University of Southern California Dissertations and Theses

Tracking fluctuations in the eastern tropical north Pacific oxygen minimum zone: a high-resolution geochemical evaluation of laminated sediments along western North America

(USC Thesis Other)

Tracking fluctuations in the eastern tropical north Pacific oxygen minimum zone: a high-resolution geochemical evaluation of laminated sediments along western North America

PDF

Download a page range

Download transcript

Copy asset link

Request this asset

Transcript (if available)

Content

TRACKING FLUCTUATIONS IN THE EASTERN TROPICAL NORTH PACIFIC OXYGEN
MINIMUM ZONE: A HIGH-RESOLUTION GEOCHEMICAL EVALUATION OF
LAMINATED SEDIMENTS ALONG WESTERN NORTH AMERICA

By

Caitlin Elizabeth Tems

______________________________________________________________________________

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

May 2016

Copyright 2016 Caitlin Elizabeth Tems
ii

This dissertation is dedicated to my parents, Robin and Cindy Tems, and grandparents June
Tems and Gene Cloninger who have instilled ethics of hard work, integrity and humility,
encouraged scientific curiosity, and provided continuous and unwavering support.

iii
Acknowledgements
Many people have contributed to the success of this dissertation and I would like to first
and foremost acknowledge the guidance and support of my advisor Dr. Will Berelson.
Transitioning from a traditional field based geology background into the realm of marine
geochemistry was both an exciting and challenging endeavor. I learned how to be an independent
and competent marine scientist through Dr. Berelson’s guidance and expertise in both the lab and
field environment, which ranged from collecting deep water sediment cores in the eastern
tropical south Pacific to constructing novel experiments and investigating unresolved questions.
I would also like to thank Dr. Frank Corsetti and Dr. Doug Capone for their guidance as
members of my dissertation committee. With one of the original goals of investigating the
connection between modern low-oxygen depositional environments and analogs in the rock
record, Dr. Corsetti provided invaluable knowledge and advice which illuminated the importance
and complexity of studying the transition from sediments to rocks and deep time. Likewise, Dr.
Doug Capone’s expertise and insights about the nitrogen cycle were integral to this work and
enabled me to gain a comprehensive understanding of this critical cycle. I would also like to
thank and show great appreciation for Dr. Donn Gorsline, who was a member of my
qualification committee and originally an additional dissertation committee member. I feel
honored to have been able to learn from Dr. Gorsline whose unprecedented wealth of knowledge
and understanding of the California Borderland, excitement about marine geology, and
encouragement greatly contributed to my Ph.D. experience. I additionally would like to thank
Dr. Maria (Masha) Prokopenko, also a member of my qualification committee, for sharing her
knowledge of nitrogen isotopes and expertise conducting oceanographic measurements at sea.
Her insight and guidance encouraged me to grow as a scientist. I would like to thank Dr. Steve
iv
Lund for sharing paleomagnetic data from the Pescadero Slope cores and for productive
discussions about factors potentially involved in oxygen minimum zone fluctuations.
I would like to gratefully acknowledge the contributions and support from collaborators
outside of the University of Southern California, specifically Dr. Robert Thunell at University of
South Carolina, Dr. Alexander Van Geen at Lamont-Doherty Earth Observatory of Columbia
University, and Dr. Xiaomei Xu at the University of California Irvine. Collaborations with each
lab group were essential to the completion of the projects in this dissertation. Dr. Thunell and
Dr. Van Geen both shared their expertise on nitrogen isotopes and sediment geochemistry,
offered invaluable laboratory collaboration and support, and provided critical insight regarding
questions being investigated in this dissertation. Dr. Xu assisted with radiocarbon
measurements, which were essential for the construction of sediment core age models.
Additionally I would like to thank Eric Tappa (University of South Carolina) and Jennifer
Lehman (University of California Irvine) for their support in the lab.
Nick Rollins, the lab manager of the Berelson Lab, played an important role in my Ph.D.
research by teaching me to run the elemental analyzer and the Picarro in the lab, providing a
Matlab crash course and assistance whenever needed with the program, encouraging pull-up
contests while in the middle of the South Pacific, and creating a positive and laid back
atmosphere in the lab. I would also like to thank former lab-mate Dr. Tim Riedel for showing
me the ropes my first year at USC and providing valuable insights into the Ph.D. process which
applied throughout my graduate school experience. I would also like to acknowledge Dr. Laurie
Chong, John Fleming, Sijia Dong, and Adam Subhas who have been great to work with in the
lab and engage in ongoing discussions about ocean geochemistry. I would also like to thank
v
former Berelson Lab member Dr. Lisa Collins. While we did not overlap in lab, encouragement
and discussions with her have been incredibly helpful.
I would also like to thank Dr. Deborah Khider, Dr. Camilo Ponton, and Danielle
Monteverde, who through scientific discussions have contributed to my growth as a scientist.
One prospective project that was preliminarily investigated involved reconstructing
oxygen minimum zone fluctuations in the Monterey Formation. While this project was not able
to be included in this dissertation, I am incredibly grateful to USC alum, Dr. Jon Schwalbach,
who provided access to cores from the Monterey Formation that allowed me to investigate the
potential of this project.
I would also like to thank the USC Earth Sciences Administrative Staff, Cindy Waite,
John Yu, Barbara Grubb, Vardui Tersimon and Karen Young who have been very supportive
throughout my graduate school career and made the dissertation process infinitely easier.
My family and friends have been an integral part of my experience at USC. While I do
not have the space to individually thank everyone here, I would like to especially thank my
parents, Cindy and Robin, and brother, Chris, who provided continuous and unwavering support
throughout my Ph.D. experience. Thank you for everything from early morning and late night
phone calls around the world, escapes to the Arabian desert and the Colorado mountains, and
offers to edit my dissertation at 3 a.m. the day it needs to be submitted! Finally, I would like to
thank Leigh and Indigo, who have kept me grounded throughout this process by providing
unconditional support, while also encouraging endless rock climbing adventures, Sierra trips, and
soccer games.

vi
Table of Contents
Dedication ii
Acknowledgements iii
Chapter 1: Introduction
1.1 Motivation 1
1.2 The Biological Pump and δ
15
N as a Proxy of Changes in Strength of OMZs 2
1. 3 Study Sites 10
1.4 Objectives 13
References 15
Figures 25

Chapter 2: δ
15
N in laminated marine sediments provide a proxy for mixing between the
California Undercurrent and the California Current: A proof of concept
Abstract 30
2.1 Introduction 30
2.2. Methods 34
2.3. Results: Downcore δ
15
N
sed
record 35
2.4. Discussion: Factors driving fluctuations in the poleward transport of nitrate δ
15
N 38
2.5. Conclusions 42
Acknowledgements 43
References 44
Figures 53

Chapter 3: Sedimentary δ
15
N reveal decadal fluctuations in the intensity of the eastern
tropical north Pacific oxygen minimum zone during the last 1200 years
Abstract 61
3.1 Introduction 62
3.2 Study Site 64
3.3 Sampling and Analytical Methods
3.3.1 Core Collection 65
3.3.2 Age model 66
3.3.3 Sediment Chemistry Analytical Techniques 67
3.3.4 Spectral Analysis 71
3.4 Results
3.4.1 δ
15
N
sed
and Periodicity 71
3.4.2 δ
15
N
sed
, bSi, and C
org
73
3.5 Discussion 73
Conclusions 80
Acknowledgements 81
References 82
Figures and Tables 92

vii
Chapter 4: Assessing Regional Variability in the Eastern Tropical North Pacific Oxygen
Minimum Zone through a Geochemical Comparison of Pescadero Slope and
Soledad Basin Sediments
Abstract 105
4.1. Introduction 106
4.2. Study Sites 109
4.3 Methods
4.3.1 Core Collection 110
4.3.2 Age Model 112
4.3.3 Nitrogen Isotope Measurements 113
4.3.4 X-ray Florescence Analysis 114
4.4 Results
4.4.1 Soledad Basin δ
15
N
sed
and Periodicity 115
4.4.2 X-Ray Fluorescence Analysis 116
4.5 Discussion 118
4.6 Conclusions 124
Acknowledgements 126
References 127
Figures and Tables 136

Chapter 5: Summary and Conclusions 148
References 159

Bibliography 163

Appendices
Appendix A: Deutsch et al., 2014
Centennial changes in North Pacific anoxia linked to tropical trade winds 187

Appendix B: Data from Chapter 2
Pescadero Slope Porosity 192
Santa Monica Basin Porosity 192
Pescadero Slope Multicore
210
Pb Age Model 193
Pescadero Slope and Santa Monica Basin Multicore δ
15
N
sed
194
Santa Monica Basin Multicore δ
15
N
sed
and Salinity Comparison 197
Δδ
15
N
sed
and Salinity Comparison 198
Δδ
15
N
sed
and Pacific Decadal Oscillation Index (1900-2009) 199

Appendix C: Data from Chapters 3 and 4
Pescadero Slope Gravity Core Varve Counts 201
Pescadero Slope Multicore & Gravity Core δ
15
N
sed
205
Pescadero Slope Multicore & Gravity Core Weight Percent Organic Carbon 220
Pescadero Slope Multicore & Gravity Core Weight Percent Biogenic Silica 222
Information about Chapter 4 Data 223

Chapter 1
Introduction

1.1 Motivation

Dissolved oxygen (DO) concentrations in the ocean are constantly fluctuating and have
varied throughout geologic history. A degree of variation is natural, however, significant
variations, especially those in sub-surface waters, can greatly impact biogeochemical cycling and
marine ecosystems. Large variations in oceanic DO in the geologic past caused ocean anoxia
events (OAE) in the Cretaceous period [Jones and Jenkyns, 2001] and are believed to have
triggered major extinctions including the end-Permian event [Benton and Twitchett, 2003;
Wignall and Twitchett, 1996]. Recent studies, utilizing both field measurements and model
predictions show that significant declines in dissolved oxygen content are occurring in portions
of the ocean today [Levin and Bris, 2015; Getzlaff and Dietze et al., 2013; Deutsch et al., 2005;
Bindoff and McDougall, 2000; Emerson et al., 2001, Garcia et al., 1998; Matear et al., 2000;
Ono et al., 2001; Watanabe et al., 2001; Stramma et al., 2008; Bograd et al., 2008]. DO
fluctuations have been accurately monitored for approximately 60 years, making it critical to find
alternative ways to investigate longer term trends in DO variations. This dissertation investigates
the utility of geochemical proxies from sediments to explore how oxygen content in the ocean
has fluctuated in the past 1200 years and the factors influencing these variations.
The surface of the ocean is generally well mixed and will come to equilibrium
concentration with atmospheric oxygen. However, the subsurface waters, once removed from
ventilation can lose oxygen due to respiration. The dynamics of ocean circulation make some
regions more susceptible to changes in DO content than others. Eastern boundary current
systems, characterized by upwelling regimes are especially susceptible to perturbations in DO
because oxygen-depleted water in these systems can upwell onto the continental shelves [Diaz
1

and Rosenberg, 2008, Stramma et al., 2010], significantly influencing biogeochemical cycling
and negatively impacting the marine biota that live in these regions. This has recently occurred
on the Central Oregon Shelf and resulted in severe inner-shelf anoxia and large-scale hypoxia
over an area of >3000 km
2
[Chan et al., 2008]. This event resulted in the loss of all fish from
rocky reef habitats, near mortality of macroscopic benthic invertebrates, and an increase in the
presence of sulfide-oxidizing bacterial mats [Chan et al., 2008]. Declines in DO have also been
observed in the Southern California Current System. Analysis of quarterly California
Corporative Oceanic Fisheries Investigation (CalCOFI) data reveals that are significant declines
in DO (as high as 1.7 µM yr
-1
at 100 m) occurring in waters below the thermocline (300 m)
[Bograd et al., 2008]. Evidence also indicates the hypoxic boundary (defined by DO ≤ 60 µM)
has shoaled by ~90 meters likely due to the advection of low DO waters into the region and
increased stratification [Bograd et al., 2008]. Likewise, time series analysis of monthly San
Pedro Ocean Time Series (SPOT) measurements over the past 15 years (unpublished data) reveal
similar declines in DO on the order of 1.7 µM yr
-1
(Figure 1.1). A decline in rockfish population
related to declining DO concentrations in the water column has also been documented off the
Southern California coast [McClatchie et al., 2010].
Variation is DO content in the subsurface is expressed as the ‘intensity’ of oxygen
minimum zones (OMZs). OMZs are midwater features are associated with expansive, highly
productive regions of the ocean (typically eastern boundary upwelling regimes) and are sites of
microbially-mediated biogeochemical cycling of redox sensitive compounds [Gilly et al., 2013;
Keeling et al., 2010]. A 50-year time series monitoring dissolved oxygen concentrations reveals
that there has been an areal and vertical expansion of DO and oxygen minimum zones (OMZs) in
the North Indian, eastern tropical Atlantic and equatorial Pacific. These regions have been losing
2

oxygen at rates between 0.09 to 0.39 μmol/kg yr [Stramma et al., 2008]. The largest rates of
decline are found in the Atlantic Basin (0.17 to 0.39
μmol/kg yr) and the Pacific Basin (0.13-0.39
μmol/kg yr) (Figure 1.2). The rate of decline is less pronounced in the North Indian Basin (0.09
μmol/kg yr), however, still represents a significant change in DO content. In the last two to three
decades this global trend has resulted in a 10 to 20% decline in oxygen in the photic zone in
these areas [Stramma et al., 2008].
Significant variations in the intensity of OMZs can alter biogeochemical cycling
(especially denitrification), climate (through the release of N
2
O), ecological relationships, and
have important implications for ocean fisheries management. If the OMZ expands, it constrains
the habitat of larger, pelagic, predatory fish by decreasing the size of the surface layer [Prince
and Goodyear, 2006]. This decrease in habitable area may disrupt predator-prey relationships,
giving large fish more foraging opportunities and therefore altering trophic food web interactions
[Stramma et al., 2010]. If the OMZ continues to expand, restriction of these fish to shallower
depths has resulted in overfishing by both commercial and privately owned vessels due to their
proximity to the surface. Expanding hypoxia will result in increases in hypoxic tolerant
organisms, such as jelly plankton, certain squid and jellyfish. Jellyfish blooms have been linked
to eutrophication and hypoxia [Purcell et al., 2007; Purcell et al., 2001]. Increasing jellyfish
populations may promote hypoxia by preying on zooplankton, which would leave unconsumed
phytoplankton organic matter to sink and degrade slowly in the water column, resulting in
further DO declines at shallower depths [Stramma et al., 2010]. DO declines will not only affect
the marine biota, they will also affect biogeochemical processes in the sediments and the water
column. Hypoxia may alter the rates at which organic matter and metals are recycled in the
ocean [Anderson et al., 2008]. This affects the efficiency of the biological pump, reducing the
3

remineralization of high quality organic matter [Stramma et al., 2010]. If the ocean becomes
more stratified as a result of global warming, the expansion of OMZs poses a serious threat to
these environments [Keeling et al., 2010].
Due to limited data available to reconstruct DO in the ocean synoptically, and the
concern of the consequences of an expanding OMZ, Global Climate Models (GCMs) have been
used to predict how DO may change in the future due to increased greenhouse gas emissions.
The GCMs predict warming of the upper ocean and freshening at high latitudes, which will cause
density stratification of the ocean interior [Keeling and Garcia, 2002]. This enhanced
stratification will reduce ventilation, gas exchange between the ocean and the atmosphere, and
result in a decline in DO in the water column [Sarmiento et al., 1998]. Additional studies suggest
that a 1°C increase in water temperature will result in a 5
μM decline in DO in the upper ocean
which will lead to a 10% expansion of hypoxic conditions and 300% increase in the volume of
suboxic water [Deutsch et al., 2011]. Even a 1% drop in DO, equivalent to ~2 μM decline in
DO, could cause a doubling in size of the current suboxic zone.
Suboxic zones (or OMZs) are areas of significant fixed nitrogen loss through
denitrification (the microbially facilitated reduction of nitrate) and anaerobic ammonium
oxidation (anammox), which is the microbial reduction of nitrite and ammonium. This is why
small changes in DO concentrations will have large impacts on ocean nitrogen budgets.
Assessing the state of the nitrogen budget has been the subject of many studies and while current
investigations have suggested that nitrogen cycle inputs and losses are balanced [Cassiotti,
2016], its been hypothesized that changes over the past 100,000 to 1,000,000 years in the
nitrogen budget might be linked to observed changes in atmospheric carbon dioxide (CO
2
) and
4

nitrous oxide (N
2
O) [Altabet et al., 1995, 1999; Ganeshram et al., 1995, 2000; Brandes and
Devol, 2002].
Fixed nitrogen losses (and inputs) are likely influenced by episodic variations that are not
well constrained [Brandes and Devol, 2002] and need to be further understood. Both sediment
and water column denitrification play an important role in fixed N loss. The eastern tropical
north Pacific (ETNP), one of three global OMZs, accounts for approximately a third of global
water column denitrification [DeVries et al., 2012] and therefore fluctuations in the intensity of
denitrification in the ETNP influences the global budget of fixed nitrogen [Brandes and Devol,
2002].
The work presented in this dissertation uses the geochemical proxy, δ
15
N, measured in
laminated sediments (where δ
15
N
sed
‰ = [(
15
N/
14
Nsample)/(
15
N/
14
Nstandard) - 1] * 1000, and the
standard is atmospheric N
2
) from sites in the ETNP. My goals were to investigate both local
(10s km
2
) and regional (1000s km
2
) fluctuations in OMZ intensity over the past 1200 years,
examine the utility of δ
15
N
sed
as a tracer of ocean circulation and changes in ocean chemistry, and
investigate the synchronicity of the drivers behind OMZ fluctuations in the late Holocene.

