediments from the most reducing basins analyzed on the California margin (Santa
Barbara and Santa Monica) have relatively invariant average down core δ98MoENRICH values
consistent with those observed on the Peru margin (1.5±0.2‰ and 1.6±0.1‰, respectively;
Table 1). Heavier δ98MoENRICH values suggesting the influence of a biogenic Mo component
are not observed. There is likely a discrete biogenic Mo component present in both the Santa
Barbara and Santa Monica basins, but it may be that this fraction is too small to exert a
detectable influence on the measured sediment Mo isotopic compositions. As stated
previously, Mo:C ratios in sediment trap materials from the California margin are only half
those reported on the Mexico margin, suggesting less Mo is associated with organic matter
deposition in the California sites (~4 nmol/mmol in Santa Barbara Basin, Zheng et al., 2000;
~9 nmol/mmol for Mazatlan margin, Nameroff et al., 1996).
Average δ98MoENRICH values from all sediment cores of the Mexican margin
(excluding Pescadero), as well as those from the Peru margin and the two inner basin
California margin sites, define a mean Mo isotope signature of δ98MoENRICH = 1.7±0.2 ‰ (1SD, n=84; Table 1). Including published data from three additional sites on the Mexican
margin (Appendix Table 6; Poulson et al., 2006), and two additional reducing inner basins of
95
the California margin (San Pedro and Santa Monica; Appendix Table 6; Siebert et al., 2006;
Poulson et al; 2006), the average is 1.63 ± 0.02 ‰ (error is standard deviation of the mean,
SDOM, n=136; Table 1). It appears that Fe/S-controlled authigenic Mo enrichments with a
seawater aqueous Mo source bear a unique Mo isotopic signature (δ98MoAUTH-S) that is
ultimately recorded in marine sediments, despite any additional variability introduced by an
organic matter-associated Mo sedimentary component. This final distinction is indeed
important; many of the sites converge on a single isotope value despite a significant range in
Mo concentrations, very distinct differences in pore water sulfide concentrations, and
presumably differences in the relative contributions of MoAUTH-S and MoBIO. This invariant
Mo isotope signature further strengthens the revised marine Mo budget described in
McManus et al. (2006). It does in fact appear that continental margins represent an important
oceanic sink for Mo, and that such deposits have a unique Mo isotopic composition. This
additional Mo sink complicates the Mo isotope paleoproxy, suggesting the Mo isotope
composition of seawater is not constrained by a simple two-component mass balance
(Arnold et al., 2004; McManus et al., 2006).
Authigenic Mo with Manganese δ98Mo(aq) Source
Sediments from the outer basins of the California margin (Santa Catalina and San
Nicolas) have the most dynamic range in measured δ98MoENRICH compositions of all sites
analyzed in this study (Table 1, Figure 11); from values consistent with organic matter
(δ98MoBIO) to values consistent with Mn-associated authigenic Mo deposits (δ98MoAUTH-Mn).
However, pore water profiles from these sites suggest Mn oxides undergo reductive
dissolution and are not preserved in sediments from both basins (Figure 11), and we
therefore anticipate the ultimate authigenic Mo phase is associated with Fe-Mo-S/Mo-S
96
precipitation (XAUTH-S). We therefore assume that the mass balance equation (Equation 8)
governing the other reducing margin sites of this study applies at these sites as well.
We propose that the same mechanisms responsible for authigenic Mo-sulfide
enrichment in other margin settings also impact Mo behavior at these sites, but that the
aqueous Mo source is not necessarily seawater (δ98MoSW). Instead, we suggest that these
sites typify environments where Mn-cycling within the sediment column influences Mo
isotopic behavior. It appears that fractionated Mo released during Mn-reduction (δ98MoMn)
and Mo associated with organic matter likely supplies the aqueous Mo that is subsequently
deposited in authigenic phases at depth, resulting in more fractionated sediment Mo isotopic
compositions than those observed in other margin settings. Both these outer basin sites have
the lowest sediment MoENRICH concentrations and the lowest MoENRICH:Corg ratios measured
in all sites from this study (Table 1). These low ratios could suggest little or no authigenic
enrichment, or a dominance of biogenic Mo in these sediments, but we suggest that the low
sediment MoENRICH:Corg ratios reflect Mo-poor organic matter preservation on the California
margin relative to that preserved on the Mexican margin. Mo isotope data from these
California sites indicate that biogenic Mo may in fact comprise an important sedimentary
component at these sites, particularly in the most surficial sediments. In addition, despite the
low overall measured sediment MoENRICH concentrations, both cores display slight increases
in solid-phase Mo with depth which presumably reflects authigenic Fe-Mo-S/Mo-S
precipitation at depth (Figure 11, Table 1). The assumption of an Fe-Mo-S/Mo-S authigenic
phase at depth is bolstered by observed increases in total reduced sulfur (up to ~0.5 wt.%) in
the uppermost 30 cm of Santa Catalina sediments (Leslie et al., 1990).
The Mo isotope profile in the Santa Catalina sediments suggests biogenic Mo is the
dominant Mo pool in the most surficial sediments; measured δ98MoENRICH values at the
97
surface are consistent with those observed in the Pescadero sediments from the Mexican
margin (Figure 11, Table 1). δ98MoENRICH values steadily decrease down core, approaching
negative values consistent with Mn-controlled authigenic Mo behavior (Figure 11, Table 1).
We propose that Mo released from Mn-oxides during Mn reduction (δ98MoMn) likely
supplies the aqueous Mo that is subsequently deposited at depth; that is, the initial source of
aqueous Mo to these sediments is isotopically fractionated (relative to seawater), altering the
ultimate Mo isotope composition of authigenic Fe-Mo-S phases.
Pore water profiles from San Nicolas basin suggest a Mn reduction zone at ~3-7 cm
depth, with an oxygenated zone above (Figure 11). Solid phase Mo profiles also suggest this
diagenetic distribution (Figure 11, Table 1), with slight Mo enrichment in the very surface
(possibly sorbed to Mn oxides), Mo depletion in the Mn reduction zone (presumably
released upon reduction of the Mn-oxide host), and gradual authigenic enrichment at depth
(presumably associated with Fe-Mo-S/Mo-S precipitation). The light δ98MoENRICH value of
the uppermost sediment sample suggests a mix of biogenic Mo (MoBIO) and Mo sorbed to
Mn-oxides (MoAUTH-Mn). Within the Mn-reduction zone, sediment δ98MoENRICH values are
heavier, suggesting MoBIO is a more dominant component of the bulk sediment Mo pool. Mo
enrichment at depth likely reflects Fe/S-controlled authigenic Mo enrichment (MoAUTH-S),
but sediment δ98MoENRICH compositions decrease to negative values (Figure 11, Table 1). As
in Santa Catalina, sediment δ98MoENRICH values suggest the aqueous source of Mo for
enrichment is not seawater Mo (δ98MoSW), but Mo released during Mn-reduction (δ98MoMn).
The data from these two sites suggest that, in certain settings, Mn cycling may exert
even greater control of Mo geochemical and isotopic behavior than previously thought.
Manganese reduction within the sediment column may release fractionated Mo back into
solution which is subsequently incorporated into authigenic phases at depth. This process
98
appears capable of generating a broad range of sediment Mo isotope values, and further
work is warranted to better constrain the impact of Mn cycling on sedimentary Mo isotope
signatures.
CONCLUSIONS
Molybdenum concentrations in marine sediments reflect a combination of multiple
primary sources: lithogenic Mo associated with detrital material (MoLITH), biogenic Mo
associated with organic matter deposition (MoBIO), and authigenic Mo deposited via either
oxic (sorption to Mn-oxides, MoAUTH-Mn) or anoxic (precipitation of Fe-Mo-S solids,
MoAUTH-S) mechanisms. These sources each appear to have distinct Mo isotope
compositions, and all modern marine sediments appear to reflect some mixture of these
sources.
MANOP site H and the Peru margin site have the highest MoENRICH concentrations of
all sites analyzed, and both sites have unique and relatively invariant core average
δ98MoENRICH compositions (Figure 12). We therefore presume the bulk sedimentary Mo pool
at these sites to be dominated by authigenic deposits, and take the core average sediment
δ98MoENRICH values to reflect the discrete Mo isotopic signatures of these authigenic Mo
phases. Many of the Mexico and California margin sites have sediment δ98MoENRICH values
consistent with the (δ98MoAUTH-S) authigenic isotope signature, and we propose that sorption
of Mo-sulfides to pyrite may be responsible for the observed fractionation.
Heavier sediment δ98MoENRICH compositions (less fractionated relative to a parent
seawater δ98Mo source) measured on the Pescadero margin (Figure 12), as well as in the
most surficial sediments of many other margin sites, suggest that an additional Mo
99
component may be controlling Mo isotope values near the sediment-water interface. We
propose that Mo associated with organic matter is a discrete source of Mo to marine
sediments that is less fractionated (relative to a seawater source) than authigenic phases, and
that biogenic Mo dominates the sedimentary Mo pool in surface sediments at many
locations. Indeed, core average sediment δ98MoENRICH values from all sites on the Mexican
margin, as well as two reducing inner basins of the California margin, appear to reflect a
mixture of both biogenic (MoBIO) and authigenic (MoAUTH-S) phases (Figure 12). Data from
the Santa Catalina and San Nicolas basin sediments suggest that when Mn reduction is an
important source of aqueous Mo within the sediment column, it is reflected in the isotopic
composition of the authigenic sediment fraction (Figure 12). It appears that different sources
of aqueous Mo (seawater [δ98MoSW] versus Mo released from Mn-oxides [δ98MoMn]) can
generate very different Mo isotopic compositions in the precipitated authigenic phases.
Bulk marine sediment δ98Mo values represent the mass balance of multiple primary
Mo source components. Because these sources have unique isotopic signatures, sediment Mo
isotope compositions reflect the dominant mechanisms responsible for Mo enrichment in
different marine environments. The utility of Mo as a proxy for reconstructing
paleodepositional environments and/or redox conditions therefore lies with sediment Mo
isotope compositions, which reveal more about the ambient geochemical conditions than
sediment Mo enrichments alone. Many aspects of Mo geochemical and isotopic behavior in
marine systems remain unconstrained, but Mo isotopes, in conjunction with elemental
abundances, show promise as a tool for interpreting the ancient record.
100
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Figure 1. Map of study areas showing approximate locations of all sites investigated.
Figure 2. Major Mo sources to modern marine sediments: 1) Lithogenic Mo terrigenous material incorporated into bulk sediment, 2) Biogenic Mo – sorbed to or
incorporated into organic material, 3) Authigenic Mo - directly precipitated as a solid phase
within the sediments (under both oxic and anoxic conditions).
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Figure 3. Published marine Mo isotope values and fractionation factors from Barling et
al., 2001, McManus et al., 2002, Siebert et al., 2003 and 2006, Poulson et al., 2006, and
Wasylenki et al., 2007.
Figure 4. Map of California margin study areas showing approximate locations of all
sites investigated. Sites reported in this study in colored symbols; sites with previously
published Mo isotope data show in grey (Siebert et al., 2006; Poulson et al., 2006).
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Figure 5. Map of Mexico margin study areas showing approximate locations of all sites
investigated. Sites reported in this study in colored symbols; sites with previously published
Mo isotope data show in grey (Poulson et al., 2006).
Figure 6. MANOP site H profiles. All pore water data (left panel) from Klinkhammer (1980). Dashed line indicates
estimated depth of Mn reduction (~12 cm). Sediment Moenrich concentrations (center panel) and δ98Moenrich values (right panel)
from Appendix Table 5. All error bars (ppm and ‰) are 1-SD errors for average replicate sample digestions.
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Figure 7. Peru margin profiles. All pore water data (left panel) from Froelich et al., 1988 (near 12oS, 183m water depth).
Sediment Moenrich concentrations (center panel) and δ98Moenrich values (right panel) from Appendix Table 5. All error bars (ppm
and ‰) are 1-SD errors for average replicate sample digestions.
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Figure 8. Sediment Moenrich concentrations and isotope compositions from Pescadero margin. Sediment
Moenrich concentrations and δ98Moenrich values from Appendix Table 5. All error bars (ppm and ‰) are 1-SD errors
for average replicate sample digestions.
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Figure 9. Sediment Moenrich concentrations and isotope compositions from Magdalena
margin. Hatched section and dashed line indicate bioturbated layer (0-3 cm). Sediment
Moenrich concentrations and δ98Moenrich values from Appendix Table 5. All error bars (ppm
and ‰) are 1-SD errors for average replicate sample digestions.
Figure 10. Sediment Moenrich concentrations, and Mo isotope compositions from Alfonso
and La Paz basins, Mexico margin. Sediment Moenrich concentrations and δ98Moenrich values
from Appendix Table 5. All error bars (ppm and ‰) are 1-SD errors for average replicate
sample digestions.
Figure 11. Pore water Fe and Mn profiles, sediment Moenrich concentrations, and Mo isotope compositions from
Santa Catalina and San Nicolas basins. Pore water Mn and Fe data (left panel) from McManus et al., (1997; 1998;
personal comm.). Sediment Moenrich concentrations (center panel) and δ98Moenrich values (right panel) from Appendix Table
5. All error bars (ppm and ‰) are 1-SD errors for average replicate sample digestions.
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Figure 12. Whole core average Moenrich concentrations versus δ98Moenrich values for all sites in this study. Shaded
regions represent whole core averages (ppm and ‰) and associated 1-SD errors (data from Appendix Table 5). The full
range in average Moenrich concentrations for Peru margin sediments (55 ±24 ppm, n=10) extends past the scale depicted
(as indicated by the red arrow).
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Table 1. General site characteristics and average sediment Moenrich data.
Bottom water oxygen data for all Mexican margin sites (and Santa Monica) from Berelson et
al., 2005. All other bottom water oxygen values compiled from Bender and Heggie (1984);
Berelson et al. (1987) and (2005); McManus et al. (2006). Average sediment Moenrich
concentrations, isotopic compositions, and Mo:C ratios from Appendix Table 5. Fractions of
lithogenic Mo (XLITH) calculated in Appendix Table 5; assumed Mo:Al ratio for lithogenic
background is 1.1 x 10-5 (McManus et al., 2006; Poulson et al., 2006) Details of lithogenic
correction (δ98Moenrich values) described in text.