1.2 The Biological Pump and δ
15
N as a Proxy of Changes in Strength of OMZs
During photosynthesis phytoplankton use carbon dioxide and nutrients in the surface
ocean to produce organic matter. The organic matter is then exported into the deep ocean where
a large amount is remineralized with a small portion (on average <10%) of the original organic
material produced reaching the sea floor, a process referred to as the biological pump. The
efficiency and strength of the biological pump, however, is dependent on the availability of
nutrients in the surface ocean; without nutrients organic matter cannot be produced. This is
5

critical as the efficiency of the biological pump ultimately impacts ocean biology and its ability
to drawdown CO
2
from the atmosphere [Gruber, 2004]. Variability in the efficiency of the
biological pump over geologic time may have altered major biogeochemical cycles [Altabet and
Francois, 1994].
The process of remineralization has important implications for nitrogen cycling. When
reminereralization occurs, organic matter is first degraded through aerobic respiration, where
oxygen is used as an electron acceptor. As oxygen concentrations decline to ~5µM, nitrate (NO
3
-
) is used by microbes to respire the organic matter through the process of denitrification. In the
denitrification reaction, NO
3
-
reacts with organic matter to produce nitrogen (N
2
)

or nitrous oxide
(N
2
O), thereby converting fixed N (in nitrate and PON) to gaseous N. N
2
and N
2
O can exchange
with the atmosphere and the dissolved nitrate that is not utilized in respiration, termed residual
NO
3
-
, becomes isotopically elevated. Denitrification preferentially fractionates the original pool
of NO
3
-
, insofar as the N
2
gas produced is depleted in
15
N, leaving the residual nitrate heavier in
15
N. When denitrification occurs in the water column it results in an isotopic enrichment (ε) of
about -20‰. ε is calculated by subtracting the value of δ
15
N of the product (N
2
) from the δ
15
N of
reactant (NO
3
-
) [Gruber, 2004].
When water column denitrification occurs close to the ocean mixed layer and in an area
that experiences upwelling, the elevated δ
15
NO
3
-
can be upwelled and incorporated into newly
produced organic matter. The process of production, export, remineralization and upwelling
continues, further enriching the residual NO
3
-
pool (Figure 1.3). In order to use δ
15
N from
sediments as a proxy for the degree of denitrification, several critical assumptions must be made:
(1) all of the nitrate that is upwelled into the surface water is incorporated into organic matter
through photosynthesis; there is complete nitrate utilization in the mixed layer on an annual basis
6

(2) the isotopic signature of the sinking particulate organic nitrogen (PON) does not change as it
sinks in the water column (3) diagenetic processes do not alter the δ
15
N signature after deposition
on the sea floor (4) denitrification (as opposed to anammox) is the primary mechanism of
nitrogen loss at the site.
When fixed N is the limiting nutrient and essentially all of the NO
3
-
brought to the mixed
layer is assimilated into organic matter, PON will take on an isotopic value of that nitrate. In
environments where nitrate is not the limiting nutrient, a fractionation occurs during nitrate
assimilation with the preferential uptake of
14
NO
3
-
by phytoplankton [Thunell and Kepple, 2004].
A study from the Arabian Sea discovered that when [NO
3
-
] is high in the surface water (25 µM)
the δ
15
N value from core top sediments was relatively low (5‰); however, when nitrate
concentrations are low in the surface water (<0.1 μM) the corresponding core top sediments were
enriched (11‰) in δ
15
N [Altabet and Francois, 1994]. Altabet and Francois were among the first
to use sedimentary δ
15
N as a paleoceanographic tracer of nitrate utilization. Their pioneering
work enabled further studies investigating nitrogen isotopes in sediment cores from the Arabian
Sea to investigate surface ocean nitrate utilization in the paleoceanographic record [Altabet et al.,
1995] and bound in diatoms [Robinson et al., 2005] and foraminifera [Straub et al., 2013]. The
critical component to investigating nitrogen isotope proxy reconstruction is the ability to
determine if complete or incomplete nitrate utilization occurs. The locations of this study are off
of Mexico and Southern California, which have shown complete nitrate utilization during
modern ocean conditions [Townsend-Small et al., 2014].
A powerful justification of using sediment
15
N as a proxy of water column nitrate isotopic
composition comes from a comparative study of the δ
15
N of NO
3
-
in the sub-photic zone,
particulate organic matter from sediment traps, and sediment core samples in the Gulf of
7

California, the San Pedro Basin, and the Monterey Bay. This work indicates that NO
3
-
, trap
material and sediments have similar δ
15
N values at all locations [Thunell et al., 2004]. These data
support several of the assumptions upon-which this study is based: (a) that PON reflects the
upwelled nitrate isotope value and (b) suggesting that δ
15
N of PON is not altered as the material
sinks in the water column is a valid hypothesis. Open ocean settings, however, do not have the
same coherence [Robinson et al., 2012]. Studies have shown that there is as much as 4‰
enrichment in the sediments compared to sediment trap material in the open ocean [Altabet,
1996; Altabet et al., 1999; Montoya, 2007]. This enrichment is absent in the coastal margin
settings due to a higher flux and accumulation of organic matter and better preservation [Altabet
et al., 1999, Robinson et al. 2012]. Many studies illustrate a 1:1 correlation between the δ
15
NO
3

in the upper water (sub thermocline) and δ
15
N in the sediments in productive areas where there is
a high sedimentation rate; sites where this is demonstrated include: Cariaco Basin [Thunell et al.,
2004], Gulf of Tehuntepec, Mexico [Thunell and Kepple, 2004], Monterey Bay [Altabet et al.,
1999], Santa Barbara Basin [Emmer and Thunell, 2000], San Pedro Basin [Altabet et al, 1999],
and Guaymas Basin [Pride et al., 1999] (Figure 1.4).
Studies have shown that in highly productive areas of the ocean, such as the Eastern
Tropical Pacific, alteration of sedimentary δ
15
N may be negligible [Kalansky et al., 2011; Altabet
et al., 1999, Kienast et al., 2002]. Likewise, Prokopenko [2004], and Sweeney and Kaplan,
[1980], found that no significant fractionation occurs as a result of diagenesis of marine organic
matter deposited in low oxygen conditions even if a substantial portion (>30%) of the organic
matter is degraded. Prokopenko [2004] compared δ
15
N of ammonium, a major metabolic product
of organic matter decomposition, in pore water to δ
15
N of bulk organic matter from the
sediments and showed that there was not preferential decomposition of nitrogen, illustrating that
8

a diagenetic signal is not overprinting δ
15
N
sed
in these environments. One study by Lehmann et
al. [2002], suggests diagenetic alteration of surficial sediments in anoxic conditions can decrease
δ
15
N. However, trends in particle flux data from the Santa Barbara Basin [Deutsch et al., 2014]
and subsurface sediments at all study sites provides strong evidence that the signal originates in
the water column, with minimal downcore diagenetic alteration.
Denitrification can result in the oxidation of proteins to CO
2
and H
2
O and the release of
an amino group in the form of NH
4
+

[Devol, 2008]. Thus, when denitrification occurs it should
also result in the accumulation of NH
4
+
, however, this is not observed in denitrification zones
[Codispoti, 1973]. This led to speculation that some form of anaerobic oxidation of ammonium
reaction (anammox) is occurring [Cline and Richards, 1972; Bender et al., 1989; Emerson et al.,
1980; Murray et al., 1995] but it was not until 1999 that a specific microbe capable of the
anammox reaction (NH
4
+
+ NO
2
-
! N
2
+ 2H
2
O) was identified [Strous et al., 1999]. Recent
studies have shown that microbially mediated anammox could be responsible for 28-48% of the
total OMZ N
2
production when anammox utilizes ammonium regenerated during organic matter
respiration by heterotrophic denitrification. However, if an additional source of ammonium is
available, anammox could account for a larger amount of total OMZ N
2
production [Hamersley
et al., 2007; Devol, 2003]. Recent studies have shown that anammox accounts for a large portion
of N-loss in the Peruvian upwelling and Bengula upwelling systems [Hamersley et al., 2007;
Kuypers et al., 2005; Lam et al., 2009]. Anammox is expected have an associated isotopic
fractionation, however, it has not been determined as of yet and likely has a similar isotope effect
as denitrification. Currently, researchers at Max Planck Institute for Marine Microbiology are
conducting experiments with a highly enriched, >98% Kuenenia stuttgartiensis (an
environmentally isolated microbe that is capable of the anammox reaction) culture to further
9

characterize the nitrogen isotope effects by anammox. Recent studies [Chong-jian et al., 2010;
Chamchoi et al., 2007], suggest that anammox is suppressed and denitrification dominates when
significant amounts of organic matter is present. Studies in the ETNP have not yet been
conducted.
Work remains to be done in understanding how much nitrogen loss in the OMZ off the
coast of Mexico and Southern California is due to anammox. Prokopenko et al. [2006]
conducted a study to elucidate the presence of anammox in sediments from sites along the
Eastern Pacific Margin based on the analysis and comparison of pore water ammonium and total
CO
2
. The study found that the Santa Barbara and Pescadero Basins exhibited only small
differences in sedimentary δ
15
N and δ
15
N of ammonium in pore water indicating that significant
amounts of anammox are not occurring in the sediments.
As mentioned previously, using δ
15
N
sed
as a proxy for water column denitrification is not
unprecedented. The magnitude of denitrification and OMZ intensity has been shown to vary on
glacial-interglacial time scales [Altabet et al., 1995; Ganeshram et al., 2000; Deutsch et al., 2005;
Thunell and Kepple, 2004]. High-resolution studies of changes in denitrification and OMZ
intensity on millennial, centennial, and decadal time scales, however, have not been investigated.

1. 3 Study Sites
This dissertation will focus on three modern localities in the eastern boundary of the
north Pacific basin (Figure 1.5). The first site, Pescadero Slope is at the eastern mouth of the
Gulf of California and in the eastern tropical north Pacific (ETNP) OMZ. Pescadero Slope
coring sites are centered at 24°15’N and 109°00’W, at a depth of 600 m on the slope that reaches
depths >2500 meters [Gonzalez-Yajimovich, 2005]. The Gulf of California is an elongate basin,
10

surrounded by land to the north, west, and east, opening to the Pacific Ocean in the south. In the
Fall, the Pacific High-Pressure System and the Inter-tropical Convergence Zone (ITCZ) migrate
south towards the equator and the winds intensify. As this occurs, the strength and prominent
direction of the winds change from weak southeastern winds to strong northwestern winds
[Douglas et al., 2007]. The strong northwestern winds drive Ekman transport in the Eastern
Gulf, resulting in intensified upwelling stimulating high productivity on the this side of the Gulf
from November to March [Staines-Urías et al., 2009; Douglas et al., 2007]. High productivity
coupled with moderate rates of deep thermohaline ventilation result in the formation of a well-
developed oxygen minimum zone between 100 and 800 meters [Fernández-Barajas et al., 1994].
Where the OMZ impinges on the basin slope, laminated sediments are deposited. At the
Pescadero site, the bottom water oxygen concentrations are below 3 µM [Berelson et al., 2005;
Chong et al., 2012]. Multicores and gravity cores were collected along the slope of the basin at a
depth of 600 m where laminated sediments are deposited. Because of the Pescadero Slope
location, the site is not only affected by the dynamics of the Gulf of California monsoon but also
influenced by Pacific Ocean circulation, making it an ideal location to study not only local
changes, but also regional phenomenon.
The second site, the Santa Monica Basin, is located in the Southern California
Borderland. The oceanographic setting of Santa Monica Basin makes this location a unique
environment in which the oxygenation record and intensity of the OMZ can be studied through
the analysis of the sediments. The intersection of the sill with low-oxygen waters, high
productivity and restricted water circulation affect the water chemistry within the basin. Water
circulation into and out of Santa Monica Basin (910 m depth) is restricted by a 740 m sill, which
traps low oxygen water [Christensen et al., 1994]. The water that intersects the sill is North
11

Pacific Intermediate Water, transported from the Eastern Tropical North Pacific by the California
Undercurrent at ~500 m depth. The water mass is characterized by low oxygen concentrations.
This combination of processes results in basin bottom water with a concentration of oxygen
between 5-15 µM, which does not support benthic macrofauna that can bioturbate (mix and
rework the upper sediments by burrowing) the basin floor. Consequently, the upper portion of
the sediment remains intact and original depositional features, such as laminations, are observed.
Analysis of the microstructure and fine lamination are excellent paleo-oxygenic proxies and can
be used to infer the severity and extent of anoxia in this and other borderland basins [Behl and
Kennett, 1996].
The third site, Soledad Basin, also known as San Lázaro Baisin, is located on the edge of
the continental shelf, 45 km to west of Baja California Sur (25° 10’ N, 112° 45’ W) between
Pescadero Slope and the Santa Monica Basin. Soledad Basin, a tectonically formed depression,
is approximately 85 km long and 35 km wide with the maximum depth near 540 m [Silverberg et
al., 2004]. It is characterized by a flat bottom and restricted water flow due to a sill at 290 m
[vanGeen et al., 2003]. Soledad Basin lies adjacent to the California Current System flowing
southward towards the equator and the California Undercurrent, which flows northward
[Sañudo-Wilhelmy et al., 2012]. Strong northerly winds induce offshore Ekman transport and
subsequent upwelling in the winter and spring months, with upwelling dissipating during late
summer and fall. The basins proximity to coastal upwelling result in relatively continuous high
productivity [Silverberg et al., 2004] which helps maintain a well-developed oxygen minimum
zone below which laminated sediments are deposited. Sediment particles from the basin are
dominated by marine snow, containing primarily fecal pellets, vacant pteropod shells, and
scattered foraminifera. Sediment trap studies have determined that there is a widely varied total
12

mass flux of sediment (63-587 mg m
-2
d
-1
) and POC fluxes range between 9 and 40 mgC m
-2
d
-1
.
Organic carbon content of the exported material ranges from 5.7 to 14% depending on the season
and C:N ratios varied between 7.4 and 12.7 [Silverberg et al., 2004]. Previous studies have
explored fluctuations in the oxygen minimum zone in the Soledad Basin and at open ocean sites
nearby by investigating variations in the deposition of laminated and bioturbated sediment along
the Pacific margin. Transitions from the last glacial period to the recent interglacial indicates
declines in O
2
content during the Holocene occurred between 11,300 to 10,000 years ago
[VanGeen et al., 2003].

1.4 Objectives
The overall objective of this dissertation is to develop high-resolution records that
reconstruct the eastern tropical north Pacific oxygen minimum zone intensity along the margin of
western North America at the Pescadero Slope, Soledad Basin, and the Santa Monica Basin. This
work investigates the range in which OMZ intensity varies in the late Holocene and assesses
what factors are driving these fluctuations on both local and regional scales. The primary mode
of ETNP OMZ intensity fluctuation reconstruction was done using nitrogen isotopes of bulk
organic sediment (δ
15
N
sed
) measured at a near annual resolution from multicores or box cores at
all three study sites and gravity cores at the Pescadero Slope and Soledad Basin sites. The age of
the sediment downcore was constrained through the development of
210
Pb (multicores and box
cores) and radiocarbon (gravity cores) age models and verified through varve counts when
possible.
Chapter 2 of this dissertation investigates the utility of δ
15
N
sed
as a tracer of changes in
ocean circulation, the strength of the California Undercurrent compared to the California
13

Current, and potential changes in mixing between these water masses. The California
Undercurrent transports elevated
15
NO
3
-
from the ETNP northwards into the California
Borderland [Castro et al., 2001], and thus by calculating the difference between δ
15
N
sed
from the
Pescadero Slope and the Santa Monica Basin, , variation in the strength of the California
Undercurrent and the factors influencing its strength can be assessed.
Chapter 3 assesses late Holocene trends in OMZ intensity at the Pescadero Slope site
through the construction and evaluation of a high-resolution 1200-year δ
15
N
sed
record. To assess
the factors driving local fluctuations in the ETNP OMZ, additional geochemical analyses are
presented. This includes measurements of sediment weight percent carbon (wt. % C) and weight
percent biogenic silica (wt. % bSi) to serve as a proxy for carbon export from the photic zone
and to investigate if increased carbon export (i.e. biological productivity) could be driving
variations in OMZ strength. The connection between OMZ intensity, ocean circulation, and
climate variability was assessed through spectral analysis of prominent periodicities present in
the record, wavelet analysis, and qualitative comparisons to dominant cycles influencing global
and regional dynamics. These include solar irradiance, the Pacific Decadal Oscillation (PDO),
and the El Niño Southern Oscillation (ENSO).
Finally, a regional evaluation of variation in OMZ strength was investigated through a
comparative analysis of δ
15
N
sed
between 1200-year records reconstructed at the Pescadero Slope
and the Soledad Basin (Chapter 4). In this chapter, X-ray florescence data analysis is used to
correlate the sedimentary records of both sites and investigate potential age model uncertainties.
Geochemical assessments and their implications of each location are also discussed.
14

References
Altabet, M. A. and R. Francois (1994), Sedimentary nitrogen isotopic ratio records surface ocean
NO3- utilization, Global Biogeochemical Cycles, 8, 103-116.
Altabet, M. A., R. Francois, D.W. Murray, and W. L. Prell (1995), Climate-related variations in
denitrification in the Arabian Sea from sediment 15N/14N ratios, Nature, 373, 506–509.
Altabet, M.A., C. Pilskaln, R. Thunell, C. Pride, D. Sigman, F. Chavez, R. Francois (1999), The
nitrogen isotope biogeochemistry of sinking particles from the margin of the Eastern North
Pacific, Deep Sea Research I, 46, 655-679.
Anderson, .H., C. Woulds, M. Schwartz, G.L. Cowie, L.A. Levin, K. Soetaert, J.J. Middelburg
(2008), Short-term fate of phytodetrtus in sediments across the Arabian Sea Oxygen
Minimum Zone, Biogeosciences, 5, 43-53.
Behl, R.J. and J.P. Kennett (1996), Brief interstadial events in the Santa Barbara basin, NE
Pacific, during the past 60 kyr, Nature, 379(6562), 243-246.
Bender, M., R. Jahnke, F. Weiss, W. Martin, D.T. Heggie, J. Orchardo, T. Sowers (1989),
Organic-carbon oxidation and benthic nitrogen and silica dynamics in San-Clemente Basin, a
Continental Borderland site, Geochimica et Cosmochimica Acta, 53(3), 685-697.
Benton, M.J., R.J. Twitchett (2003), How to kill (almost) all life: the end-Permian extinction
event, Trends in Ecology and Evolution, 18, 358–365.
Berelson, W.M., M. Prokopenko, F.J. Sansone, A.W. Graham, J. McManus, J.M. Bernhard
(2005), Anaerobic diagenesis of silica and carbon in continental margin sediments: discrete
zones of TCO2 production, Geochimica et Cosmochimica Acta, 69(19), 4611-4629.
Bindoff, N. L. and T. J. McDougall (2000), Decadal changes along an Indian Ocean section at
32 degrees S and their interpretation, Journal of Physical Oceanography, 30, 1207– 1222.
15

Bograd, S.J., C.G. Castro, E. Di Lorenzo, D.M. Palacios, H. Bailey, W. Gilly, and F.P. Chavez
(2008), Oxygen declines and the shoaling of the hypoxic boundary in the California Current,
Geophysical Research Letters, 35, L12607, doi: 10.1029/2008GL034185.
Brandes, J. A., and A. H. Devol (2002), A global marine-fixed nitrogen isotopic budget:
Implications for Holocene nitrogen cycling, Global Biogeochemical Cycles, 16(4), 1120,
doi:10.1029/2001GB001856.
Casciotti, K. L. (2016), Nitrogen and oxygen isotopic studies of the marine nitrogen cycle,
Annual Reviews of Marine Science, 8, 379-407, doi:
10.1146/annurev-marine-010213-
135052.
Castro, C. G., F. P. Chavez, and C. A. Collins (2001), Role of the California Undercurrent in the
export of denitrified waters from the eastern tropical North Pacific, Global Biogeochemical
Cycles, 15(4), 819–830, doi:10.1029/2000GB001324.
Chamchoi, N., S. Nitisoravut, and J.E. Schmidt (2008), Inactivation of ANAMMOX
communities under concurrent operation of anaerobic ammoium oxidation (ANAMMOX)
and denitrification, Bioresource Technology, 9 (99), 3331-6.
Chan, F., J.A. Barth, J. Lubchenco, A. Kirincic, H. Weeks, W.T. Peterson, and B.A. Menge
(2008), Emergence of anoxia in the California Current large marine ecosystem, Science, 319,
920.
Chong, L. S., M. G. Prokopenko, W. M. Berelson, A. Townsend-Small, and J. McManus (2012),
Nitrogen cycling within suboxic and anoxic sediments from the continental margin of
western North America, Marine Chemistry, 128–129, 13–25,
doi:10.1016/j.marchem.2011.10.007.
16

Chong-jian, T., P. Zheng, C-h Wang, and Q. Mahmood (2010), Suppression of anaerobic
ammonium oxidizers under high organic content in high-rate Anammox UASB reactor,
Bioresource Technology, 101(6), 1762-1768.
Christensen, C.J., D.S. Gorsline, D.E. Hammond, S.P. Lund (1994), Non-annual laminations and
expansion of anoxic basin-floor conditions in Santa Monica Basin, California Borderland,
over the past four centuries, Marine Geology, 116, 399-418.
Cline, J.D. and F.A. Richards (1972), Low oxygen concentrations and nitrate reduction in the
eastern tropical North Pacific ocean, Limnology and Oceanography, 17, 885-900.
Codispoti, L.A. (1973), Denitrification in the Eastern Tropical North Pacific ocean, Ph.D.
Dissertation, University of Washington, Seattle.
Deutsch, C., S. Emerson, L. Thompson (2005), Fingerprints of climate change in North Pacific
oxygen, Geophysical Research Letters, 32, L16604, doi:10.1029/2005GL023190.
Deutsch, C., H. Brix, T. Ito, H. Frenzel, and L. Thompson (2011), Climate-Forced Variability of
Ocean Hypoxia, Science, 333(6040), 336-339, doi:10.1126/science.1202422.
DeVries, T., C. Deutsch, F. Primeau, B. Chang, and A. Devol (2012), Global rates of water-
column denitrification derived from nitrogen gas measurements, Nature Geoscience, 5(8),
547–550, doi:10.1038/ngeo1515.
Devol, A.H. (2003), Solution to a marine mystery, Nature, 422, 575-576.
Devol, A.H. (2008), Denitrification including anammox, in Nitrogen in the Marine Environment
(ed. Capone, D.), 263-301, doi: 10.1016/B978-0-12-372522-6.00006-2.
Diaz, R.J., R. Rosenberg (2008), Spreading dead zones and consequences for marine
ecosystems, Science, 321, 926-929.
17

Douglas, R., O. Gonzalez-Yajimovich, J. Ledesma-Vazquez, and F. Staines-Urias (2007),
Climate forcing, primary production and the distribution of Holocene biogenic sediments in
the Gulf of California, Quaternary Science Reviews, 26(1–2), 115–129,
doi:10.1016/j.quascirev.2006.05.003.
Emerson, S.R., R. Jahnke, M. Bender, P. Froelich, G. Klinkhammer, C. Bowser, F. Setlock
(1980), Early diagenesis in sediments from the eastern equatorial Pacific. I. Pore water
nutrient and carbonate results, Earth and Planetary Science Letters, 49, 57-80.
Emerson, S., S. Mecking, and J. Abell (2001), The biological pump in the subtropical North
Pacific Ocean: Nutrient sources, Redfield ratios, and recent changes. Global
Biogeochemical Cycles, 15, 535–554.
Emmer, E. and R. Thunell (2000), Nitrogen isotope variations in Santa Barbara Basin sediments:
implications for denitrification in the eastern tropical North Pacific during the last 50,000
years, Paleoceanography, 15, 377-387.
Fernández-Barajas, M.E, M.A. Monreal-Gómez, A. Molina-Cruz (1994), Thermohaline structure
and geostrophic flow in the Gulf of California, during 1992, Ciencias Marinas, 20, 267–286.
Ganeshram, R.S., T.F. Pedersen, S.E. Calvert, and J.W. Murray (1995), Large changes in oceanic
nutrient inventories from glacial to interglacial periods, Nature, 376, 755-758.
Ganeshram, R. S., T. F. Pedersen, S. E. Calvert, G. W. McNeil, and M. R. Fontugne (2000),
Glacial-interglacial variability in denitrification in the world’s oceans: Causes and
consequences, Paleoceanography, 15(4), 361–376, doi:10.1029/1999PA000422.
Garcia, H., A. Cruzado, L. Gordon, and J. Escanez (1998), Decadal-scale chemical variability in
the subtropical North Atlantic deduced from nutrient and oxygen data, Journal of
Geophysical Research, 103, 2817– 2830.
18

Getzlaff, J. and H. Dietze (2013), Effects of increased isopycnal diffusivity mimicking the
unresolved equatorial intermediate current system in an earth system climate model,
Geophysical Research Letters, 40, 2166-2170, doi:10.1002/grl.50419.
Gilly, W.F., J.M. Beman, S.Y. Litvin, and B.H. Robinson (2013), Oceanographic and biological
effects of shoaling of the oxygen minimum zone, Annual Reviews in Marine Science, 5,
21.1-21.28, doi:10.1146/annurev-marine-120710-100849.
Gonzalez-Yajimovich, O., R. G. Douglas, and D. S. Gorsline (2005), The preserved carbonate
record in Holocene sediments of the Alfonso and Pescadero basins, Gulf of California,
Mexico, Proc. Geol. Assoc., 116, 315–330, doi:10.1016/S0016-7878(05)80050-1.
Gruber, N. (2004), The dynamics of the marine nitrogen cycle and its influence on the
atmospheric CO
2
variations, In: Carbon-Climate Interactions NATO ASI (eds., Follows, M.
and Oguz, T.), John Wiley and Sons, New York: 1-39.
Hamersley, M.R., G. Lavik, D. Woebken, J.E. Fattray, P. Lam, E.C Hopmans, J.S.S. Damste, S.
Kruger, M. Graco, D. Gutierrez, and M.M.M. Kuypers (2007), Anaerobic ammonium
oxidation in the Peruvian oxygen minimum zone, Limnology and Oceanography, 52(3), 923-
933.
Jones, C.E. and H.C. Jenkyn (2001), Seawater Strontium Isotopes, Oceanic Anoxic Events, and
Seafloor Hydrothermal Activity in the Jurassic and Cretaceous American Journal of Science,
301, 112–149.
Kalansky, J.F., R.S. Robinson, B.N. Popp (2011), Insights into nitrogen cycling in the western
Gulf of California from the nitrogen isotopic composition of diatom‐bound organic matter,
Geochemistry, Geophysics, Geosystems, 12(6), Q06015, doi:10.1029/2010GC003437.
19

Keeling, R.F. and H.E. Garcia (2002), The change in oceanic O
2
inventory associated with recent
global warming, Proceedings of the National Academy of Sciences, 99 (12), 7848-7853.
Keeling, R.F., A. Körtzinger, N. Gruber (2010), Ocean Deoxygenation in a Warming World,
Annual Review of Marine Science, 2, 199-229.
Kienast, S. S., S. E. Calvert, and T. F. Pederson (2002), Nitrogen isotope and productivity
variations along the northeast Pacific margin over the last 120 kyr: Surface and subsurface
paleoceanography, Paleoceanography, 17(4), 1055, doi:10.1029/2001PA000650.
Kuypers, M.M.M, G. Lavik, D. Woebken, M. Schmid, B.M. Fuchs, R. Amann, BB Jorgensen,
and m.S.M. Jetten (2005), Massive nitrogen loss from the Benguela upwelling system
through anaerobic ammonium oxidation, Proceedings of the National Academy of Sciences,
102(18): 6478-6483, doi︎10.1073︎pnas.0502088102.
Lam, P., G. Lavik, M.M. Jensen, J. van de Vossenberg, M. Schmid, D. Woebken, D. Gutiérrez,
R. Amann, M.S.M. Jetten, and M.M.M. Kuypers (2009), Revising the nitrogen cycle in the
Peruvian oxygen minimum zone, Proceedings of the National Academy of Sciences, 106(12),
4752-4757.
Levin, L.A., and N. LeBris (2015), The deep ocean under climate change, Science, 350(6262),
766-768, doi:10.1126/science.aad0126.
Matear, .J., A.C. Hirst, and B.I. McNeil (2000), Changes in dissolved oxygen in the Southern
Ocean with climate change, Geochemistry, Geo. Geophysics, Geosystems, 12000GC000086.
McClatchie, S., R. Goericke, R. Cosgrove, G. Auad, R. Vetter (2010), Oxygen in the Southern
California Bight: Multidecadal trends and implications for demersal fisheries, Geophysical
Research Letters, 37, L19602, doi: 10.1029/2010GL044497.
20

Montoya, J.P. (2007), Natural abundance of
15
N in marine planktonic ecosystems, Stable
isotopes in ecology and environmental science, 176.
Murray, J.W., L.A. Codispoiti, G.E. Friederich (1995), Oxidation-reduction environments: the
suboxic zone in the Black Sea, In: Aquatic Chemistry: Interfacial and Interspecies Processes
(eds. Huang, C.P., C.R. O’Melia, and J.J. Morgan), American Chemical Society,
Washington, DC, 157-176.
Pride, C., R. Thunell, D. Sigman, L. Keigwin, M. Altabet, E. Tappa (1999), Nitrogen isotopic
variations in the Gulf of California since the Last Deglaciation: response to global climate
change, Paleococeanography, 14(3), 397-409, doi: 1.1029/1999PA900004.
Prince, E.D., C.P. Goodyear (2006), Hypoxia-based habitat compression of tropical pelagic
fishes, Fisheries Oceanography, 25: 451-464.
Prokopenko, M. (2004), Fractionation of nitrogen isotopes during early diagenesis, Ph.D
Dissertation, University of Southern California, Los Angeles.
Prokopenko, M.G., D.E. Hammond, W.M. Berelson, J.M. Bernhard, L. Stott, R. Douglas
(2006), Nitrogen cycling in the sediments of Santa Barbara basin and the Eastern Subtropical
North Pacific: Nitrogen isotopes, diagenesis and possible chemosymbiosis between two
lithotrophs (Thioploca and Anammox)—“riding on a glider.” Earth and Planetary Science
Letters, 242, 186-204.
Purcell, J.E., D.L. Breitburg, M.B Decker, W.M Graham, M.J. Youngbluth, K.A. Raskoff
(2001), Pelagic cnidarians and ctenophores in low dissolved oxygen environments. In:
Costal Hypoxia: Consequences for Living Resources and Ecosystems (eds. Rabalais, N.B.
and R.E. Turner), American Geophysical Union, Washington, DC, 77-100.
21

Purcell, J.E., S. Uye, W.T. Lo (2007), Anthropogenic causes of jellyfish blooms and their direct
consequences for humans: a review, Marine Ecology Progress Series, 350, 153-174.
Robinson, R.S., D.M. Sigman, P.J. DiFiore, M.M. Rohde, T.A. Mashiotta, and D.W. Lea (2005),
Diatom-bound
15
N/
14
N: new support for enhanced nutrient consumption in the ice age
subantarctic, Paleoceanography, 20, PA3003, doi:10.1029/2004PA001114.
Robinson, R. S., et al. (2012), A review of nitrogen isotopic alteration in marine sediments,
Paleoceanography, 27, PA2403, doi:10.1029/ 2012PA002321.
Sañudo-Wilhelmy, S.A., L.S. Cutter, R. Durazo, E.A. Smail, L. Gómez-Consarnau, E.A. Webb,
M.G. Prokopenko, W.M. Berelson, and D.M. Karl (2012), Multiple B-vitamin depletion in
large areas of the coastal ocean, Proceedings of the National Academy of Sciences, 109(35),
14041-14045.
Sarmiento, J. L., T.M.C. Hughes, R.J. Stouffer, S. Manabe (1998), Simulated response of the
ocean carbon cycle to anthropogenic climate warming, Nature, 393; 245– 249.
Silverberg, N., A. Martıınez, S. Aguíñiga, J. D. Carriquiry, N. Romero, E. Shumilin, and S. Cota
(2004), Contrasts in sedimentation flux below the southern California Current in late 1996
and during the El Niño event of 1997–1998, Estuarine, Coastal and Shelf Science, 59(4),
575–587, doi:10.1016/j.ecss.2003.11.003.
Staines-Urías, F., R.G. Douglas, D.S. Gorsline (2009), Oceanographic variability in the southern
Gulf of California over the past 400years: Evidence from faunal and isotopic records from
planktic foraminifera. Palaeogeography, Palaeoclimatology, Palaeoecology, 284 (3-4), 337-
354.
Stramma, L., GC. Johnson, J. Sprintall, V. Mohrholz (2008), Expanding oxygen-minimum zones
in the tropical oceans. Science, 320, 655-658.
22

Stramma, L., S. Schmidtko, L.A. Levin, G.C. Johnson (2010), Ocean oxygen minima expansions
and their biological impacts. Deep-Sea Research I, 57, 587-595.
Straub, M., D.M. Sigman, H. Ren, A. Martinez-Garcia, A.N. Meckler, M.P. Hain, and G.H.
Haug (2013), Changes in North Atlantic nitrogen fixation controlled by ocean circulation,
Nature, 501(7466), 200-203, doi:10.1038/nature12397.
Strous, M., J.A. Fuerst, E.H.M. Kramer, S. Logemann, G. Muyxer, K.T. Van de Pas-schoonen,
R. Webb, G.C. Kuenen, M.S.M. Jetten (1999), Missing lithotroph identified as new
planctomycetes, Nature, 400, 446-449.
Sweeney, R.E. and I.R. Kaplan (1980), Natural abundances of N-15 as a source indicator for
near-shore marine sedimentary dissolved nitrogen, Marine Chemistry, 9, 81-94.
Thunell, R. and A. Kepple (2004), Glacial-Holocene δ15N record from the Gulf of Tehuantepec,
Mexico: Implications for denitrification in the eastern equatorial Pacific and changes in
atmospheric N
2
O, Global Biogeochemical Cycles, 18(1), doi:10.1029/2002GB002028.
Thunell, R., D.M. Sigman, F.E. Muller-Karger, Y. Astor, R. Varela (2004), Nitrogen Isotopic
Dynamics of the Cariaco Basin, Venezuela, Global Biogeochemical Cycles, 18,GB3001, doi:
10.1029/2003GB002185.
Townsend-Small, A., M. G. Prokopenko, and W. M. Berelson (2014), Nitrous oxide cycling in
the water column and sediments of the oxygen minimum zone, eastern subtropical North
Pacific, Southern California, and Northern Mexico (23°N–34°N), Journal of Geophysical
Research Oceans, 119, 3158–3170, doi:10.1002/2013JC009580.
VanGeen, A., Y. Zheng, J.M. Bernhard, K.G. Cannariato, J. Carriquiry, W.E. Dean, B.W.
Eakins, J.D. Ortiz, and J. Pike (2003), On the preservation of laminated sediments along the
23

western margin of North America, Paleoceanography, 18(4), 1098,
doi:10.1029/2003OA000911.
Watanabe, Y. W., T. Ono, A. Shimamoto, T. Sugimoto, M. Wakita, and S. Watanabe (2001),
Probability of a reduction in the formation rate of the subsurface water in the North Pacific
during the 1980s and 1990s, Geophysical Research Letters, 28, 3289– 3292.
Wignall, P.B. and R.J. Twitchett (1996), Oceanic anoxia and the end Permian mass extinction,
Science, 272, 1155-1158.