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Table 1. (Continued)
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Table 1. (Continued)
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Table 1. (Continued)
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SEDIMENT GEOCHEMISTRY ALONG A CHEMOCLINE TRANSECT,
LAKE TANGANYIKA, EAST AFRICA
Rebecca L. Poulson, James McManus, and Silke Severmann
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ABSTRACT
This study presents sediment geochemical data from five sites along a depth transect
in Lake Tanganyika, East Africa. Permanent stratification of the lake waters leads to the
production of a chemocline at ~150 m water depth. We present a variety of radiochemical,
biogenic (organic carbon, carbonate, total reduced sulfur), and trace metal (Mo, U, Re, Cd,
and V) data to characterize the burial response of these elements to changes in sediment
reducing character across this transition. Sediment accumulation rates are derived from 210Pb
and 137Cs profiles, and the resulting age models are verified by reasonable alignment of the
sedimentary calcium carbonate records from all sites. The CaCO3 profiles exhibit peak
concentrations of 60-70% that appear to be centered ~1870 AD, which is roughly coincident
with the termination of the Little Ice Age (c.1850) and the most recent lake high stand
(c.1880). Sediment C/S ratios agree with reported freshwater C/S values at the shallowest
site, but decrease to lower values at the deeper locations. The lacustrine-like shallow
sediment C/S ratios (~22 wt./wt.) are consistent with these sediments generally being sulfur
limited, which is not surprising in this freshwater system. However, the deeper sites exhibit
lower C/S ratios (6± 2), perhaps indicative of a system where diagenesis is dominated by
sulfate reduction. The distribution of redox-sensitive trace metals, specifically U and Mo,
suggest authigenic metal enrichment at these sites, particularly in sediments from depths
below the chemocline. These data, along with the sediment C/S ratios, imply that overlying
water column oxygenation limits both sulfate reduction and authigenic trace metal
enrichment. Sites closest to the reported chemocline depth appear to be governed by similar
geochemistries; that is, there are similarities in the observed trends of carbon, sulfur, and
trace metal behavior at these sites. Despite sulfur limitation in this freshwater system,
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conditions appear to be sufficiently reducing (particularly at depths below the chemocline)
for substantial reduced sulfur accumulation and authigenic metal enrichment within the
sediments.
INTRODUCTION
Lake Tanganyika is the largest of the East African Rift Valley lakes and is the
second largest (by volume) body of freshwater in the world, with an estimated total volume
of ~18,940 km3 (Hutchinson, 1957). The lake occupies a ~650 km north-south trough that is
~50 km wide and ~1.4 km deep at its deepest point (Figure 1, Edmond et al., 1993). The
equatorial "endless summer" climate leads to permanent thermal stratification in the lake
(e.g. Hecky, 2000) with temperatures of 25-27oC typical in the surface waters and a uniform
deep water temperature of 23.3±0.05 oC (Degens et al., 1971; Edmond et al., 1993). Recent
climate warming is reflected in increasing lake temperatures, with a warming trend of 0.1oC
per decade in surface waters over the last century, and an overall increase of ~0.3oC in
measured deep water temperatures since ~1940 (O'Reilly et al., 2003).The thermocline is
positioned at ~40-50 m water depth much of the year, but during the cool trade-wind season
(April-September) mixing of surface waters deepens the thermocline to ~150 m (Huc et al.,
1990; Edmond et al., 1993).
The combined effects of stratification and organic matter degradation produce a
strong chemocline at ~150 m water depth (Figure 2; Degens et al., 1971; Edmond et al.,
1993). Water column profiles are marked by a sharp decrease in oxygen concentrations to <5
μM and depletion of detectable nitrate above the chemocline (Degens et al., 1971; Edmond
et al., 1993). Below ~150 m, ammonia and sulfide concentrations increase steadily, and
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phosphate concentrations increase with depth throughout the entire water column (Figure 2;
Edmond et al., 1993). Despite increasing water column sulfide concentrations, reported
sulfate concentrations are relatively constant at ~40 μM throughout the water column,
decreasing slightly only below ~1000 m water depth (Degens et al., 1971). The pH of the
lake is ~8.5 (Edmond et al., 1993), and waters are supersaturated with respect to calcium
carbonate (Cohen et al., 1997).
This study investigates the geochemistry of surface sediments from depths (~72 m to
332 m) transecting the chemocline off the Luiche Platform, a deltaic deposit off the eastern
shoreline just south of Kigoma, Tanzania (Figure 1, Table 1). A full suite of analytical
techniques are employed to characterize the sediment geochemical conditions above, within,
and below the chemocline. The microbially-mediated breakdown of organic matter, both in
the water column and within the sediments, proceeds through a well-known sequence of
electron donors, similar to what is observed in marine sediments (e.g., Froelich et al., 1979).
Organic matter is oxidized by sequential reduction of the available oxidant with the greatest
free energy change (O2, NO3-, MnO2, Fe2O3, and SO42-). These reactions (in conjunction with
stratification) are responsible for the observed water column chemocline (Figure 2, Degens
et al., 1971; Edmond et al., 1993), and directly affect the behavior of key reaction
constituents within the sediments. Variations in the sediment distribution of these elements
(e.g. C, S, Fe, Mn) are therefore indicative of the dominant diagenetic reactions and the
extent of reducing conditions along the transect.
The sediment distributions of redox-sensitive trace metals (e.g. V, Mo, Cd, Re, U)
are investigated herein to further constrain sediment geochemistry across the chemocline
transect. In general, these metals are more soluble under oxidizing conditions, but changes in
solubility and oxidation state lead to authigenic sedimentary enrichments under reducing
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conditions (e.g. Tribovillard et al., 2006 and references therein). The specific geochemical
mechanisms responsible for authigenic enrichment are different for each of these metals, and
these differences impact the accumulation patterns of these metals across the transect. In
particular, this work investigates the observed trends in authigenic U and Mo accumulation,
and the relationship of these metals to organic carbon delivery and degradation within the
sediments. Previous work in Lake Malawi utilized changes in the distribution of redoxsensitive trace metals over several meters of the sediment column to infer paleo-excursions
in the depth of the chemocline (Brown et al., 2000). This study investigates the accumulation
patterns of redox-sensitive trace metals in surficial Lake Tanganyika sediments (<30 cm),
providing a modern context for future interpretation of metal distributions in longer sediment
records.
METHODS
All sediment cores from the Luiche platform were collected using a multi-corer,
which is a smaller version of that described by Barnett et al., (1984). Two companion cores
for each multi-core deployment were sectioned shipboard; one was stored for 210Pb and 137Cs
determination, the other was sectioned shipboard under nitrogen and samples were
centrifuged for pore water separation. All sediment samples were freeze dried and ground
before further analysis. 210Pb and 137Cs were determined by γ-ray spectroscopy (e.g., Gilmore
and Hemingway, 1995) as described in Wheatcroft and Sommerfield (2005). In short, ~30 g
of dried ground sediment were counted for >24 hr on two equivalent Canberra GL2020RS
LEGe planar (2000 mm2) γ-ray detectors and activities were corrected to the date of core
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collection; detector efficiency and self-absorption corrections are described in Wheatcroft
and Sommerfield (2005).
Solid-phase metal analyses were performed on ~100 mg of dry ground bulk
sediment samples from the cores processed shipboard. Sediments were first digested using a
series of HCl, HNO3, and HF digestions (either hot plate or microwave (CEM, MARS
5000)). These two methods are generally analytically indistinguishable (Tables 2-4). For Re
and Cd, there appears to be some loss in samples digested via the hot plate method
(presumably during combustion in an 800oC furnace before acid digestion); only microwave
digestions are reported for Re and Cd (Tables 3 and 4).
Major element compositions (Ca, Fe, Mn, and Ti) were measured on total sample
digestions by inductively-coupled optical emission spectrometry (ICP-OES; Teledyne
Leeman Prodigy) (Table 2). For the same bulk sediment sample digestions, trace metal
concentrations (V, Mo, Cd, Re, U) were determined by inductively-coupled plasma mass
spectrometry (ICPMS; Thermo PQ ExCell) (Table 3). The reproducibility of analytical
techniques was evaluated by performing replicate analyses of multiple standard reference
materials (Table 4). Major element concentrations analyzed by ICP-OES for all standard
reference materials are typically reproducible within ~5%, and agree reasonably with
published values (Table 4). Trace metal concentrations determined by ICPMS in the
standard reference materials were typically reproducible within ~10% and agree with
reported values (Table 4). Replicate digestions and ICP-OES and ICPMS analyses were
performed on ~15% of all natural sediment samples in this study (Tables 2 and 3). The
average reproducibility for most elements typically is better than ~10%, which is consistent
with our other measures of reproducibility. However, it is worth noting that the total
measured sediment concentrations of Mo, Cd, and Re are at the low end of our typical
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analytical range, adding to the uncertainty of measured values (averaging ~20% for these
metals). For example, the average standard deviation on replicate Cd analyses is only 0.01
ppm, but on average this error represents ~35% of the total sample Cd concentration (Table
3).
Additional data for these sediments was supplied by Silke Severmann (University of
California - Riverside). Calcium concentrations were determined on carbonate extractions
using sodium acetate-acetic acid buffer (pH 4.5; Tessier et al., 1979; Table 2). Total carbon
(TC) and total inorganic carbon (TIC) were measured using an elemental analyzer (Eltra CS500). For TIC, acidification was done online (20% HCl); total organic carbon (TOC) was
calculated by difference (Table 2). Total reduced inorganic sulfur (TRIS) was measured by
single-step chrome reduction (Fossing and Jørgensen, 1989; Table 2).
RESULTS & DISCUSSION
Age Model
Sediment age models were generated from 210Pb- and 137Cs-derived sedimentation
rates and confirmed by correlation of all sediment calcium carbonate records (Figures 3 and
4). To derive linear sedimentation rates (LSR) from the 137Cs data, the depth of maximum
137
Cs was estimated in each sediment profile, and assigned an age of 1961 (Putyrskaya and
Klemt, 2007; orange dashed line in Figure 3; Table 1). All 137Cs-LSRs were then calculated
as 137Csmax depth (cm)/ 45 yr (2006-1961) (Table 1). The 1961 date represents the average
date of origin for two 137Cs maxima reported in lake sediments from the Southern Alps
(Putyrskaya and Klemt, 2007). There is uncertainty associated with the application of
northern hemisphere 137Cs timing to the southern hemisphere sites of this study; however, it
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appears to be a reasonable estimate for the 137Cs maxima observed in the Tanganyika
sediments. Only one sediment interval from the 232 m site contained detectable levels of
137
Cs (Figure 3); the 137Cs-LSR from an estimated ~2 cm 137Cs maximum is therefore poorly
constrained at best.
The depth of the modern sediment mixed layer was estimated from near-surface
perturbations in 210Pbxs and 137Cs profiles (black dashed lines in Figure 3). When possible, a
single best fit ln 210Pbxs data regression was used to calculate sedimentation rates below the
mixed layer (e.g. Nittrouer et al., 1979; Figure 3, Table 1). For the 335 m site, the average of
two separate 210Pb-derived sedimentation rates determined above and below an apparent
discontinuity at ~3 cm depth was used (Figure 3, Table 1). The 210Pb profiles from the 72 m
and 107 m sites also suggest a possible disruption in sedimentation at ~6-8 cm depth, but the
data are interpreted herein with a single regression (Figure 3, Table 1). Due to low 210Pb
activities in the 232 m core, the 210Pb-LSR for this site was calculated from only the
uppermost ~3 cm of data (Figure 3), and the estimated sedimentation rate for this site
therefore has considerable associated uncertainty.
For each site, the 137Cs- and 210Pb-derived LSRs were averaged and a single
sedimentation rate was applied to the entire core length (Figure 3, Table 1). There is limited
data available for comparison, but McKee et al. (2005) report 210Pb-derived sedimentation
rates from two nearby river deltas that are generally consistent with our estimates. These
authors report sedimentation rates of ~0.15cm/yr for two shallow (<100 m) cores taken in
the Gombe and Mwamgongo River deltas, located to the north of our study area. McKee et
al. (2005) suggest that sedimentation rates have increased in recent decades (reporting a
value of 0.25 cm/yr for recent sediments in the Mwamgongo delta), perhaps due to increased
rainfall in the early 1960s. Division of the 210Pbxs profiles in this study into shorter sediment
125
intervals could provide a more detailed history of temporal changes in sedimentation.
However, application of a single average sedimentation rate to each full core length
reasonably aligns the sedimentary calcium carbonate records from all sites (Figure 4),
suggesting the estimated sedimentation rates are an adequate approximation of
sedimentation over the core histories analyzed for the purposes of this study despite the
assumption of constant compaction.
Regional Sedimentation
Bulk sediment mass accumulation rates (MARs) were calculated as: MAR
(mg/cm2yr) = [LSR (cm/yr) * (1-ϕ) * ρ * 1000] using the average 137Cs- and 210Pb-derived
LSRs, average porosities (ϕ) estimated from wet and dry sediment weights, and the densities
of bulk sediment solids (ρ) calculated from the relative fractions of principal sedimentary
components (after Davis et al., 1999; Table 2). Sedimentary carbonate fractions were
calculated from both solid-phase Ca and TIC concentrations assuming CaCO3 stoichiometry
(CaCO3 = [Ca] * (100/40); CaCO3 = [TIC] * (100/12)). The average value from these two
separate estimates was taken to represent the total sediment carbonate fraction (XCARB) and
assigned an average density of 2.71 g/cm3 (Carmichael, 1982; Table 2). Similarly, the
sediment Corg content was used to estimate the bulk sediment organic matter fraction (XOM):
OM (wt. %) = Corg (wt. %) * (1.7); the value of 1.7 used approximates the typical H, C, and
O content of organic matter (e.g. Caplan and Bustin, 1996). The organic fraction was
assigned a density of 1.0 g/cm3 (Carmichael, 1982; Table 2). The remaining sediment (1 XCARB - XOM) was assumed to represent the detrital component of lithogenic materials
(XLITH), with an average density of 2.65 g/cm3 (Carmichael, 1982; Table 2). Bulk sediment
densities were then calculated as: Bulk ρ = (XCARB*(2.71 g/cm3)) + (XOM*(1.0 g/cm3)) +
126
(XLITH*(2.65 g/cm3)). This estimate ignores a sedimentary biogenic silica component which
cannot be constrained with the available data. Nevertheless, the lithogenic sedimentary
fractions calculated from the carbonate and organic matter estimates co-vary with measured
sediment Ti concentrations, suggesting these are reasonable approximations of the bulk
sedimentary components (Figure 5).
Average whole-core bulk sediment mass accumulation rates are highest in the
shallow 72 m site and decrease with site depth (Figure 6, Table 2). Not surprisingly, average
carbonate and Ti mass accumulation rates also generally decrease with site depth (Figure 6).