24
Oxygen Declines in the Southern California Bight at 100 Meters
220
200
180
160
140
120
100
80
60
40
Dissolved [O
2
] (μM)
1984 1989 1994 1999 2004 2009 2014
Year (AD)
Bograd et al., 2008
y = -1.712 +3560
r
2
= 0.1790
SPOT
y = -1.712 +3560
r
2
= 0.1067
Figure 1.1. Oxygen declines at 100 meters in the water column from Bograd et al., 2008 analysis of California Cooperative Oceanic
Fisheries Investigations (CalCOFI) program between 1984 and 2006 are represented in black. A linear regression indicates a decline of
1.712 μM O2 per year. Time series oxygen measurements from the San Pedro Time Series station, located in San Pedro Basin of the
Southern California Bight, between 1998 and 2014 are shown in gray. San Pedro Basin, an inner basin in the California Borderland,
indicates the same overall decline of 1.712 μM O2 per year.
25
Pressure
(dbar)
Year (AD)
[O
2
] (μmol kg
-1
)
at 400 m
[O
2
]
(μmol kg
-1
)
a.
b.
Study Site D
Figure 1.2. This figure is slightly modified from Stramma et al. [2008]. a) shows the distribution of [O
2
] (μmol kg
-1
) at 400 meters in the water
column in the Pacific Basin. Two prominent oxygen minimum zones are see in the eastern tropical north Pacific and eastern tropical south Pacific.
b) Analysis of recent and historical O
2
measurements at Study Site D from Stramma et al. [2008], in the eastern tropical Pacific and closest to the field
sites of this study, shows an expansion of the OMZ for the past 30 years. One dbar (unit of pressure) is approximately equivalent to 1 meter. The
while lines labeled 60, demark the upper and lower bounds of the OMZ, in between which [O
2
] is below 60 μmol kg
-1
and considered hypoxic.
26
N Assimilation
C
org
N
org
Sediments
C
org
+ NO
3
-
N
2
+ CO
2
Denitrification
15
NO
3
-
OMZ
C
org
+ O
2
!CO
2
+H
2
O
Aerobic Respiration
OLZ Photic

15
N
15
NO
3
!
Upwelling
Figure 1.3. A cartoon illustrating the utility of δ
15
N
sed
as a proxy for fluctuations in denitrification and OMZ intensity. Organic
matter is created through photosynthesis in the photic zone but quickly begins to sink and degrade. It is first degraded through
aerobic respiration. Once [O
2
] reaches sufficiently low levels, denitrification takes over. Denitrification results in the preferetial
fractionation of the residual pool of NO
3
-
. Elevated
15
NO
3
-
is upwelled to the surface and completely assimilated into organic
matter. As the cycle continues, the elevated
15
NO
3
-
signal will be recorded in the sedimentary organic matter. Therefore,
elevated δ
15
N
sed
values indicate more denitrification a more intense OMZ.
27
14
12
10
8
6
4
2
0
14 12 10 8 6 4 2 0
δ
15
NO
3
-
(‰)
δ
15
N
sed
(‰)
Cariaco Basin
San Pedro Basin
Guaymas Basin
Gulf of Tehuantepec
Santa Barbara Basin
Monterey Bay
1:1
Figure 1.4. This figure is adapted from Thunell et al. [2004]. This illustrates that in the Cariaco Basin, Gulf of Tehuantepec,
Monterey Bay, Santa Barbara Basin, San Pedro Basin, and Guaymas Basin the δ
15
NO
3
-
at the base of the photic zone is very
similar to δ
15
N in surface sediments. Guaymas Basin is in the Gulf of California and has similar environmental conditions as
the Pescadero Slope and the Soledad Basin. San Pedro Basin is also adjacent to the Santa Monica Basin, also with a similar
depositional environment. These locations are all characterized by low oxygen content, high sediment accumulation rates, and
high preservation rates.
28
20˚N
25˚N
30˚N
120˚W 115˚W 110˚W 105˚W
Pescadero
Slope
Pacic Ocean
Baja California
Mexico
Gulf of California
United States
Soledad Basin
Figure 1.5. A map showing the three study sites: Santa Monica Basin, Soledad Basin and
Pescadero Slope in the eastern tropical north Pacific.
Santa
Monica Basin
29

Chapter 2
δ
15
N in laminated marine sediments provide a proxy for mixing between the
California Undercurrent and the California Current: A proof of concept

Abstract
Measurements of particulate δ
15
N in coastal marine laminated sediments provide a high-
resolution proxy for fluctuations in the intensity of denitrification in the water column. In the
Eastern Tropical North Pacific oxygen minimum zone, this denitrification signal is transported
northward by the California Undercurrent, thus serving as a tracer of ocean circulation. This is
verified through comparisons between salinity in the thermocline off Southern California (Santa
Monica Basin) and the difference between δ
15
N
sed
within age equivalent sediments from a
southern (Pescadero Slope) and northern (Santa Monica Basin) site. Trends in this parameter,
Δδ
15
N
sed
, relate to Pacific Decadal Oscillation (PDO) phase changes between 1900 and 1990.
We hypothesize that the decline in Δδ
15
N
sed
during warm PDO phases is due to a strengthening
of the California Undercurrent transporting
15
N-enriched nitrate from the ETNP northward. The
deviation from this trend after 1990 suggests recent changes in circulation and/or California
Current water nutrient biogeochemistry.

2.1. Introduction
Laminated sediments off Southern California and Baja California serve as a high-
resolution archive of paleoceanographic proxies [Schimmelmann and Tegner, 1991; Christensen
et al., 1994; Behl and Kennett, 1996; Hagadorn, 1996; Kemp, 1996; Schimmelman and Lange,
1996; Pride et al., 1999; Kemp et al., 2000; van Geen et al., 2003; Thunell and Kepple, 2004;
González-Yajimovich, 2005, 2007; Hendy and Pedersen, 2006; Schimmelmann et al., 2012], one
30

of which is the δ
15
N of marine organic matter. The sedimentary nitrogen isotope value (δ
15
N
sed
)
in this region is sensitive to fluctuations in the intensity of water column denitrification in the
core of Eastern Tropical North Pacific (ETNP) Oxygen Minimum Zone (OMZ) [Altabet et al.,
1999; Pride et al., 1999; DeVries et al., 2012; Deutsch et al., 2011, 2014]. The ETNP, one of
three global OMZs, accounts for approximately a third of global water column denitrification
[DeVries et al., 2012] and therefore fluctuations in the intensity of denitrification in the ETNP
influences the global budget of fixed nitrogen [Brandes and Devol, 2002]. The Pescadero Slope,
located at the mouth of Gulf of California and near the core the ETNP OMZ, provides an ideal
study location for reconstructing high-frequency fluctuations in water column denitrification,
while Santa Monica Basin (SMB), also laminated, receives the signal of denitrification through
water mass transport by the California Undercurrent.
Water column denitrification results in an isotopic fractionation imprinting a regional
signature of elevated
15
N/
14
N on the residual water column nitrate, δ
15
N
nitrate
(expressed as δ
15
N
per mil = [(
15
N/
14
N
sample
)/(
15
N
/14
N
standard
) -1]*1000, where the standard is atmospheric N
2
). As the
residual nitrate bearing the characteristically high δ
15
N signal is upwelled along the coast, fueling
algal production, the isotopic signal of denitrification is imparted on the algal particulate organic
matter (POM). As long as the algal consumption of nitrate within the euphotic zone is complete
[e.g. Altabet and Francois, 1994], and the δ
15
N of sinking POM is not diagenetically altered, as is
the case in the low O
2
water-column and sediments of the ETNP region [Thunell et al., 2004;
Collins et al., 2011], the δ
15
N of exported POM will reflect the magnitude/intensity of regional
denitrification [Altabet et al., 1999; Ganashram et al., 1995; Thunell and Kemple, 2004].
31

It is common practice to measure bulk δ
15
N
sed
(containing both organic and inorganic N
components), using simple evaluation tools to test whether the observed δ
15
N signal is primarily
organic [Robinson et al., 2012]. A 9-11 C:N ratio, a zero intercept plot of weight percent
nitrogen vs. weight percent carbon, and POM δ
13
C values of -23.18 to -22.15‰ all indicate that
the signal represented is organic matter. While diagenetic alteration of δ
15
N signature deposited
on the ocean floor is also a concern in more oxic depositional environments [Robinson et al.,
2012, Galbraith et al., 2012], multiple studies have demonstrated the absence of alteration of the
δ
15
N in rapidly accumulating organic-rich sediments of the coastal upwelling regimes [Altabet,
1999; Pride et al., 1999; Prokopenko et al., 2006a, b, c]. These measures and metrics support
δ
15
N
sed
as a robust proxy for reconstructing the regional history of denitrification [Altabet et al,
1999; Thunell et al, 2004; Thunell and Kemple, 2004; Haug et al., 1998; De Pol-Holz et al.,
2006; Martinez and Robinson, 2010]. We do note, however, that Lehmann et al. [2002] have
found some diagenetic alteration of
15
N values of PON during short-term degradation
experiments conducted under anoxic conditions. Nonetheless, we rely on the tight relationship
between sedimentary and water column nitrate δ
15
N found by Thunell et al. [2004] as the
foundation upon which we build our interpretations.
Previously published records of bulk sedimentary δ
15
N within ETNP indicate variability
in the intensity/magnitude of denitrification on glacial/interglacial time scales [Altabet et al.,
1995; Ganeshram et al., 2000]. A recent study by Deutsch et al. [2014], based on high-resolution
δ
15
N
sed
records from the Gulf of California and the California Borderland, demonstrated links
between the strength of trade winds and ETNP OMZ intensity. While these studies focused on
ETNP, our study delves into developing δ
15
N
sed
as a potential tool for monitoring changes in the
degree of transport and mixing between the northern sourced California Current and southern
32

sourced California Undercurrent. We interpret the mixing signal in the context of recent
hydrological and biogeochemical variability likely driven by changing climate and/or ocean
circulation.
Altabet et al. [1999] hypothesized that not only would changes in δ
15
N
sed
downcore
record the intensity of water column denitrification, it would also indicate the transport of the
isotopic signal along the western coast of North America if examined along the path of
transported water masses. The δ
15
N of water column nitrate along the Baja peninsula show
higher δ
15
N
nitrate
values in comparison to lower isotopic values of nitrate in northern waters off
southern California [Sigman et al., 2003, 2005; Townsend-Small et al. 2014]. This trend can be
interpreted as reflecting two-end-member mixing between ETNP and waters derived from the
northern California Current [Liu and Kaplan, 1989; Kienast et al., 2002]. Examining the isotopic
value of PON at two locations, one southern (Pescadero Slope) and one more northern (SMB),
we assess the degree of mixing between the southern and the northern sourced water masses.
Recent modeling studies predict changes in the sources and hence biogeochemistry of northern
water and intensification of these changes in the future [Rykaczewski and Dunne, 2010]. Our
two-end member δ
15
N
sed
mixing model provides a tool to further investigate these predicted
changes.
We make the case that the δ
15
N
sed
values at the Pescadero Slope site are set by changes in
denitrification intensity in the ETNP OMZ and that the similar patterns observed in SMB reflect
the advection of this signal northward by the California Undercurrent [Liu and Kaplan, 1989;
Castro et al., 2001; Kienast et al., 2002]. We expand this view to account for subtle, decadal-
scale differences in δ
15
N
sed
values between the southern and northern sites.
33

2.2. Methods
Measurements of δ
15
N
sed
were made on sediment splits from multicores [Barnett et al.,
1984] obtained at two sites. The Pescadero Slope core was collected in 2009 at 24°16.88’N and
108°11.79’W at 616 m water depth. Overlying water at this site has [O
2
] = 0.4 µM. The SMB
core was collected in 2011 at 33°50.21’N and 119°01.75’W at 892 m water depth with bottom
water [O
2
] ~3 µM [Chong et al., 2012]. Both cores are laminated as evidenced in x-radiographs
and high-resolution photo-scans (Figure 2.1). The Pescadero Slope core was split length-wise
and sampled at a 3 mm resolution. The SMB core was sampled using a core extruder and
samples were cut at 0.8 mm intervals. Water column and sediment pore water chemistry at these
two sites are described in Chong et al. [2012], Prokopenko et al. [2006a, 2011], and Townsend-
Small et al. [2014].
Individual samples were prepared for bulk δ
15
N analysis by drying and grinding to a fine
powder. The samples were analyzed by an elemental analyzer interfaced to a continuous flow
isotope ratio mass spectrometer (IRMS) at the UC Davis Stable Isotope Facility. The references
used to normalize the data were G-11 (nylon, δ
15
N of -9.77‰), G-12 (glutamic acid-enriched,
δ
15
N of 45.31‰), G-13 (bovine liver, δ
15
N of 7.72‰), and G-9 (glutamic acid, δ
15
N of -4.26‰).
The average standard error between replicate Pescadero Slope samples was 0.03‰, and was
0.11‰ for SMB samples. Standard error was calculated by dividing the average standard
deviation by n-1 where n represents the number of samples with replicates (Pescadero Slope
n=13, SMB n=5). We also examined the uncertainty in the difference in δ
15
N
sed
between these
sites, by individually squaring the standard errors at each site, followed by summing, and then
taking the square root of this sum. The resulting standard error of the difference is 0.12‰.
34

Lead-210 from the Pescadero Slope was measured via gamma spectroscopy and porosity
data were used to establish the pattern of integrated mass with depth [Deutsch et al., 2014].
Integrated mass plotted with excess
210
Pb was fit with an age model for the Pescadero Slope
multicore (Figure 2.2) and interpretation of this profile yields a mass accumulation rate (MAR)
of 77.4 mg cm
-2
yr
-1
. This rate is 20% lower than the accumulation rate found by Staines-Urías et
al. [2009], however, Staines-Urías et al. found their accumulation rate to be higher than studies
by Barron and Bukry [2007] and Douglas et al. [2007] on cores located nearby. The age model
for the SMB core was constructed using porosity data from previous studies and an assumed
accumulation rate of 16 mg cm
-2
yr
-1
as reported from analyses of 9 cores by Christensen et al.
[1994] who found the rates to be consistent (±15%) throughout much of the basinal region of the
SMB. Uncertainty was assigned a value of ±10% to the best-estimate of MAR; Pescadero Slope
MAR = 77.4±7.7 mg cm
-2
yr
-1
SMB MAR = 16±2 mg cm
-2
yr
-1
.