In all but the shallowest (72 m) site, carbonate accumulation rates are lowest (<5 mg/cm2yr)
over the last 50 yrs (Figure 6). Average carbonate accumulation rates during the latter part of
the Little Ice Age (c. 1550-1850 AD) are higher than modern rates, but average carbonate
accumulation rates are highest at all core locations during the 100 year period following the
Little Ice Age (Figure 6). This period (1850-1950 AD) includes the most recent maximum
lake high stand (c. 1878 AD; Alin and Cohen, 2003). The trend in Ti mass accumulation
rates (taken to represent terrigenous inputs) is opposite that of carbonate, with the highest
accumulation rates in the modern (1950 to present) and Little Ice Age periods, and the
lowest rates in between (Figure 6). The peak in carbonate sedimentation appears to be
centered ~1870 AD (Figure 4), which is roughly coincident with the termination of the Little
Ice Age and the most recent lake high stand.
Sediment Distributions of Diagenetic Reactants
Core average sediment organic carbon concentrations increase with site depth, from
~3% at the 72 m site, to ~6% at 335 m (Figure 7, Table 2). Average total reduced inorganic
sulfur (TRIS, presumably produced by sulfate reduction) also increases with site depth, from
127
~0.1% at the 72 m site, to ~1% at the deepest sites (Figure 7, Table 2). Despite similar trends
in total sediment concentration, the accumulation patterns for these two sediment
components differ along the transect (Figure 7). Organic carbon accumulation is highest in
the shallowest site (~0.14 mg/cm2yr), and burial rates generally decrease with site depth
(Figure 7). In contrast, TRIS accumulation is lowest at the 72m site (0.06 mg/cm2yr) but
averages ~0.2 mg/cm2yr for all other sites (Figure 7).
Berner and Raiswell (1984) argued that sediment C/S ratios can be used to
differentiate between sediments deposited in marine and freshwater environments, because
sulfur limitation in lacustrine settings drives sediment C/S ratios to high values. Indeed,
sediment C/S ratios (Corg/TRIS) for the shallowest study site are high (22±8), consistent with
reported values for freshwater systems (Berner and Raiswell, 1984; Figures 7 and 8). The
high average C/S ratio for the 72 m site likely reflects the combined effects of higher oxygen
concentrations (>30 μM; Edmond et al., 1993) at this depth inhibiting sulfide formation, and
the high rate of Corg burial at this site (Figure 7).
Whole core average C/S ratios for all other sites decrease with site depth, however,
approaching values more consistent with marine C/S ratios (average C/S = 6±2, n = 35;
marine C/S = 2.8±1.5, Berner and Raiswell, 1984; Figures 7 and 8). In fact, whole core
average C/S ratios for all sites generally trend along a slope similar to the reported marine
relationship, but the regression has a non-zero intercept (Figure 8). This non-zero intercept
implies higher carbon burial relative to sulfur compared to marine systems (Figures 7 and 8).
The average sediment C/S in the 335 m site may reflect the increasing impact of sulfur
limitation with depth; however, given the relatively large associated errors, it is consistent
with the general trend (Figure 8). All average C/S ratios reflect sulfur limitation in this
system, but the low sediment C/S ratios measured in all but the shallowest site suggest a
128
similar relationship between sulfate reduction and organic carbon oxidation for all the deeper
sites (Figures 7 and 8).
Sediment Fe and Mn contents were also measured as potential indicators of the
dominant sedimentary redox conditions, as both Fe(III) and Mn(IV) oxides are known
electron acceptors for organic carbon degradation (e.g., Froelich et al., 1979). Sediment
Fe:Ti and Mn:Ti ratios do not appear to be affected by changes in the sedimentary carbonate
fraction, suggesting these metals are dominantly associated with the terrigenous detrital
fraction (Table 2). Sediment Mn:Ti ratios generally decrease with sediment depth at all sites,
and whole core average Mn:Ti ratios also decrease with site depth across the transect (Figure
9, Table 2). Samples of the regional bedrock were not analyzed in this work, but others have
characterized the dominant bedrock lithology as Proterozoic metasedimentary and late
Paleozoic-early Mesozoic non-marine sedimentary rocks (Cohen et al., 1997). Mn:Ti ratios
in samples from the shallowest site are consistent with reported Proterozoic sedimentary
rock values (average Mn:Ti ~0.15; Taylor and McLennan, 1985 and references therein),
suggesting a detrital source. Lower Mn:Ti ratios in deeper sites suggest Mn reduction is an
active process, both within the sediments (particularly in the deepest sites) and in the
overlying water column (e.g. Klinkhammer and Bender, 1980). Pore water Mn profiles
provide further evidence of Mn reduction, with apparent recycling (loss) of Mn to pore
waters and deep lake waters (Figure 10).
Sediment Fe:Ti ratios are relatively invariant at all sites, averaging 11.9±0.5 for all
samples analyzed (1-SD, n=52; Figure 9; Table 2). Whole core average Fe:Ti ratios are
consistent with, but slightly higher than, the reported values for Proterozoic fine-grained
sedimentary rocks (average Fe:Ti ~9; Taylor and McLennan, 1985 and references therein).
The apparent invariance of measured Fe:Ti ratios suggests the sedimentary Fe pool is likely
129
dominated by detrital Fe, with no evidence of Fe reduction occurring within the sediment
column. However, total reduced inorganic sulfur (TRIS) in these sediments increases with
site depth, reflecting the increasingly reducing character of the sediments with depth (e.g.
Raiswell et al., 1988; Table 2). Pore water profiles suggest Fe reduction may be an active
process in these sediments, particularly below ~10 cm depth in the sediment column (Figure
10). The observed increase in sedimentary reduced sulfur without an observable change in
sediment Fe:Ti ratios suggests that if detrital Fe is reduced within the sediment column, it is
likely precipitated with sulfide as pyrite. Unlike Mn, the data suggest reduced Fe is not lost
to the overlying water column through diffusion.
Sediment Trace Metal Distributions
To quantify the authigenic enrichments of trace metals along the study transect, it is
necessary to consider the multiple sources potentially contributing these metals to the bulk
sediment pool. The fraction of sedimentary carbonate varies greatly over the depth ranges
analyzed at all sites (from <5% to ~70%, Figure 4, Table 2); any metal incorporation into
carbonate could potentially affect sediment metal contents. These metals may also be
contributed to the lake sediments in association with terrigenous detrital material delivered
by riverine or eolian inputs. To better constrain the authigenic metal fraction, the carbonate
and detrital metal contributions were estimated from sediment carbonate and Ti
concentrations ("predicted" metal concentrations in Figure 11). Sediment metal
concentrations in excess of the predicted carbonate and lithogenic contributions are
considered the "enriched" fraction, and are taken to primarily represent the authigenic
deposition of these metals (XENRICH, Table 4). However, some metals are considered
biologically essential, or form known associations with organic matter (e.g. Tribovillard et
130
al., 2004; Lane et al., 2005; Mendel and Bittner, 2006), and additional contributions
associated with organic matter deposition may impact sediment metal distributions. The
estimated metal enrichments therefore also potentially include metals directly associated
with organic matter, but no attempt was made to separately estimate "organic" metal
contributions.
Sediment Cd:Ti ratios also do not appear to be affected by changes in the carbonate
fraction, but Metal:Ti ratios for the remaining suite of trace metals (V, Mo, Cd, and Re) all
appear to be influenced (to varying degrees) by carbonate inputs; coincident maxima in
Metal:Ti and carbonate profiles suggest some contribution of carbonate-associated metals to
the total sedimentary metal pool. Carbonate metal contents taken from the literature were
used to estimate the carbonate fraction of the total bulk sediment metal concentration
(XCARB; Table 3). Average carbonate metal concentrations for V (20 ppm), Mo (0.4 ppm),
and U (2.2 ppm) were taken from Turekian and Wedepohl (1961). These authors did not
report a value for carbonate Re; instead, 1.2 ppb Re reported for whole rock limestone was
used (Pierson-Wickmann et al., 2000). Estimated carbonate metal contributions typically
represent <30% of the bulk sedimentary metal pool, though higher metal-carbonate fractions
are estimated for the shallow 72 m site (Table 3).
Average Metal:Ti values for Proterozoic sedimentary rocks were used to estimate
the sediment detrital metal fractions of V, Mo, and Cd at our study sites (XLITH, Table 3;
V:Ti = 0.027; Mo:Ti = 1.4x10-4; Cd:Ti = 1.1x10-5; Taylor and McLennan, 1985 and
references therein). For these same Proterozoic formations, the average reported U
concentrations (4-6 ppm) are higher than values measured in this study (average U = 2.4±0.6
ppm, n=52; Table 3) and no Re data are reported (Taylor and McLennan, 1985). The lowest
measured U:Ti and Re:Ti ratios from this study (U:Ti = 1.1x10-4; Re:Ti = 1.6x10-7) were
131
used to estimate the lithogenic fraction of these metals. The estimated lithogenic metal
fractions typically represent about half of the total sedimentary metal pool (Table 3).
The combined estimated carbonate and lithogenic contributions of each metal are
shown in Figure 11 ("predicted" values). Predicted metal concentrations agree reasonably
well with measured values for all metals in the 72 m site, suggesting little or no
distinguishable authigenic enrichment of these metals at this shallow depth (Figure 11). The
estimates reasonably predict the entire measured bulk sediment V pool at all site depths
(Figure 11, Table 3), suggesting V is primarily associated with the detrital sedimentary
component; there appears to be little or no authigenic vanadium enrichment in any of the
study locations. There are differences in the specific behavior of the other measured trace
metals, but in all cases the difference between predicted and measured metal concentrations
(the presumed "enriched" fraction, Table 3) is greatest at the deepest site (Figure 11). It is
worth noting here that the "predicted" Re concentrations rely on poorly constrained
carbonate and lithogenic Re values; we are therefore hesitant to further interpret the small
(~1 ppb) estimated "enriched" Re fractions (Figure 11).
Estimated authigenic Cd enrichments are low for all sites in this study; the
maximum average estimated enrichment is only ~0.06 ppm in the deepest site (Figure 11,
Table 3). Nevertheless, the estimates suggest some amount of Cd enrichment at all depths,
increasing below the chemocline (Figure 11). Cadmium exists in a single coordination state
(Cd2+), and is rapidly immobilized in sediments under suboxic conditions (e.g. McCorkle
and Klinkhammer, 1991). Authigenic Cd enrichment is thought to be primarily achieved
through CdS formation; Cd is efficiently immobilized in sediments in the presence of trace
levels of free sulfide (<5 μM), though remobilization of Cd upon subsequent reoxidation of
sediments has been suggested (e.g. van Geen et al., 1995; Rosenthal et al. 1995; and
132
references therein). However, Cd is also known to form associations with organic matter;
recent work has identified a Cd-containing carbonic anhydrase in marine diatoms (Lane et
al., 2005). Association of Cd with biological material is further suggested by the relatively
invariant Cd:Corg ratio reported for particulate marine organic matter (~3x10-6; Rosenthal et
al., 1995). In light of this reported Cd:Corg relationship, delivery of Cd to the sediments in
association with organic carbon can adequately account for the entire measured bulk Cd pool
in the Tanganyika sediments. Based on the available data, it is not possible to discern
whether the estimated "enriched" Cd fraction in these sediments reflects authigenic CdS
precipitation or delivery of Cd associated with organic matter; in truth, both processes likely
impact the observed Cd distributions.
Authigenic Accumulation of U and Mo
Authigenic accumulation of U is primarily achieved through the reduction of soluble
U(VI) to the more insoluble U(IV), at sediment Eh conditions similar to those required for
Fe(III) reduction (e.g. Anderson et al., 1989; Klinkhammer and Palmer, 1991). The exact
mechanism for U reduction remains unclear; microbial reduction of U has been observed
(e.g. Lovley et al., 1991), but abiotic reduction of U linked to sulfate reduction has also been
proposed (Klinkhammer and Palmer, 1991). In addition, some amount of U may be
delivered to the sediments through sorption to organic matter in the water column (e.g.
Anderson et al., 1989; Klinkhammer and Palmer, 1991; Zheng et al., 2002a). Maximum
authigenic U accumulation appears to be taking place in the Tanganyika sediments just
above the chemocline depth; calculated accumulation rates are slightly lower in the deepest
sites (Figure 12). However, given the uncertainties associated with these estimates, enriched
133
U accumulation rates for all sites below 72 m may be essentially the same; only the lowest
accumulation rate estimated for the shallowest location is markedly different (Figure 12).
High C/S ratios measured in the shallow (72 m) site sediments suggest a suppression
of sulfate reduction, presumably due to the higher bottom water oxygen concentrations at
this depth (Figure 8). Assuming a link between sulfate and U reduction, this may explain the
low U accumulation rate at the 72 m site relative to deeper locations (Figure 12). However, it
is also possible that some amount of U may be reoxidized in these sediments, such that the
low U accumulation reflects loss of U from the sediments in the oxygenated shallow site
(Zheng et al., 2002b). The linear relationship between core average Uenrich and TRIS across
all sites on the transect further suggests that the reduction of U is primarily tied to sulfate
reduction in this environment (Figure 13). A similar linear relationship between core average
Uenrich and Corg is suggested for all but the shallowest site (Figure 13). As stated previously,
the low Uenrich value for the 72 m site likely reflects lower rates of sulfate reduction relative
to the high Corg burial, as well as some potential loss of U to overlying water due to
reoxidation. The Uenrich/Corg burial ratio observed in the deeper sites (0.36 Uenrich ppm: Corg %;
Figure 13) is about one-third of the values observed in marine continental margin
environments (~1.2 Uauth ppm: Corg %; McManus et al., 2006). Nevertheless, the positive
correlation between U and C burial in Lake Tanganyika sediments suggests similar
preservation of the diagenetic relationship between authigenic U accumulation, sulfate
reduction, and organic carbon degradation in this freshwater system.
Molybdenum is considered a biologically essential element (e.g. Mendel and Bittner,
2006 and references therein), and may be delivered directly to the lake sediments in
association with organic matter. Therefore some amount of the estimated "enriched" Mo
may be of biogenic origin. However, because the estimated Moenrich concentrations are near
134
zero for the three shallowest sites, the relationships between core average Moenrich and Corg
(or TRIS) along the transect are not easily constrained (Figure 13).