2.3. Results
Downcore δ
15
N
sed
record
The downcore pattern of δ
15
N
sed
plotted with sediment age shows coherence between the
SMB and Pescadero Slope records (Figure 2.3), although there is an offset of 1-2 ‰ between the
two sites. Sedimentary δ
15
N
sed
values from the Pescadero Slope vary between 8.5 and 10‰
whereas the SMB range is 6.7-9‰. Both records show general similarity in the direction of
change, notably the marked increase (towards higher values) in δ
15
N
sed
between 1993 and the
present. Both records also show a decrease in δ
15
N
sed
values between 1900 and 1990.
Although the general pattern of δ
15
N
sed

vs. time is similar between the two sites, there are
intervals when these records are within <1‰ and other times when they differ by >2‰. It is
35

likely that the variability in δ
15
N
nitrate
(and consequently δ
15
N
sed
) at the Pescadero Slope site is
driven by fluctuations of the intensity of denitrification in the ETNP [Deutsch et al., 2014], while
the California Current waters have a more constant δ
15
N
nitrate
value due to the absence of
denitrification in these waters. Thus we hypothesize, as southern waters carried by the California
Undercurrent mix with northern waters, it is the degree of mixing that sets the δ
15
N
nitrate
value of
waters off Southern California [Liu and Kaplan, 1989], and consequently is recorded in the SMB
sedimentary δ
15
N
sed
.
Testing this premise, we compared 50 years of California Cooperative Oceanic Fisheries
Investigations (CalCOFI) salinity data from a site in SMB (line 86.7 station 40) to the age-
corresponding δ
15
N
sed
value from the SMB sediment core (Figure 2.4a). The yearly average
salinity was calculated by averaging salinity values obtained quarterly between 150-250 m, the
depth of the core of the California Undercurrent [Hickey, 1992]. Another comparison was made
between SMB δ
15
N
sed
in a given year and the depth of water defined by sigma t = 26.4-26.5 (the
core of the California Undercurrent) for that year. This comparison yielded no correlation (r
2

=0.0041), which rules out the possibility that changes in δ
15
N
sed
were due to a change in depth of
the California Undercurrent and hence the depth of upwelled waters.
Southern water has characteristically higher salinity than northern water [Castro et al.,
2001] and higher δ
15
N
nitrate
[Liu and Kaplan, 1989] than northern water, hence, we predict a
positive correlation between salinity and SMB δ
15
N
sed
. The observed relationship (Figure 2.4a)
supports our premise. It is important to take into consideration that the δ
15
N
nitrate
-mixing trend is
non-linear, as each end member must be weighted by its nitrate concentration. Kienast et al.
[2002], compiling δ
15
N
nitrate
and salinity data collected along the coast of western North America,
illustrated that when these parameters are directly compared they plot on the ideal mixing line
36

developed by Liu and Kaplan [1989], confirming that southern ETNP water is an important
source of
15
N enriched nitrate to the western coast of North America.
To further evaluate the robustness of δ
15
N
sed
as an indicator of the transport of subsurface
δ
15
N
nitrate
values into the California Borderland, we utilize the parameter Δδ
15
N
sed
. We define
Δδ
15
N
sed
as the difference in δ
15
N
sed
between the Pescadero Slope and SMB during a given year
and compare this parameter to salinity in the SMB (Figure 2.4b). The advantage of using
Δδ
15
N
sed
is that it accounts for changes in the values of transported δ
15
N
nitrate
driven solely by
denitrification in the southern waters. The plot of Δδ
15
N
sed
vs. salinity (Figure 2.4b) has a better
correlation coefficient and ANOVA p value than does the SMB δ
15
N
sed
vs. salinity plot, thus
supporting our hypothesis that there is mixing between two end-member water masses, but that
the southern water end-member δ
15
N
nitrate
value fluctuates.
The model presented above works for several reasons: (1) the oxygen concentrations in
the California Borderland California Current are not low enough for denitrification to actively
occur in the sub-thermocline water column, and (2) during upwelling in the spring, the algal
organic matter faithfully records the δ
15
N
nitrate
from sub-thermocline waters [Collins et al., 2011,
Altabet et al., 1999, Price et al., 1999] indicating complete nitrate consumption. Yet, this model
certainly oversimplifies the physics of oceanographic processes; for example, it is unlikely that
nitrate isotope values in the Southern California sub-thermocline waters reflect simple two-end-
member mixing between northern and southern waters. Nor is it likely that northern water nitrate
isotopic values in sub-thermocline waters have remained constant for the last 100 years. While a
significant correlation between salinity in the SMB vs. δ
15
N
sed
in SMB and between salinity in
SMB vs. Δδ
15
N
sed
exists from 1953 through 1990 (p = 0.0170 and 0.0098 respectively), both of
37

these correlations weaken during the time window 1993 to 2009 (Figure 2.4c/d). This suggests
that the ideal mixing model may not apply with respect to the recent 20 years.

2.4. Discussion
Factors driving fluctuations in the poleward transport of nitrate δ
15
N

There have been many studies focusing on forcing of the California Current [Lynn and
Simpson, 1987; Hickey, 1992; DiLorenzo, 2003], however, studies specifically about the
California Undercurrent are limited. It is well established that the Pacific Decadal Oscillation
(PDO) influences the dynamics of the California Current [Mantua and Hare, 2002]. As the
Pacific High intensifies relative to the Aleutian Low, the California Current strengthens along the
western coast of North America. This results in increased upwelling and colder surface water
temperatures, indicated by negative PDO index values. By analyzing Δδ
15
N
sed
between the
southern and northern sites we gain insight into the strength of the California Undercurrent and
how it relates to changes in and mixing with the California Current.
The Δδ
15
N
sed
parameter is a metric based on the similarity in sub-thermocline nitrate
isotopic values between the Pescadero Slope and SMB (Figure 2.5). As Δδ
15
N
sed
values get
smaller, more southern water is influencing waters off Southern California. When Δδ
15
N
sed

values increase, we propose that less southern water is transported into the Southern California
Bight. Our analyses of salinity vs. Δδ
15
N
sed
correlations between 1950-1990 (Figure 2.4a/b)
suggest that this proxy is valid, and may be useful if extended to periods where we have no water
column data.
We relate distinct periods of variability in Δδ
15
N
sed
to changes in the degree of transport
and mixing associated with PDO phase (Figure 2.5). Between 1900 and ~1945 the PDO is near
38

neutral or slightly positive (warm). This interval corresponds to Δδ
15
N
sed
values that are low,
1.4‰ ± 0.02. Around 1945, the PDO shifted to a strongly negative (cold) phase, which persisted
until 1977. This interval is characterized by the highest values of Δδ
15
N
sed
(1.8‰ ± 0.02). The
reversal in PDO from cold to warm that occurred 1977-1984 is matched by a reversal in the trend
of Δδ
15
N
sed
whereby the difference between isotopic values at Pescadero Slope and SMB
decrease (1.5‰ ± 0.10). A direct comparison of running averages of PDO (1900-1990) to
Δδ
15
N
sed
shows anti-correlation with a 97 percent significance (Figure 2.6a/b), while the same
plots for the time period of 1993-2009 exhibit a positive correlation with 97 and 99 percent
probability (Figure 2.6c/d). It appears that for a major portion of the historical record of PDO
phase changes, the sediment proxy we have measured provides a reasonably coherent and
quantitative metric of this ocean-atmosphere phenomenon. When PDO is positive, more
southern water moves northward and the isotopic values at Pescadero Slope and SMB converge.
When PDO is negative, the difference between these sites is magnified as SMB likely receives
more of its nitrate from northern-sourced waters.

However, since shortly after the last phase change, 1993 to the present, the coupling
between Δδ
15
N
sed
and PDO does not follow the same pattern. The PDO index is declining, while
the Δδ
15
N
sed
value is also declining, opposite to the pattern of the earlier three periods of
increasing California Current influence corresponding to a larger Δδ
15
N
sed
. Moreover, the recent
21 years have been anomalous insofar as δ
15
N
sed
values in the Pescadero region have become
persistently higher. This increase in δ
15
N
sed
is indicative of more intense denitrification in an
expanding OMZ [Altabet et al., 1995; Emmer and Thunell, 2000; Thunell and Kepple, 2004,
Deutsch et al., 2011, 2014]. We note here that the coincident increase in δ
15
N
sed
at both locations
indicate that the change at this time is not due to early diagenesis of δ
15
N
sed
as shown by Deutsch
39

et al. [2014]. The convergence between Pescadero Slope and SMB δ
15
N
sed
values (Figure 2.5a),
in light of both records showing δ
15
N
sed
increase

with time, could be the result of an increase in
the δ
15
N
nitrate
in northern waters. It is improbable that this increase is due to denitrification in the
California Current source water, since oxygen levels in this region are not low enough to support
denitrification. A more plausible reason for the decrease in Δδ
15
N
sed
between the Pescadero
Slope and the SMB is southward transport of partially consumed nitrate with elevated δ
15
N
nitrate

[Waser et al., 1998; Granger et al., 2010]. In addition, this water could also have higher nitrate
concentrations. Some combination of these processes could explain the mixing relationship
observed between the California Undercurrent and California Current in the Santa Monica Basin
from 1993-2009 where there appears to be no correlation between either δ
15
N
sed
or Δδ
15
N
sed
and
salinity (r
2
values of 0.0901 and 0.0893 and p values of 0.2971 and 0.3455 respectively) (Figure
3.4c/d).
A recent study by Rykaczewski and Dunne [2010] examined changes in nutrient supply
and primary production in the California Current Ecosystem (CCE) under projected conditions of
increased greenhouse gas emissions and future climate change. Their study indicated that
reduced meridional gradients in zonal winds and a lower magnitude of wind-stress curl would
occur throughout the Pacific Basin. This would result in basin-wide declines of both
downwelling and upwelling, altering the trajectories of water parcels prior to their arrival in the
CCE. Consequently, water parcels will have prolonged transport, originate from more northerly
locations and from deeper sources when compared to pre-industrial conditions resulting in the
delivery of a greater amount of nitrate to the CCE. The model predicts that water entering the
CCE is sourced from ~40-50°N in the North Pacific, shifted from 30-40°N. Profiles of δ
15
N
nitrate

from the Gulf of Alaska (50°N 145°W) do exhibit
15
N-enrichment (at depths 0-~125 m) when
40

compared to more southern locations in the mid-Pacific [Casciotti et al., 2008]. A change in the
source of California Current waters could lead to a more positive nitrogen isotopic value and
higher concentrations of nitrate. Changes in the nutrient biogeochemistry of water being
transported into the northern CCE may already be taking place (Rykaczewski and Dunne’s ESM
2.1 model calculates a 28% increase in nitrate concentrations from pre-industrial conditions to
the time period of 1981-2000). The breakdown in the relationship between Δδ
15
N
sed
, salinity,
and PDO we observe over the last 20 years could thus be explained by this change in gyre
circulation.
Other environmental and analytical factors could also contribute to the patterns observed
in Δδ
15
N
sed
. Widespread and regional changes in pH and oxygen concentration have been
observed for the ETNP and off Southern California (Feely et al., 2008; Bograd et al., 2008;
Booth et al., 2014). However, since oxygen declines off southern California have not yet
reached the water column denitrification threshold (< 5 µM), this will not likely alter the
δ
15
N
nitrate
value of thermocline waters. El Nino/La Nina can also have a large impact on water
properties off Baja and Southern California. However, a consistent relationship between ENSO
fluctuations and Δδ
15
N
sed
between 1900 and present is not observed (Figure 2.7). Analytically,
uncertainty in the age models, which increases downcore (Figure 2.2), could also shift Δδ
15
N
sed

patterns relative to PDO. To account for this, we calculated Δδ
15
N
sed
vs. time with uncertainties
of ±10% of the calculated sediment accumulation rates and compared these to patterns in the
PDO index (Figure 2.8). In all cases, the data exhibit similar trends to the model we developed
assuming our best estimates of MARs, thus conclude that the patterns are robust through a range
of age model uncertainty.

41

2.5. Conclusions

Values of δ
15
N
sed
determined in laminated coastal marine sediments provide insight into
changes in the intensity of denitrification, and therefore, changes in the intensity of the Oxygen
Minimum Zone (OMZ) in locations characterized by low-oxygen and high sediment
accumulation rates. High-resolution records from the Pescadero Slope and the Santa Monica
Basin indicate a decline in the intensity of the OMZ from 1900 to 1990, followed by a significant
increase in denitrification from 1993 to present. Fluctuations in the difference between δ
15
N
sed

values at Pescadero Slope and Santa Monica Basin, a parameter termed Δδ
15
N
sed
, offers an
intriguing new proxy for assessing changes in the degree of transport and mixing between water
masses with distinctive nitrogen biogeochemistries. This investigation illustrates the transport of
15
N-enriched nitrate from the Eastern Tropical North Pacific (ETNP) into the California
Borderland correlates with phase changes of the Pacific Decadal Oscillation (PDO) over three
distinct PDO transitions during 1900-1993. During PDO positive (warm), the difference
between sedimentary nitrogen isotope values at Pescadero Slope and Santa Monica Basin are
minimized, suggesting stronger California Undercurrent influence. During PDO negative (cold),
Δδ
15
N
sed
values increase, indicative of greater California Current influence on δ
15
N
nitrate
in
Southern California sub-thermocline waters. However, the relationship between PDO and
Δδ
15
N
sed
changes post-1993 deviating from patterns in the preceding intervals. We interpret this
recent change as reflecting changes in circulation and sources of waters feeding the California
Current.

42

Acknowledgements
This work has been published in Geophysical Research Letters with Will Berelson and
Maria Prokopeko as coauthors [Tems, C. E., W.M. Berelson, and M.G. Prokopenko (2015),
Particulate δ
15
N in laminated marine sediments as a proxy for mixing between the California
Undercurrent and the California Current: A proof of concept, Geophysical Research Letters, 42,
doi:10.1002/ 2014GL061993].
This study was supported by a National Science Foundation grant (OCE-0727123 to
W.B.). We acknowledge the use of publicly available salinity data from the CalCOFI Program
(http://www.data.calcofi.org/ctddata.html) from 1952 through 2010 (Figure 3), Pacific Decadal
Oscillation data sourced from the UK Historical SST data set for 1900-81, Reynold’s Optimally
Interpolated SST (V1) for January 1982-Dec 2001, and OI SST Version 2 (V2) beginning
January 2002-present compiled by the University of Washington
(http://jisao.washington.edu/pdo/PDO.latest) (Figure 2.5 and Figure 2.7), and extended
multivariate ENSO index data from: Wolter, K., and M. S. Timlin, (2011) El Niño/Southern
Oscillation behaviour since 1871 as diagnosed in an extended multivariate ENSO index
(MEI.ext). International Journal of Climatology, 31(7), 1074-1087, doi:10.1002/joc.2336
(http://www.esrl.noaa.gov/psd/enso/mei.ext/table.ext.html) (Figure 2.8). Additionally, we
would like to thank James McManus, Nick Rollins, Donn Gorsline, Doug Hammond, Oscar
Gonzalez-Yajimovich, Curtis Deutsch, Robert Thunell, Alexander van Geen, UC Davis Stable
Isotope Facility, Captain and Crew of the RV Horizon, and Captain and Crew of the RV
Yellowfin.

43

References
Altabet, M.A., and R. Francois (1994), Sedimentary nitrogen isotopic ratio as a recorder for
surface ocean nitrate utilization, Global Biogeochemical Cycles, 8(1), 103-116,
doi:10.1029/93GB03396.
Altabet, M.A., R. Francois, D.W. Murray, and W.L. Prell (1995), Climate-related variations in
denitrification in the Arabian Sea from sediment
15
N/
14
N ratios, Nature, 373(6514), 506-509,
doi:10.1038/373506a0.
Altabet, M.A., C. Pilskaln, R. Thunell, C. Pride, D. Sigman, F. Chavez, and R. Francois (1999),
The nitrogen isotope biogeochemistry of sinking particles from the margin of the Eastern
North Pacific, Deep-Sea Research I, 46, 655-679, doi:10.1016/S0967-0637(98)00084-3.
Barnett, P.R.O., J. Watson, and D. Connelly (1984), A multiple corer for taking virtually
undisturbed samples from shelf, bathyal and abyssal sediments, Oceanologica Acta,
7(4), 399-408.
Barron, J.A., and D. Bukry (2007), Solar forcing of Gulf of California climate during the past
2000 yr suggested by diatoms and silicoflagellates, Marine Micropaleontology, 62, 115-139,
doi:10.1016/j.marmicro.2006.08.003.
Behl, R.J., and J.P. Kennett (1996), Brief interstadial events in the Santa Barbara basin, NE
Pacific, during the past 60 kyr, Nature, 379(6562), 243-246, doi:10.1038/379243a0.
Bograd, S.J, C.G. Castro, E. Di Lorenzo, D.M. Palacios, H. Bailey, W. Gilly, and F.P. Chavez
(2008), Oxygen declines and the shoaling of the hypoxic boundary in the California Current,
Geophysical Research Letters, 35(12), L12607.
Booth, J.A.T., C.B. Woodson, M. Sutula, F. Micheli, S.B. Weisberg, S.J. Bograd, A. Steele, J.
Schoen,7 and L. B. Crowder (2014), Patterns and potential drivers of declining oxygen content
44

along the southern California coast, Limnology Oceanography, 59(4), 1127-1138, doi:
10.4319/lo.2014.59.4.1127.
Brandes, J.A., and A.H. Devol (2002), A global marine-fixed nitrogen isotopic budget: Implications
for Holocene nitrogen cycling, Global Biogeochemical Cycles, 16(4), 1120,
doi:10.1029/2001GB001856.
Casciotti, K.L, T.W. Trull, D.M. Glover, and D. Davies (2008), Constraints on nitrogen cycling
at the subtropical North Pacific Station ALOHA from isotopic measurements of nitrate and
particulate nitrogen, Deep-Sea Research II, 55, 1661-1672, doi: 10.1016/j.dsr2.2008.04.017.
Castro, C.G., F.P. Chavez, and C.A Collins (2001), Role of the California Undercurrent in the
export of denitrified waters from the eastern tropical North Pacific, Global Biogeochemical
Cycles, 15(4), 819-830, doi:10.1029/2000GB001324.
Chong, L.S., M.G. Prokopenko, W.M. Berelson, A. Townsend-Small, and J. McManus (2012),
Nitrogen cycling within suboxic and anoxic sediments from the continental margin of
Western North America, Marine Chemistry, 128-129, 13-25,
doi:10.1016/j.marchem.2011.10.007.
Christensen, C.J., D.S. Gorsline, D.E. Hammond, and S.P. Lund (1994), Non-annual laminations
and expansion of anoxic basin-floor conditions in Santa Monica Basin, California
Borderland, over the past 4 centuries, Marine Geology, 116(3-4), 399-418.
Collins, L.E., W. Berelson, D.E. Hammond, A. Knapp, R. Schwartz, D. Capone (2011), Particle
fluxes in San Pedro Basin, California: A four-year record of sedimentation and physical
forcing, Deep-Sea Research I, 58(8), 898-914, doi:10.1016/j.dsr.2011.06.008.
De Pol-Holz, R. O. Ulloa, L. Dezilau, J. Kaiser, F. Lamy, D. Hebbeln (2006), Melting of the
Patagonian Ice Sheet and deglacial perturbations of the nitrogen cycle in the eastern South
Pacific, Geophysical Research Letters, 33(4), L04704, doi:10.1029/2005GL024477.
45