Authigenic Mo accumulation in reducing environments is primarily associated with
the transformation of the soluble molybdate ion (MoO42-) to less soluble thiomolybdates
(MoOxS4-x2-) that may be scavenged by sulfidized organic matter or Fe-sulfide phases such
as pyrite (Helz et al., 1996; Zheng et al., 2000). Experimental work has shown that, in the
presence of both H2S and S0- electron donors, thiomolybdate Mo(VI) may be reduced to
Mo(V) or Mo(IV) polysulfide anions (Vorlicek et al., 2004); it remains unclear whether
reduction of the metal itself is necessary for authigenic Mo accumulation. Helz et al. (1996)
proposed a sulfide-controlled geochemical “switch” for Mo at ~10 μM H2S(aq), where the
dominant dissolved Mo phase abruptly transitions from molybdate (MoO42-) to
tetrathiomolybdate (MoS42). The marine pore water work of Zheng et al. (2000) proposed
two thresholds for Mo-sulfide formation; at H2S(aq) concentrations of ~0.01 μM these authors
proposed that Mo is removed from solution via coprecipitation of Fe-Mo-S phases, whereas
at higher H2S(aq) concentrations (~10 μM) they postulate that Mo precipitates independent of
iron.
Estimates of Mo enrichment in the Tanganyika sediments suggest authigenic Mo
deposition is inhibited above the chemocline depth (Figures 11 and 12). The estimated
authigenic Mo accumulation rates are essentially zero for the three shallowest sites, but
authigenic Mo appears to be accumulating in the sediments of the two deepest sites (Figure
12). The observed pattern of Mo accumulation suggests that sulfide concentrations sufficient
for authigenic Mo-sulfide precipitation are limited to sediments below the chemocline depth,
where oxygen is no longer detectable, and free sulfide >10 μM persists in the water column
(Figure 2, Edmond et al., 1993).
135
CONCLUSIONS
The sediment C/S ratios for all study sites are consistent with these sediments
generally being sulfur limited, as would be expected in this freshwater system. At the
shallowest site, C/S ratios agree with reported freshwater values; however, the deeper sites
exhibit lower C/S ratios. The low C/S ratios observed in deeper sediments likely reflect the
importance of sulfate reduction for organic matter degradation in this system. Accumulation
of authigenic U occurs at all sites along the study transect, while significant authigenic Mo
accumulation is limited to sites below the chemocline. These metal distributions likely
reflect the different geochemical mechanisms responsible for enrichment, and the
relationship of these metals to organic carbon delivery and degradation within the sediments.
Sediment trace metal distributions, along with the observed C/S ratios, imply sulfate
reduction and authigenic metal enrichment are suppressed by higher oxygen availability in
the shallowest site relative to the deeper locations. Conditions within the sediments become
increasingly reducing along the depth transect; authigenic metal enrichments and sediment
reduced sulfur contents increase with site depth, while C/S ratios decrease well below
reported freshwater values. Despite sulfur limitation in this lacustrine system, there is
evidence of substantial reduced sulfur accumulation and authigenic metal enrichment within
the lake sediments, particularly at depths below the chemocline.
136
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Zheng, Y., Anderson, R.F., van Geen, A., and Kuwabara, J. (2000) Authigenic molybdenum
formation in marine sediments: A link to pore water sulfide in the Santa Barbara
Basin. Geochim. Cosmochim. Acta 64, 4165-4178.
139
Figure 1. Map of study area: Lake Tanganyika, Tanzania. Inset shows chemocline
transect on Luiche Platform. Base map generated using Online Map Creator at
www.planiglobe.com.
Figure 2. Water column profiles from Northern Lake Tanganyika, Kigoma Basin; all data from
Edmond et al. (1993). Grey dashed line indicates approximate depth of chemocline (150 m).
140
141
Figure 3. Down-core profiles of 210Pb and 137Cs used to create age models for all Luiche
Platform sites. 137Cs data plotted in orange diamonds, 210Pb data included in sedimentation
rate regressions shown in grey squares (additional data 210Pb shown in black circles). Black
dashed lines indicate estimated depths of sediment mixing, orange dashed lines indicate
estimated depths of 137Cs maxima. 137Cs-derived linear sedimentation rates in orange, 210Pbderived linear sedimentation rates in black. Linear sedimentation rates outlined in boxes are
average values used in age models.
142
143
Figure 4. Down-core profiles of sediment carbonate contents for all Luiche Platform
cores. Left panel: carbonate versus depth in the sediment column as measured. Right panel:
profiles set to ages from site-specific 210Pb- and 137Cs-derived linear sedimentation rates.
Figure 5. Sediment Ti concentrations (%) versus estimated sedimentary lithogenic
fractions (XLITH) for all study sites. Lithogenic sediment fractions estimated from % total
sediment in excess of carbonate and organic matter fractions.
144
Figure 6. Average total sediment, carbonate, and Ti mass accumulation rates for all
study sites. Errors are 1-SD for average values.
Figure 7. Whole core average organic carbon (Corg) and total reduced inorganic sulfur (TRIS), Corg and TRIS
accumulation rates, and sediment C/S (Corg /TRIS) ratios for all study sites. All errors are 1-SD for whole core averages.
Marine and freshwater C/S values from Berner and Raiswell, 1984).
145
Figure 8. Individual sample (left panel) and whole core average (right panel) sediment organic carbon (Corg) versus total
reduced iron sulfide (TRIS) for all study sites. Errors are 1-SD for whole core averages. Marine and freshwater C/S values from
Berner and Raiswell, 1984).
146
147
Figure 9. Whole core average sediment Mn:Ti and Fe:Ti ratios for all study sites. All
errors on Mn:Ti and Fe:Ti ratios are 1-SD for whole core averages.
Figure 10. Pore water Mn and Fe profiles for all study sites. OLW indicates values
measured in the overlying water samples.
averages. "Predicted" values are estimated metal contributions from carbonate and terrigenous detrital inputs. All
sediment "enriched" metal concentrations estimated from difference between predicted and measured (see text).
Figure 11. Whole core average predicted and measured trace metal concentrations. Errors are 1-SD for whole core
148
149
Figure 12. Whole core average "enriched" U and Mo accumulation rates. Error bars are
propagated from 1-SD uncertainties in whole core averages of the two terms (MetalENRICH
and MAR). Dashed line indicates chemocline depth (~150 m).
Figure 13. Whole core average "enriched" U and Mo versus whole core average
organic carbon (%Corg) and total reduced inorganic sulfur (%TRIS). All errors are 1-
SD for whole core averages.
Table 1. Water depths, locations, and sedimentation data for all study sites.
See text for detailed description of linear sedimentation rate (LSR) and sediment mass accumulation rate (MAR) calculations.
150
151
Table 2. Sediment data: major elements, carbon and sulfur contents, carbonate
fractions, estimated densities, and mass accumulation rates. All reported values are from
separate total sediment digestions. Errors (1-SD) listed when more than one replicate
digestion was performed. Sediment total inorganic carbon (TIC), total organic carbon
(TOC), and total reduced inorganic sulfur (TRIS) data from Silke Severmann (UCRiverside). See text for detailed description of carbonate, density, and mass accumulation
rate (MAR) calculations.
152
Table 2.
153
Table 2 (Continued)
154
Table 2 (Continued)
155
Table 3. Sediment trace metal data. All reported values are from separate total sediment
digestions. Errors (1-SD) listed when more than one replicate digestion was performed.
Carbonate (XCARB) and lithogenic metal contributions (XLITH) estimated from published
values (see text). Estimated sediment "enriched" metal fractions (XENRICH) calculated from
difference of measured and predicted metal concentrations (see text for details).