Deutsch, C., H. Brix, T. Ito, H. Frenzel, and L. Thompson (2011), Climate-forced variability of
ocean hypoxia, Science, 333(6040), 336-339, doi:10.1126/science.1202422.
Deutsch, C. W. Berelson, R. Thunell, T. Webber, C. Tems, J. McManus, J. Crusius, T. Ito, T.
Baumgartner, V. Ferreira, J. Mey, and A. VanGeen (2014), Centennial changes in North
Pacific anoxia linked to tropical trade winds, Science, 345(6197), 665-668,
doi:10.1126/science.1252332.
DeVries, T., C. Deutsch, F. Primeau, B. Chang, and A. Devol (2012), Global rates of water-
column denitrification derived from nitrogen gas measurements, Nature Geosciences, 5(8),
547-550, doi:10.1038/ngeo1515.
Di Lorenzo, E. (2003), Seasonal dynamics of the surface circulation in the Southern California
Current Systems, Deep-Sea Research II, 50, 2371-2388, doi:10.1016/S0967-0645(03)00125-
5.
Douglas, R., O. Gonzalez-Yajimovich, J. Ledesma-Vazquez, and F. Staines-Urias (2007),
Climate forcing, primary production and the distribution of Holocene biogenic sediments in
the Gulf of California, Quaternary Science Review, 26(1-2), 115-129,
doi:101016/j.quascirev.2006.05.003.
Emmer, E., and R. C. Thunell (2000), Nitrogen isotope variations in Santa Barbara Basin
sediments: Implications for denitrification in the eastern tropical North Pacific during the last
50,000 years, Paleoceanography, 15(4), 377–387, doi:10.1029/1999PA000417.
Feely, R.A., C.L. Sabine, J.M. Hernandez-Ayon, D. Ianson, and B. Hales (2008), Evidence for
upwelling of corrosive “acidified’ water onto the continental shelf, Science, 320(5882),
1490-1492, doi: 10.1126/science.1155676.
46

Galbraith, E.D., M. Kienast and the NICOPP working group members (2013), The acceleration
of oceanic denitrification during deglacial warming, Nature Geoscience, 6, 579-584,
doi:10.1038/NGEO1832.
Ganashram, R.S., T.F. Pederson, S.E. Calvert, and J.W. Murray (1995), Large changes in
oceanic nutrient inventories from glacial to interglacial periods, Nature, 376, 755-758,
doi:10.1038/376755a0.
Ganeshram, R.S., T.F. Pedersen, S.E. Calvert, G.W. McNeil, and M.R. Fontugne (2000),
Glacial-interglacial variability in denitrification in the world's oceans: Causes and
consequences, Paleoceanography, 15(4), 361-376, doi:10.1029/1999PA000422.
Gonzalez-Yajimovich, O., R.G. Douglas, D.S. Gorsline (2005), The preserved carbonate record
in Holocene sediments of the Alfonso and Pescadero basins, Gulf of California, Mexico,
Proceedings of the Geologists’ Association, 116, 315-330, doi:10.1016/S0016-
7878(05)80050-1.
Gonzalez-Yajimovich, O., R.G. Douglas, and D. Gorsline (2007), Holocene record of ocean-
climate variation in the Gulf of California: rainfall and productivity, Proceedings of the
Twenty-first annual Pacific Climate Workshop, Asilomar Conference Grounds, California,
Technical Report 77 of the Interagency Ecological program for the San Francisco Estuary,
59-68.
Granger, J., D.M. Sigman, M.M. Rohde, M.T. Maldonado, P.D. Tortell (2010), N and O isotope
effects during nitrate assimilation by unicellular prokaryotic and eukarotic plankton cultures,
Geochimica et Cosmochimica Acta, 74(3), 1030-1040, doi:10.1016/j.gca.2009.10.044.
47

Hagadorn, J.W. (1996), Laminated sediments of Santa Monica Basin, California continental
borderland, Geological Society, London, Special Publications, 116, 111-120,
doi:10.1144/GSL.SP.1996.116.01.11.
Haug, G.H., T.F.Pedersen, D.M. Sigman, S.E. Calvert, B. Nielsen, and L.C. Peterson (1998),
Glacial/interglacial variations in production and nitrogen fixation in the Cariaco Basin during
the last 580 kyr, Paleoceanography, 13(5), 427-432.
Hendy, I.L., and T.F. Pedersen (2006), Oxygen minimum zone expansion in the eastern tropical
North Pacific during deglaciation, Geophysical Research Letters, 33, L20602,
doi:10.1029/2006GL025975.
Hickey, B.M. (1992), Circulation over the Santa Monica-San Pedro Basin and shelf, Progress in
Oceanography, 30, 37-115, doi:10.1016/0079-6611(92)90009-O.
Kemp, A.E.S. (1996), Laminated sediments as palaeo-indicators, Geological Society, London,
Special Publications, 116, vii-xii, doi:10.1144/GSL.SP.1996.116.01.01.
Kienast, S.S., S.E. Calvert, and T.F. Pederson (2002), Nitrogen isotope and productivity
variations along the northeast Pacific margin over the last 120 kyr: surface an subsurface
paleoceanography, Paleoceanography, 17(4), 1055, doi:10.1029/2001PA000650.
Lehmann, M.F., S.M. Bernasconi, A. Barbiere, and J.A. McKenzie (2002), Preservation of
organic matter and alteration of its carbon and nitrogen isotope composition during simulated
and in situ early sedimentary diagenesis, Geochimica et Cosmochimica Acta, 66(20), 3573-
3582, doi:10.1016/S0016-7037(02)00968-7.
Liu, K-K, and I.F. Kaplan (1989), The Eastern Tropical Pacific as a source of 15N-enriched
nitrate in seawater off Southern California, Limnology and Oceanography, 34(5), 820-830.
48

Lynn, R.J., and J.J. Simpson (1987), The California Current System: The seasonal variability of
its physical characteristics, Journal of Geophysical Research, 92(C12), 12,947-12,966.
Mantua, N.J., and S.R. Hare (2002), The Pacific Decadal Oscillation, Journal of Oceanography,
58, 35-44, doi:10.1023/A:1015820616384.
Martinez, P.M., and R.S. Robinson (2010), Increase in water column denitrification during the
last deglaciation: the influence of oxygen demand in the eastern equatorial Pacific,
Biogeosciences, 7, 1-9, doi:10.5194/bg-7-1-2010.
Pride, C., R. Thunell, D. Sigman, L. Keigwin, M. Altabet, and E. Tappa (1999), Nitrogen
isotopic variations in the Gulf of California since the last deglaciation: response to global
climate change, Paleoceanography, 14(3), 397-409, doi:10.1029/1999PA900004.
Prokopenko, M., D. E. Hammond, A. J. Spivack, and L. Stott (2006a), Impact of long-term
diagenesis on δ
15
N of organic matter in marine sediments: Sites 1227 and 1230, Proceedings
from the Ocean Drilling Program, Scientific Results, 201, 30.
Prokopenko, M., D. E. Hammond, and L. Stott (2006b), Lack of isotopic fractionation of d15N
of organic matter during long-term diagenesis in marine sediments, ODP Leg 202, Sites 1234
and 1235, Proceedings from the Ocean Drilling Program, Scientific Results, 202, 22.
Prokopenko, M.G., D.E. Hammond, W.M. Berelson, J. Bernhard, L. Stott, and R. Douglas
(2006c), Nitrogen cycling in the sediments of Santa Barbara basin and Eastern Subtropical
North Pacific: Nitrogen isotopes, diagenesis and possible chemosymbiosis between two
lithothrophs (Thioploca and Anammox – “riding on a glider”, Earth and Planetary Science
Letters, 242(1-2), 186-204, doi:10.1016/j.epsl.2005.11.044.
Prokopenko, M.G., D.M. Sigman, W.M. Berelson, D.E. Hammond, B. Barnett, L. Chong, and A.
Townsend-Small (2011), Denitrification in anoxic sediments supported by biological nitrate
49

transport, Geochimica et Cosmochimica Acta, 75, 7180–7199,
doi:10.1016/j.gca.2011.09.023.
Prokopenko, M.G., M.B. Hirst, L. De Brabandere, D.J.P. Lawrence, W.M. Berelson, J. Granger,
B.X. Chang, S. Dawson, E.J. Crane III, L. Chong, B. Thamdrup, A. Townsend-Small, and
D.M. Sigman (2013), Nitrogen losses in anoxic marine sediments driven by Thioploca-
anammox bacterial consortia, Nature, 500, 194-200, doi:10.1038/nature12365.
Robinson, R.S., M. Kienast, A.L. Albuquerque, M. Altabet, S. Contreras, R. De Pol Holz, N.
Dubois, R. Francois, E. Galbraith, T-C. Hsu, T. Ivanochko S. Jaccard, S-J. Kao, T. Kiefer, S.
Kienast, M. Lehmann, P. Martinez, M. McCarthy, J. Möbius, T. Pedersen, T.M. Quan, E.
Ryabenko, A. Schmittner, R. Schneider, A. Schneider-Mor, M. Shigemitsu, D. Sinclair, C.
Somes, A. Studer, R. Thunell, and J-Y. (2012), A review of nitrogen isotopic alteration in
marine sediments, Paleoceanography, 27, PA2403, doi:10.1029/2012PA002321.
Rykaczewski, R.R., and J.P. Dunne (2010), Enhanced nutrient supply to the California Current
Ecosystem with global warming and increased stratification in an earth system model,
Geophysical Research Letters, 37(21), L21606, doi:10.1029/2010GL045019.
Schimmelmann, A. and M.J. Tegner (1991), Historical oceanographic events reflected in
13
C/
12
C
ratio of total organic carbon in laminated Santa Barbara Basin sediment, Global
Biogeochemical Cycles, 5(2), 173-188.
Schimmelmann A. and C. Lange (1996), Tales of 1001 varves: a review of Santa Barbara Basin
sediment studies, Geological Society, London, Special Publications, 116, 121-141,
doi:10.1144/GSL.SP.1996.116.01.12.
Schimmelmann, A., I.L. Hendy, L. Dunn, D.K. Pak, and C.B. Lange (2013), Revised ∼2000-year
chronostratigraphy of partially varved marine sediment in Santa Barbara Basin, California,
50

GFF Special Issue: Varve Genesis, Chronology and Paleoclimate, 135(3-4),
doi:10.1080/11035897.2013.773066.
Sigman, D. M., R. Robinson, A. N. Knapp, A. van Geen, D. C. McCorkle, J. A. Brandes, and R.
C. Thunell (2003), Distinguishing between water column and sedimentary denitrification in
the Santa Barbara Basin using the stable isotopes of nitrate, Geochemistry, Geophysics,
Geosystems, 4(5), 1040, doi:10.1029/2002GC000384.
Sigman, D.M., J. Granger, P.J. Difiore, M.M. Lehmann, R. Ho, G. Cane, and A. Van Geen
(2005), Coupled nitrogen and oxygen isotope measurements of nitrate along the eastern
North Pacific margin, Global Biogeochemical Cycles, 19(4), GB4022,
doi:10.1029/2005GB002458.
Smith, K. L., H.A. Ruhl, M. Kahru, C. L. Huffard, and A.D. Sherman (2013), Deep ocean
communities impacted by changing climate over 24 y in the abyssal northeast Pacific Ocean,
Proceedings of the National Academy of Sciences, 110(49), 19838-19841,
doi:10.1073/pnas.1315447110.
Staines-Urías, F., R.G. Douglas, and D.S. Gorsline (2009), Oceanographic variability in the
southern Gulf of California over the past 400 years: Evidence from faunal and isotopic
records from planktic foraminifera, Palaeogeography, Palaeoclimatology, Palaeoecology,
284(3-4), 337-354, doi:10.1016/j.palaeo.2009.10.016.
Thunell, R.C., and A.B. Kepple (2004), Glacial-Holocene δ
15
N record from the Gulf of
Tehuantepec, Mexico: Implications for denitrification in the eastern equatorial Pacific and
changes in atmospheric N
2
O, Global Biogeochemical Cycles, 18(1), GB1001,
doi:10.1029/2002GB002028.
51

Thunell, R.C., D.M. Sigman, F. Muller-Karger, Y. Astor, and R. Varela (2004), Nitrogen isotope
dynamics of the Cariaco Basin, Venezuela, Global Biogeochemical Cycles, 18(3), GB3001,
doi:10.1029/2003GB002185.
Townsend-Small, A., M.G. Prokopenko, and W.M. Berelson (2014), Nitrous oxide cycling in the
water column and sediments of the oxygen minimum zone, eastern subtropical North Pacific,
Southern California, and Northern Mexico (23°N–34°N), Journal of Geophysical Research:
Oceans, 119, 3158–3170, doi:10.1002/2013JC009580.
Van Geen, A., Y. Zheng, J. M. Bernhard, K. G. Cannariato, J. Carriquiry, W. E. Dean, B. W.
Eakins, J. D. Ortiz, and J. Pike (2003), On the preservation of laminated sediments along the
western margin of North America, Paleoceanography, 18(4), 1098,
doi:10.1029/2003PA000911.
Wasser, N.A.D., P.J. Harrison, B. Nielsen, and S.E. Calvert (1998), Nitrogen isotope
fractionation during uptake and assimilation of nitrate, nitrite, ammonium, and urea by a
marine diatom, Limnology and Oceanography, 43(2), 215-224,
doi:0.4319/lo.1998.43.2.0215.
Wolter, K., and M. S. Timlin (2011), El Niño/Southern Oscillation behaviour since 1871 as
diagnosed in an extended multivariate ENSO index (MEI.ext). International Journal of
Climatology, 31(7), 1074-1087, doi:10.1002/joc.2336.

52
1 cm
. b . a
Figure 2.1. a) X-radiograph of the top of a Santa Monica multicore; clear laminations are present.
b) a high-resolution scan of a section of the Pescadero multicore, showing millimeter-scale
laminations. Laminations are preserved at a higher resolution on the Pescadero Slope than in the
Santa Monica Basin, nonetheless, their presence in both cores indicate anoxic
conditions at the time of deposition.
53
Integrated Mass (g ● cm
-2
)
Total
210
Pb (Bq kg
-1
)
y = 44.822 + 712.7 * e
(-0.4015*x)
r
2
= 0.9856
P = <0.001
0
10
8
6
4
2
600 200 500 300 700 100 400 0
Figure 2.2. Total
210
Pb plotted against the integrated mass (calculated from porosity depth profiles)
from a multicore collected at the Pescadero Slope. An exponential decay equation
(y = m1 + m2*exp(-m3*x)) was fit to the data using KaleidaGraph 4.03 to attain the following
equation: y = 44.82 + 712.7*exp(-0.4015*x)) with an r
2
value of 0.986 and an ANOV A p value
<0.001. Axes appear reversed in the figure. The sediment accumulation rate was found to be
77.42 mg cm
-2
yr
-1
, which was calculated by dividing the decay rate of
210
Pb by the m3 value
(0.4015).
54
1750 1800 1850 1900 1950 2000
6.5
7
7.5
8
8.5
9
9.5
10
10.5
Year AD

PS MAR 77.42 mg cm
-2
yr
-1
PS MAR 69.67 mg cm
-2
yr
-1
PS MAR 85.16 mg cm
-2
yr
-1
SMB MAR 16 mg cm
-2
yr
-1
SMB MAR 18 mg cm
-2
yr
-1
SMB MAR 14 mg cm
-2
yr
-1
δ
15
N
sed
(‰)
Figure 2.3. δ
15
N
sed
records from the Pescadero Slope (PS) (black) and the Santa Monica Basin (SMB) (gray)
plotted against year of deposition. Solid bold lines represent the age model calculated year, while dashed and
thin lines represent ±10% of the calculated Mass Accumulation Rate (MAR).
55
1993-2010 1993-2009
1952-1990 1952-1990
34.00 34.05 34.10 34.15
0.5
1.0
1.5
2.0
Salinity

y = -3.663*x+126.4
r
2
= 0.0893
p = 0.3455
34.00 34.05 34.10 34.15
0.5
1.0
1.5
2.0
Salinity

y = -3.416*x+118.2
r
2
= 0.4691
p = 0.0098
34.00 34.05 34.10 34.15
6
7
8
9
Salinity

y = 6.923*x-228.5
r
2
= 0.0901
p = 0.2971
δ
15
N
sed
(
0
/
00
)
Δδ
15
N
sed
(
0
/
00
) Δδ
15
N
sed
(
0
/
00
)
34.00 34.05 34.10 34.15
6
7
8
9
Salinity

y = 3.410*x-109.0
r
2
= 0.3898
p = 0.0170
δ
15
N
sed
(
0
/
00
)
Santa Monica Basin (SMB)
a.
c.
b.
d.
Pescadero Slope - SMB
Figure 2.4. a) δ
15
N
sed
from SMB plotted against salinity from the CalCOFI data from 1952-1990.
b) Δδ
15
N
sed
between the Pescadero Slope and the Santa Monica Basin (SMB) plotted against
salinity from CalCOFI surveys from 1952-1990. c) δ
15
N
sed
from SMB against salinity from the
CalCOFI data from 1993-2010. d) Δδ
15
N
sed
between the Pescadero Slope and the SMB plotted
against salinity from 1993-2009. P-values are calculated through ANOV A analysis.
56
1900 1920 1940 1960 1980 2000
0.5
1
1.5
2
2.5
Year
1900 1920 1940 1960 1980 2000
-2
0
2
Year
PDO Index
a.
b.
∆δ
15
N
sed
(‰)
Figure 2.5. a) Δδ
15
N
sed
between the Pescadero Slope and Santa Monica Basin from 1900-2009.
Uncertainty is ±0.12‰. The horizontal dashed line is the mean value for entire time period. Dashed
vertical lines separate Pacific Decadal Oscillation (PDO) phase periods. b) Annual average of the
PDO index (data from http://jisao.washington.edu/pdo/PDO.latest). Also included is a five-year
running average (gray line). Horizontal dashed lines separate PDO positive from negative and
dashed vertical lines separate PDO phases.
57
1
0
-1
0.95 1.35 1.75 2.15
0.95 1.35 1.75 2.15
1
0
-1
1
0
-1
1
0
-1
0.95 1.35 1.75
0.95 1.35 1.75
PDO Index
PDO Index PDO Index
PDO Index
1900 - 1990 (5-year Running Average of PDO)
1993 - 2009 (5-year Running Average of PDO)
1900 - 1990 (10-year Running Average of PDO)
1993 - 2009 (10-year Running Average of PDO)
Δδ
15
N
sed
(‰)
Δδ
15
N
sed
(‰)
Δδ
15
N
sed
(‰) Δδ
15
N
sed
(‰)
y = -0.5537x + 0.9284
r
2
= 0.1070
p = 0.0393
y = 0.8694x - 1.095
r
2
= 0.3142
p = 0.0314
y = -0.5502x + 0.9284
r
2
= 0.1105
p = 0.0361
y = 0.6495x - 0.085
r
2
= 0.4934
p = 0.0109
a.
b.
c.
d.
Figure 2.6. Cross plots of PDO index vs. Δδ
15
N
sed
. To show the robustness of the fit, an ANOV A
p-value has been calculated and appears on each subplot. a) A 5-year running average of the
PDO index is compared to Δδ
15
N
sed
from 1900-1990 and indicates anti-correlation. b) A 5-year
running average of the PDO index is compared to Δδ
15
N
sed
from 1993-2009, and indicates a
change in correlation (to positive correlation opposed to anti-correlation). c) A 10-year running
average of the PDO index vs. Δδ15Nsed from 1900-1990 is anti-correlated, similar to trends seen
in (a). d) A 10-year running average of the PDO index vs. δ
15
N
sed
from 1993-2009 shows
correlation during the time period between 1993-2009, similar to the trends in (b).
58
1900 1920 1940 1960 1980 2000
0.5
1
1.5
2
2.5
1900 1920 1940 1960 1980 2000
-2
-1
0
1
2
Δδ
15
N
sed
(‰) ENSO Index
a.
b.
Year
Year
Figure 2.7. a) Δδ
15
N
sed
between the Pescadero Slope and SMB from 1900-2009.
Uncertainty plotted as error bars is ±0.12‰. The black arrows indicate significant El Niño
events. b) Annual average of the Multivariate ENSO index (MEI) [Wolter, K., and M. S. Timlin
(2011), El Niño/Southern Oscillation behaviour since 1871 as diagnosed in an extended multi
variate ENSO index (MEI.ext). International Journal of Climatology, 31(7), 1074-1087,
doi:10.1002/joc.2336]. Also included is a five-year running average (gray line).
59
1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 2010
-2
0
2
PDO Index
1900 1920 1940 1960 1980 2000 2010
0.5
1.75
2.75
δΔ
15
N
sed
(
o
/
oo
)

1900 1920 1940 1960 1980 2000 2010
0.5
1.75
2.75

1900 1920 1940 1960 1980 2000 2010
0.5
1.75
2.75

SMB MAR 18 mg cm
-2
yr
-1
SMB MAR 14 cm
-2
yr
-1
SMB MAR 16 mg cm
-2
yr
-1
PS MAR 85.19 mg cm
-2
yr
-1
PS MAR 69.71 mg cm
-2
yr
-1
PS MAR 77.42 mg cm
-2
yr
-1
PS MAR 69.71 mg cm
-2
yr
-1
SMB MAR 18 mg cm
-2
yr
-1

PS MAR 85.19 mg cm
-2
yr
-1

SMB MAR 14 mg cm
-2
yr
-1
PS MAR 77.42 mg cm
-2
yr
-1

SMB MAR 16 mg cm
-2
yr
-1
Year
δΔ
15
N
sed
(
o
/
oo
)
Year
δΔ
15
N
sed
(
o
/
oo
)
Year
Year
Yearly Average 5-Year Running Average
a.
b.
c.
d.
Figure 2.8. The top three panels show Δδ
15
N
sed
between the Pescadero Slope and
SMB plotted against year. a) The mass accumulation rate (MAR) of the Pescadero
Slope is kept constant (at 77.4 mg cm-2yr-1) while the MAR of SMB is allowed to
deviate by ±12.5% from the assumed MAR of 16 mg cm-2yr-1. b) The MAR of the
SMB is kept constant (at 16 mg cm-2yr-1) while the Pescadero Slope MAR is
allowed to deviate by ±10%. c) Δδ15Nsed is compared between sites where MAR of
both sites deviates in opposite directions, with the black line representing the MAR
calculated from the assumed age model. d) Yearly average of the Pacific Decadal
Oscillation (PDO) index. Also included is a five-year running average (gray line).
60

Chapter 3
Sedimentary δ
15
N reveal decadal fluctuations in the intensity of the eastern
tropical north Pacific oxygen minimum zone during the last 1200 years

Abstract
Oxygen minimum zones (OMZs), located below highly productive marine regions, are
sites of microbially-mediated denitrification and biogeochemical cycling that have global
significance. The intensity of OMZs fluctuate naturally, however, the degree of these
fluctuations and a comprehensive understanding of the factors that drive these fluctuations on
interannual to centennial time scales is lacking. Our high-resolution (near annual) record of
δ
15
N
sed
from the Pescadero Slope in the Gulf of California (eastern tropical north Pacific, ETNP)
suggests that the OMZ is self-regulating, capped by maximum (10.5‰) and minimum (8.0‰)
δ
15
N
sed
values which create hard ceilings and floors between which OMZ intensity has varied
over the past 1200 years. A comparative analysis of the relationship between δ
15
NO
3
-
and [O
2
] in
Pescadero and nearby sites suggests that the observed range of δ
15
N
sed
values is equivalent to a 45
µM fluctuation in O
2
content and that these changes can occur in less than 40 years. Our findings
show that the OMZ typically intensifies quickly and contracts gradually; the average rate of
OMZ intensification (-1.2 µM O
2
/year) is twice as fast as the rate of OMZ re-oxygenation.
Spectral analyses suggest Gleissberg and Suess (deVries) solar cycles, solar irradiance, and the
Pacific Decadal Oscillation could influence the internal variability in the intensity of the OMZ
during the late Holocene. Intensity fluctuations are also associated with changes in productivity
as evidenced by a significant positive correlation between δ
15
N
sed
and weight percent C
org
. ETNP
OMZ intensity regulation has significant implications for nitrogen cycling on a global scale.
61

3.1. Introduction
Oceanic oxygen minimum zones (OMZs) are mid-water features that are associated with
expansive, highly productive regions of the ocean and are also sites of microbially-mediated
biogeochemical cycling of redox sensitive compounds [Gilly et al., 2013; Keeling et al., 2010].
These regions have important implications for nitrogen loss (through denitrification and
anammox), ecological relationships (predators and their prey and the range of hypoxia sensitive
organisms), marine economic resourses and ocean management since the world’s most
productive fisheries are found above OMZs. In the geologic record, variations in dissolved
oxygen concentrations [O
2
] and OMZ intensity are hypothesized to have triggered the end-
Permian mass extinction [Benton and Twitchett, 2003; Wignall and Twitchett, 1996] and the
great ocean anoxia events in the Cretaceous period [Jones and Jenkyns, 2001], which had
devastating impacts on marine communities and changed the cycling of nutrients and trace
metals on regional and global scales. The magnitude and intensity of OMZs has been shown to
vary on glacial-interglacial time scales [Altabet et al., 1995; Ganeshram et al., 2000; Deutsch et
al., 2004; Thunell and Kepple, 2004], suggesting a relationship between the climate system and
the expanse of OMZs. This connection, however, has not been studied at finer temporal
resolution in the ETNP during an interglacial period and knowledge of how OMZs fluctuate on a
near annual basis and the exact drivers responsible for the changes are lacking. Quantifying the
degree to which OMZ intensity fluctuates and the time scales of these variations is critical in
assessing how they may change in response to future changes in climate [Deutsch et al., 2014].
A common paradigm is that as the climate warms, O
2
becomes less soluble in seawater
and the upper ocean becomes more stratified, reducing ventilation, with both of these processes
resulting in a decline in dissolved O
2
in the ocean interior [Keeling et al., 2010]. In the eastern
62

tropical Pacific, a 10-20% decline in dissolved O
2
has been observed in data collected over the
last several decades [Stramma et al., 2008; Bograd et al., 2005; McClatchie et al., 2010] and a
50-year time series of dissolved O
2
also reveals a recent expansion of OMZs globally beginning
~1990 [Stramma et al., 2008]. This has prompted investigations examining fluctuations in the
intensity of the eastern tropical north Pacific (ETNP) OMZ using paleoceanographic proxies to
evaluate decadal and centennial trends over longer time periods [Deustch et al., 2014; Tems et al.,
2015]. Recent studies indicate that over the last 150 years the ETNP OMZ has been contracting,
not expanding, with the exception of the last 20 years. Deutsch et al. [2014] attribute this
contraction to a reduction in the intensity of the trade winds in a warming climate, which results
in a deepening of the low latitude thermocline. The study presented here builds on the work of
Deustch et al. [2014] and Tems et al. [2015] to evaluate if the recent trends and range of OMZ
intensity variations are representative of the ETNP on centennial and millennial scales and what
additional factors may influence fluctuations in the OMZ on these longer time scales.
The δ
15
N of bulk sedimentary organic matter (where δ
15
N
sed
‰ =
[(
15
N/
14
Nsample)/(
15
N/
14
Nstandard) - 1] * 1000, and the standard is atmospheric N
2
) is used as a
proxy to monitor changes in water column denitrification. Water column denitrification is tightly
coupled to OMZ oxygen concentration along the Baja California margin [Sigman et al., 2005].
The δ
15
N
sed
proxy is thus a useful tracer of OMZ conditions if complete nitrate utilization occurs
in the photic zone [Altabet and Francios, 1994], sedimentation rates are high [Thunell et al.,
2004], bottom water [O
2
] is low [Prokopenko et al., 2006], if δ
15
N
sed
is not diagenetically altered
[Altabet et al., 1999; Lehmann et al., 2002; Thunell et al., 2004; Prokopenko et al., 2006,
Deutsch et al., 2014; Tems et al., 2015], and nitrite oxidation is not significantly influencing the
63

isotopic signature [Casciotti et al., 2013]. As discussed in Deutsch et al. [2014] and Tems et al.
[2015] we believe all these criteria are met at the Pescadero Slope site.
To understand the factors influencing ETNP OMZ fluctuations over the past 1200 years,
we delve into an assessment of primary productivity proxies and examine δ
15
N
sed
temporal
patterns. Identifying periodicities in the δ
15
N
sed
record through spectral analysis provides a
quantitative framework for evaluating mechanisms controlling changes in the OMZ. Changes in
oceanic primary productivity are commonly inferred from the accumulation of biogenic
components, which include diatoms and other siliceous phytoplankton in tropical Pacific
upwelling zones [Ragueneau et al, 2000]. In these regions, biogenic silica accumulation can be
used to establish a lower limit of primary productivity and carbon export [Pichevin et al., 2012].
We hypothesize that when δ
15
N
sed
is elevated due to increased denitrification, more POM would
have been produced and exported out of the photic zone and degraded in the OMZ. To test this
hypothesis, we present records of weight percent (wt. %) organic carbon (C
org
)

and wt. %
biogenic silica (bSi) in the sediments as proxies for primary productivity and directly compare
these records to δ
15
N
sed
over the past 1200 years.

3.2. Study Site
This study focuses on cores collected from the slope of the Pescadero Basin in the mouth
of the Gulf of California (Figure 3.1). The Gulf of California is an elongate basin, surrounded by
land to the north, west, and east, and opens to the Pacific Ocean in the south. In the fall, the
Pacific High-Pressure System and the Inter-tropical Convergence Zone (ITCZ) migrate south
towards the equator and the winds intensify. As this occurs, the strength and prominent direction
of the winds change from weak southeastern winds to strong northwestern winds [Douglas et al.,
64

2007]. The strong northwestern winds cause Ekman transport in the Gulf, resulting in intensified
upwelling and stimulating high productivity on the eastern side of the Gulf from November to
March [Staines-Urías et al., 2009; Douglas et al., 2007]. Biogenic silica fluxes in the Gulf are
highest during this time of year [Thunell et al., 1994]. High productivity coupled with moderate
rates of deep thermohaline ventilation result in the formation of a well-developed oxygen
minimum zone between 100 and 800 meters [Fernández-Barajas et al., 1994]. The OMZ
impinges on the basin slope where organic-rich (2-5 wt. % C) laminated sediments with elevated
15
N/
14
N are found. C:N ratios of 9–11, a zero intercept plot of weight percent nitrogen versus
weight percent carbon [Tems et al., 2015], and particulate organic matter (POM) δ
13
C values of
-22.1 to -21.2‰ all indicate that the signal is predominantly marine organic matter. The location
of Pescadero Slope is not only influenced by local dynamics associated with the Gulf of
California monsoon but is also influenced by regional Pacific Ocean circulation and the ETNP
OMZ, making it an ideal location to study local as well as regional variability.

3.3 Sampling and Analytical Methods
3.3.1 Core collection
One multicore (PESC-MC1, 64.5 cm) and two gravity cores (PESC-GC1, 137.7 cm, and
PESC-GC3, 204.5 cm) were collected in 2009 from the slope of Pescadero Basin. PESC-MC1
was collected at 24°16.88’N and 108°11.79’W in 616 m water depth. PESC-GC1 and PESC-
GC3 were collected at 24°16.160’N, 108°11.599’W and 24°16.759’N, 108°11.700’W and at 601
m and 620 m, respectively. Bottom water [O
2
] is ~0.4µM at these locations [Tems et al., 2015].
Water column and sediment pore water chemistry at this site is described in Chong et al. [2012],
Prokopenko et al. [2011], and Townsend-Small et al. [2014].
65

3.3.2 Age model
The multicore, PESC-MC1, preserved the sediment-water interface and the deposition of
the most recent sediments. To align the δ
15
N
sed
records from PESC-MC1 and PESC-GC3, weight
percent total carbon, δ
13
C
sed
total carbon, and paleomagnetic data were used to correlate PESC-
MC1 to PESC-GC1 and PESC-GC1 to PESC-GC3 (Figure 2.2). The age model for the cores
was constructed by combining and cross-correlating
210
Pb measurements (obtained by gamma
spectroscopy), accelerator mass spectrometry AMS
14
C measurements, and varve counts. The
210
Pb age model, developed by plotting integrated mass and
210
Pb and fitting an exponential
decay regression to the data [Deutsch et al., 2014; Tems et al., 2015], resulted in a sediment
accumulation rate of 77.4 mg

cm
-2
yr
-1
, which was applied to the first 135 years of PESC-MC1.
The age model for the remainder of PESC-MC1 and all of PESC-GC3 is based on a single linear
regression of calibrated
14
C calibrated ages.
The
14
C age model is based on the analysis of 26 samples of sedimentary organic matter.
The samples were pre-treated with an acid-base-acid washing procedure [Olsson, 1986],
graphitized [Xu et al., 2007] and analyzed for organic matter AMS dating [Southon and Santos,
2004, 2007] at the Keck Carbon Cycle AMS Facility at the University of California (UC) Irvine.
Due to poor preservation of foraminifera in the cores only one bulk benthic foraminifera date
was measured and showed an offset with respect to the age of organic mater similar to the
average global ocean reservoir age (450 ± 20 years). Additional samples were also measured
without undergoing acid/base pre-treatment procedures and dates of treated and non-treated
samples of the same interval were found to be within analytical error of each other. The
14
C
measurements were corrected for preparation backgrounds and isotopic fractionation according
to conventions of Stuiver and Polach [1977] and a marine correction using Calib 7.1 [Stuiver and
66

Reimer, 1993], including a correction for local reservoir age (relative to the global ocean
reservoir age, ΔR) of 508 ± 30 years was applied (Table 3.1). This correction was estimated by
averaging the radiocarbon ages obtained from PESC-MC1 older than 1950 AD (n=3) after each
was corrected by the global reservoir age for the year the sediment was deposited as determined
by the
210
Pb age model. The global reservoir age was determined from Marine13 [Reimer et al.,
2013] and error in ΔR was calculated as the standard deviation between the average reservoir
corrections from PESC-MC1. Independent analysis of radiocarbon ages obtained from PESC-
GC3 suggested a reservoir age within error of that calculated for PESC-MC1. The calibrated
ages were compared to the integrated mass (related to depth in the core by porosity, which
corrects for the effect of compaction) and fit with a linear model (r
2
= 0.98, P < 0.001), which
yielded an average sediment accumulation rate of 77.4 mg cm
-2
yr
-1
(Figure 3.3a). This value is
identical to the sedimentation rate derived from the
210
Pb analysis of PESC-MC1. The reservoir
age correction determined from the post-1950 sediments was assumed to be constant and apply
through the depth of the gravity core.
The validity of the radiocarbon age model is further supported by laminae counts. We
counted 796 laminae couplets, defined as one light and one dark band as viewed on high-
resolution images and x-radiographs, in what is determined by our radiometric isotope age model
to be 721 years (Figure 3.3b), which illustrates that laminae couplets are annual (±10%).
Multiple varve counts were completed and the counting was reproducible within 1.5% (12 years).