156
Table 3 (Continued)
157
Table 3 (Continued)
Table 4. Average standard reference material compositions. Average values are from separate replicate digestions (n=# of
samples). For all but Cd and Re, averages represent mix of hot plate and microwave digestion techniques; reported Cd and Re
averages from microwave digests only (see text).
158
159
CONCLUSION
The data from this thesis suggest that early diagenetic processes can generate Mo
isotope compositions in marine sediments that span the full range of values previously
observed in natural environments. There appear to be discrete isotopic signatures for endmember oxic (~-0.5‰) and anoxic (~1.6‰) authigenic Mo deposits; however, these
signatures can be obscured by additional sedimentary Mo contributions. Notably lithogenic
material (~0‰), while typically representing only a small portion of the total sediment bulk
Mo pool, can effectively "dilute" the authigenic signal. In addition, heavier sediment Mo
isotope compositions (less fractionated relative to a parent seawater δ98Mo source) measured
in the surface sediments of many sites suggest that an additional Mo component may be
controlling Mo isotope values near the sediment-water interface. I propose that Mo
associated with organic matter is a discrete source of Mo to marine sediments that is less
fractionated (relative to a seawater source) than authigenic phases, and that biogenic Mo
dominates the sedimentary Mo pool in surface sediments at many locations. Data presented
here further suggest that when Mn reduction is an important source of aqueous Mo within
the sediment column, it is reflected in the isotopic composition of the authigenic sediment
fraction. It appears that different sources of aqueous Mo (seawater [δ98MoSW] versus Mo
released from Mn-oxides [δ98MoMn]) can generate very different Mo isotopic compositions
in the precipitated authigenic phases. Bulk marine sediment δ98Mo values therefore represent
the mass balance of multiple primary Mo source components with an observed range over all
study sites of -0.8 to +2.3‰.
Trace metal concentrations in lake sediments were also investigated in this work, to
investigate whether metal limitation in a freshwater system inhibits the formation of
160
authigenic metal enrichments as observed in marine environments. Sediment geochemical
data from a depth transect in Lake Tanganyika, East Africa, were investigated to characterize
the burial response of these elements to changes in sediment reducing character across a
chemocline transition. Accumulation of authigenic U is suggested at all sites along the study
transect, while significant authigenic Mo accumulation appears to be limited to sites below
the chemocline. These metal distributions likely reflect the different geochemical
mechanisms responsible for enrichment, and the relationship of these metals to organic
carbon delivery and degradation within the sediments. Despite sulfur limitation in this
lacustrine system, there is evidence of substantial reduced sulfur accumulation and
authigenic metal enrichment within the lake sediments, particularly at depths below the
chemocline.
161
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APPENDIX A: ADDITIONAL DATA TABLES FOR CHAPTER 3
Appendix Table 1. Standard reference materials for Mo, Al, Ca, Fe, Mn, and Ti. Average values are from separate replicate
digestions (n=# of samples); mix of hot plate and microwave digestion techniques (see text).
171
172
Appendix Table 2. Sediment Mo concentration data. All reported values are from
separate total sediment digestions (analyzed by either ICPMS or MC-ICPMS as indicated).
Errors on Average Mo concentrations are 1-SD for all analyses. For some samples, aliquots
of the same sediment digestion were analyzed by both methods; the average of these two
analyses (and 1-SD error) are listed under "Both" so as not to place undue weight on a single
digestion when calculating the final sample Mo concentration average. Samples listed in
bold font were processed by microwave digestion; all other samples processed by hot plate
digestion.
173
Appendix Table 2.
174
Appendix Table 2 (Continued)
175
Appendix Table 2 (Continued)
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Appendix Table 2 (Continued)
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Appendix Table 2 (Continued)
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Appendix Table 2 (Continued)
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Appendix Table 2 (Continued)
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Appendix Table 2 (Continued)
181
Appendix Table 3. Sediment major element compositions and lithogenic Mo fractions.
All reported values are from separate total sediment digestions. Errors (1-SD) listed when
more than one replicate digestion was performed. Al, Ca, Fe, and Mn for MANOP H
(Vulcan BC37, same core) and MANOP M (Pluto 25BC, companion to Mo core) from Lyle
et al. (1984); data not corrected for carbonate dilution. Samples listed in bold font were
processed by microwave digestion; all other samples processed by hot plate digestion.
Appendix Table 3.
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Appendix Table 3 (Continued)
183
Appendix Table 3 (Continued)
184
Appendix Table 3 (Continued)
185
Appendix Table 3 (Continued)
186
Appendix Table 3 (Continued)
187
Appendix Table 3 (Continued)
188
Appendix Table 3 (Continued)
189
Appendix Table 3 (Continued)
190
Appendix Table 3 (Continued)
191
192
Appendix Table 4. Sediment Mo isotope compositions. All reported values are from
separate total sediment digestions. Errors on Average values are 1-SD for all analyses; errors
reported for individual Mo isotope analyses are 2SE instrumental errors from individual
runs.
193
Appendix Table 4 (Continued)
194
Appendix Table 4 (Continued)
195
Appendix Table 4 (Continued)
196
Appendix Table 5. Average sediment Moenrich concentrations and isotopic
compositions. Average bulk sediment Mo concentrations from Appendix Table 2, average
bulk Mo isotope compositions from Appendix Table 3, average %Al from Appendix Table
4. Assumed Mo:Al ratio for lithogenic background is 1.1 x 10-5 (McManus et al., 2006;
Poulson et al., 2006). %Al (Appendix Table 4) and %Corg values for MANOP H (Vulcan
BC37, same core) and MANOP M (Pluto 25BC, companion to Mo core) from Lyle et al.
(1984); data not corrected for carbonate dilution. %Corg value for Peru is average from
McManus et al., (2006); companion core. MoLITH and XLITH values in italics do not have
corresponding %Al data; reported values calculated from core average %Al values.
197
Appendix Table 5.
198
Appendix Table 5 (Continued)
199
Appendix Table 5 (Continued)
200
Appendix Table 5 (Continued)
201
Appendix Table 6. Detailed sediment data from previously published California and
Mexico margin sites. All previously published data listed in italics (Siebert et al., 2006;
Poulson et al., 2006); all other data are from replicate recent sample digestions. All Mo
values listed represent separate total sediment digestions (analyzed by either ICPMS or MCICPMS as indicated). Errors on Average Mo concentrations are 1-SD for all analyses. For
some samples, aliquots of the same sediment digestion were analyzed for Mo by both
methods; the average of these two analyses (and 1-SD error) are listed under "Both" so as
not to place undue weight on a single digestion when calculating the final sample Mo
concentration average. All Mo isotope values reported are separate total sediment digestions.
Errors on Average values are 1-SD for all analyses; errors reported for individual Mo isotope
analyses are 2SE instrumental errors from individual runs. All Al, Ca, Fe, and Mn values are
replicate digestions and analyses of the same bulk sediment sample. Errors (1-SD) listed
when more than one replicate digestion was performed. Assumed Mo:Al ratio for lithogenic
background is 1.1 x 10-5 (McManus et al., 2006; Poulson et al., 2006).
202
Appendix Table 6.
203
Appendix Table 6 (Continued)
204
Appendix Table 6 (Continued)
205
Appendix Table 6 (Continued)
206
Appendix Table 6 (Continued)
207
Appendix Table 6 (Continued)
208
Appendix Table 6 (Continued)
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Appendix Table 6 (Continued)
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Appendix Table 6 (Continued)
211
Appendix Table 6 (Continued)