3.3.3 Sediment Chemistry Analytical Techniques
PESC-MC1, PESC-GC1, and PESC-GC3 were split lengthwise and sampled at 3 mm
resolution. PESC-MC1 and PESC-GC1 were digitally scanned at high-resolution and PESC-GC3
67

was x-radiographed confirming the presence of laminated sediments throughout all three cores.
Measurements of δ
15
N, percent total organic carbon (%TOC) and weight percent biogenic silica
were made on sediment splits from PESC-MC1 and PESC-GC3. Measurements of total carbon
δ
13
C and wt. % total carbon were made on sediment splits from PESC-MC1, PESC-GC1, and
PESC-GC3 to correlate the three cores.
Individual samples were prepared for bulk δ
15
N
sed
by drying and grinding to a fine
powder. The samples from PESC-MC1 were analyzed with a PDZ
Europa
ANCA-‐GSL
elemental
analyzer interfaced to a PDZ
Europa
20-‐20
continuous flow isotope ratio mass spectrometer
(IRMS) at the UC Davis Stable Isotope Facility. The references used to normalize the data were
G-11 (nylon, δ
15
N of -9.77‰), G-12 (glutamic acid enriched, δ
15
N of 45.31‰), G-13 (bovine
liver, δ
15
N of 7.72‰), and G-9 (glutamic acid, δ
15
N of-4.26‰). PESC-GC3 samples were
measured for δ
15
N
sed
at the University of South Carolina on a Euro Elemental Analyzer
elemental interfaced to a GV Isoprime continuous flow IRMS. The reference standards used to
normalize the data were N-1 (δ
15
N = 0.4‰), N-2 (δ
15
N = 20.41‰), N-3 (δ
15
N = 4.7‰), and
USGS-40 (δ
15
N = -4.52‰). An inter-laboratory calibration showed that the measurements at
each facility were within analytical error of each other. Uncertainty was determined by
averaging the standard deviation between replicate samples (n=15) and was calculated to be
0.17‰.
Biogenic silica (bSi) analyses were run at the University of Southern California (USC)
using a hot sodium bicarbonate leaching method [DeMaster, 1979; DeMaster, 1991; Strickland,
1968]. For each sample, ten to twelve milligrams of powdered sediment were measured into 50
mL plastic centrifuge tube and 50 mL of 5% (weight/volume) sodium bicarbonate was added.
The centrifuge tube was then placed in an 80°C bath for 5 hours. Each sample was mixed every
68

30 minutes and 0.5 mL aliquot subsamples were taken at 3, 4, and 5 hour intervals. Each
subsample was neutralized with 0.235 M hydrochloric acid and analyzed colorimetrically on a
spectrophotometer. Measured samples were corrected to account for dilution factors and
changes in volume in the centrifuge tube due to sampling. The corrected concentration was
computed for each subsample time point and fitted with a linear regression. The y-intercept was
used to calculate the biogenic silica concentration and subsequently the weight percent biogenic
silica. Error was calculated by averaging the standard deviation between replicate samples and
was found to be ±0.15% as wt. % bSi.
Five to eight milligrams of powdered sediment were packaged into tin capsules and were
analyzed for total carbon δ
13
C and weight percent total carbon at USC on a Costech ECS 4010
elemental analyzer interfaced to a Picarro G2131-i cavity ring down spectrometer. For this
analysis the standard used was L-glutamic acid (USGS 40) with δ
13
C
VPDB
= -26.39± 0.04‰ and a
carbon mass fraction of 40.8%. Weight percent inorganic carbon was also measured and used to
calculate weight percent organic carbon. The inorganic carbon measurements were made on the
Picarro G2131-i cavity ring down spectrometer using an Automate device to acidify the sample.
For this analysis, 14 to 335 mg of powdered sample were placed in a glass test tube, which was
evacuated using a vacuum pump, and pre-acidified with 1 mL of 10% phosphoric acid (from
which CO
2
had been removed by bubbling the solution with N
2
gas for 20 minutes). When
analyzed, the Automate injects an additional 3 mL of 10% phosphoric acid and the CO
2
that is
released from the sample is collected and analyzed using the Picarro. The standards used for the
inorganic carbon analyses were OPT Calcite (δ
13
C
VPDB
= 2.47 ± 0.01 and 12.002% C) and AR15
(δ
13
C
VPDB
= -9.65 ± 0.03). Weight percent organic carbon was calculated as the difference
between wt. % total and wt. % inorganic C.
69

Due to very small amounts of inorganic carbon in many samples, an alternative acid
fumigation method was also used to measure wt. % C
org
. This method involved weighing 7-10
mg of ground sediment into silver capsules, acidifying the samples by adding 0.15 mL of
deionized water to each sample and placing the samples in a desiccator containing a beaker with
50 mL of reagent grade HCl for 8 hours, after which the samples where placed in an oven to dry,
packaged in an additional tin cup, and measured on a Costech ECS 4010 elemental analyzer
interfaced to a Picarro G2131-i. For this analysis the standard used was L-glutamic acid (USGS
40, 40.8% mass fraction C). The two methods for measuring wt. % C
org
and δ
13
C
org
are within
error of each other.
We assessed the concentration of magnetic particles, their grain size distribution, and
variations in the percentage of various magnetic minerals in order to accurately match PESC-
MC1 and PESC-GC1. Such measurements have been used routinely [Bloemendal and King,
1988; King and Channell, 1991] as proxies for variations in bulk clastic sedimentology (clastic
concentration, grain size, mineralogy). For this analysis, PESC-MC1 and PESC-GC1 were
sampled discretely and contiguously using 2-cm-cube samples. The bulk magnetic susceptibility
(chi) of each sample was initially measured at USC. Two artificial remanences were then applied
in sequence, an anhysteretic remanence (ARM, applied in 0.05 mT applied field and 100 mT af
field), and a saturation isothermal remanence (SIRM, applied at 1000 mT steady field). Both
remanences were af-demagnetized sequentially in fields of 10 mT, 20 mT, and 40 mT.
Variations in chi, ARM, and SIRM intensity were used to assess the concentration of magnetic
particles in the sediments (and by proxy, clastic sediment concentration). The ratios ARM/chi,
ARM20/0, and SIRM20/0 were used to estimate the percentage of finer-grained magnetic
particles (and by proxy, clastic grain size variability). All six parameters were plotted for cores
70

PESC-MC1 and PESC-GC1 and compared. Twelve clear parameter highs/lows could be
correlated between the cores and there was a simple linear relationship identified between them.
The correlations indicated that PESC-MC1 had 30 ± 2 cm of sediment not present in PESC-GC1.
Presumably, that sediment was lost from PESC-GC1 during coring.

3.3.4 Spectral Analysis
To identify periodicities in the Pescadero δ
15
N
sed
data, we performed spectral analysis
using a Lomb-Scargle Fourier Transform [Lomb, 1976; Scargle, 1982, 1989] in combination
with a Welch-Overlapped-Segment-Averaging (WOSA) procedure [Schulz and Stattegger, 1997;
Welch, 1967; Khider et al., 2014]. We chose this method since our record is not uniformly
spaced in time and other methods require the data to be evenly spaced in time. This is often
achieved through interpolation of the time series resulting in dependent data points and
potentially biased results [Schulz and Stattegger, 1997; Khider et al., 2014]. The Lomb-Scargle
periodogram weights the data on a “per point” basis instead of a “per time interval,” eliminating
the need for interpolation. Additionally, we performed wavelet analysis to investigate the
consistency of the periodicity through time. Here we apply the weighted wavelet Z-transform
(WWZ) for unevenly spaced time series [Foster, 1996; Schumann, 2005]. A mathematical
description of both methods is summarized in Khider et al. [2014].

3.4. Results
3.4.1 δ
15
N
sed
and Periodicity
Our δ
15
N
sed
record from the Pescadero Slope fluctuates between 8.1 and 10.4‰ over the
past 1121 years, with an overall mean value of 9.28 ± 0.34‰ (n=739, Figure 3.4). We find no
71

consistent long-term trend through the record. Variations between low δ
15
N
sed
values (interpreted
as a less intense OMZ) and high δ
15
N
sed
(a more intense OMZ) occur quickly, and are generally
followed by more gradual declines in OMZ intensity (from high δ
15
N
sed
to low δ
15
N
sed
) (Table
3.2). We assessed this quantitatively by dividing the δ
15
N
sed
record into ten distinct cycles and
calculating the rate of increase in δ
15
N
sed
to maximum values and the rate of decrease to minima.
Each cycle was defined as encompassing the full transition between an intensification/relaxation
of the OMZ that lasted longer than 40 years after a 20-year weighted average was applied to the
record. A 20-year weighted average was applied to define δ
15
N
sed
minima and maxima
objectively. In nine of ten intervals OMZ expansion or intensification occurred more quickly
than OMZ contraction. The one interval in which this pattern was not consistent (cycle 3) is an
atypical fluctuation in the record and represents the largest variation in the intensity of the OMZ.
In cycle 3, the rate of OMZ expansion is quite similar to the OMZ contraction rate. Considering
all 10 cycles, the average rate of δ
15
N
sed
increase was 0.025 ± 0.011 ‰ yr
-1
and the average rate
of δ
15
N
sed
decrease was 0.015 ± 0.012 ‰ yr
-1
. Thus, the increase in isotopic values occurs 1.7
times quicker than the rate of isotope decline. If gradients were calculated using minima and
maxima in the data (rather than the smoothed curve), the results similarly show rapid rates of
δ
15
N
sed
increase and slower rates of δ
15
N
sed
decrease (Table 3.2). In this case the rate of increase
is more than two times greater than the rate of decrease.
Spectral analysis of the Pescadero Slope δ
15
N
sed
reveals decadal, multi-decadal, and
centennial periodicities. Peaks at or above the 95% confidence level (Figure 3.5a) are found at
65, 86, 102 and 230 years. Wavelet analysis indicates these periodicities are strongest between
1250-1450 AD (Figure 3.5b).

72

3.4.2 δ
15
N
sed
, bSi, and C
org

Weight percent (wt. %) organic carbon in the sediments varies between 2.8 to 4.3% with
an average of 3.5% and uncertainty of ± 0.05% (n=31, Figure 3.6b). As with the δ
15
N
sed
values,
there is no overall temporal trend in wt. % C
org
. Biogenic silica (bSi) content of the sediments
varied between 2.6 and 21.8% over the last ~1200 years, however, as seen in the wt. % C
org
record there is no systematic nor long-term trend over this time period (Figure 3.6c). The single
high wt. % bSi value is ~3 times greater than the 7% mean bSi value as all other measurements
were between 3-11 ± 0.15%.
The relationships amongst all three geochemical proxies were further evaluated by
directly comparing each parameter to the other two. There is a positive, significant correlation (P
= 0.007) between wt. % C
org
and δ
15
N
sed
; high wt % C
org
is associated with elevated δ
15
N
sed
(Figure 3.7a). To assess the consistency of both productivity proxies, wt. % C
org
and wt. %bSi
were also directly compared and found to be highly significant (P < 0.001), with increased wt. %
C
org
corresponding to higher wt. % bSi (Figure 3.7b). Since correlations between wt. % C
org
and
both δ
15
N
sed
and wt. % bSi were significant, a correlation between δ
15
N
sed
and wt. % bSi was
expected, however, this relationship was not particularly significant (P = 0.35; Figure 3.7c).

3.5. Discussion
Climate influences denitrification and the extent of OMZs on glacial to interglacial time
scales [Altabet et al., 1995; Ganeshram et al., 1995, 2000], however, high-resolution fluctuations
and the drivers of the fluctuations have not been assessed on decadal to centennial time scales.
The data presented here suggests that the intensity of denitrification in OMZs could be self-
73

regulating over the past 1200 years and appears to be associated with climate processes, such as
solar irradiance and oscillations in ocean circulation.
The fluctuations in δ
15
N
sed
values from Pescadero Slope during the last millennium range
from a maximum of 10.4‰ to a minimum of 8.1‰, reflecting changes in the strength of the
OMZ in this region of the ETNP (Figure 3.4 and Figure 3.6a). The data shows that there are hard
ceilings and floors between which denitrification intensity varied during the late Holocene. The
consistency of these maximum and minimum values suggests internal feedbacks within the
system may be influencing the amount of denitrification and thus OMZ intensity. The
mechanism underpinning this regulation is not currently fully understood, however, we
hypothesize that organic carbon export could play a significant role in the extent of
denitrification in the ETNP OMZ. Organic carbon export and subsequent consumption of
oxygen through remineralization has been shown to be a dominant factor controlling
denitrification [Ward et al., 2008, 2009; Babbin et al., 2014]. Thus increased production and
export of organic carbon from the photic zone would result in increased denitrification, while
also reducing the availability of nitrate. This would consequently reduce primary production and
export of organic matter from the photic zone, since upwelling of nutrients from the OMZ drives
primary production in this region. The positive correlation between wt. %C and δ
15
N
sed
(Figure
3.7a) support this hypothesis, however, we recognize that other feedbacks and natural
fluctuations add complexity to self-regulation of the ETNP OMZ.
Lower resolution, longer-term records indicate that while mean δ
15
N
sed
values are lower
during glacial periods a similar range of

values (2-3‰) is observed. The ETNP region is
characterized by values ranging between 5.5 and 8.2‰ [Ganeshram et al., 1995] and the Arabia
Sea records values range from 5-7‰ during glacial periods. This suggests that denitrification
74

self-regulation mechanisms are present during both the recent Holocene interglacial and glacial
periods, with each period defined by different limits of denitrification intensity.
We interpret δ
15
N
sed
values as reflecting changes in the oxygen content of the OMZ. To
predict how oxygen within the ETNP changes relative to fluctuations in δ
15
N
sed
in the ETNP, we
compared water column δ
15
NO
3
-
and O
2
measurements obtained by Townsend-Small et al.
[2014] from the Pescadero, Soledad, and Magdalena Basins (Figure 3.8). We constrained the
analysis to these three sites closest to Pescadero and did not include data from basins above 25°N
due to the biological regime shift that occurs off the coast of Baja California at sites north of
25°N [Sañudo-Wilhelmy et al., 2012; Lluch-Belda et al., 2003]. We also avoided using data
where O
2
measurements <3 µM due to uncertainty in oxygen measurements at this low level
[Riser and Johnson, 2008; Bograd et al., 2008; Prokopenko et al., 2011; Pierce et al., 2012;
Townsend-Small et al., 2014]. The significant relationship between δ
15
NO
3
-
and O
2
(δ
15
NO
3
-
= -
0.05*[O
2
] + 15; r
2
= 0.86, P < 0.001), if directly applied to our δ
15
N
sed
values, indicates that a 1‰
change in δ
15
NO
3
-
is equivalent to a 20 ± 3.1 µM change in [O
2
]. This interpretation is a
simplification of the exact nature of the oxygen-nitrate relationship, and we do not exclude the
possibility that other factors may affect the isotope value of nitrate that exits the photic zone as
particulate N. These other factors include upwelling from different depths, partial utilization of
nitrate, ecosystem structure and extent and intensity of N fixation. To reduce uncertainty in this
relationship, the water depths analyzed ranged from 100 to 200 m. Water from this range has
measureable nitrate and is likely of a source of upwelled water to the surface ocean in this region
[Feely et al., 2008]. Additionally, minimal N
2
fixation has been found in close proximity to the
Pescadero site [White et al., 2013]. In keeping with a simple interpretation of the N cycle in this
coastal system, our data implies that over the past 1200 years there have been ~45 µM changes in
75

oxygen content in the upper portion of the ETNP OMZ. Some of these swings in oxygen content
occurred in as short as 40 years.
Recent observations from the ETNP indicate that O
2
is declining at a rate of 1.7 µM yr
-1

[Bograd et al., 2008]. This rate is similar to the average O
2
decline (1.2 µM yr
-1
) based on our
interpretation of changes in δ
15
N
sed
values during periods when the ETNP OMZ was expanding.
Average ETNP
15
N
sed
values during the last three glacial periods ranged from ~6.5 to 8.2‰
[Ganeshram et al., 1995]. Using our observed relationship between δ
15
NO
3
-
and O
2
we estimate
a 54 µM change in [O
2
] during each of these glacial periods. This is similar to the range of
fluctuation seen during the most recent interglacial. Hence, recent rates of oxygen concentration
change within global OMZ’s are comparable to rates of change that have occurred many times in
the past 1200 years.
Spectral analysis provides a method to determine if there are significant periodicities in
the data and this in turn can be used to assess what factors drive OMZ cyclicity during the late
Holocene. Marine and lacustrine sediments, tree-ring, fish scale deposition/populations, and
North American surface temperature studies have observed periodicities between 50 and 70
years [Baumgartner et al., 1992; Meko, 1992; Sharp, 1992; Schlesinger and Ramankutt, 1994;
Thurow and Schaaf, 1995; Minobe, 1997; Pike and Kemp, 1997; MacDonald and Case, 2005],
similar to the highly significant 65 year frequency identified in this study (Figure 4a). This
frequency of variability is thought to relate to Pacific basin wide changes in atmospheric and
oceanic circulation, known as the Pacific Decadal Oscillation (PDO) [Minobe, 1997]. When the
Aleutian Low intensifies relative the Pacific High, the southward flowing California Current
along the coast of Western North America weakens and moves offshore, resulting in reduced
upwelling, reduced biological productivity, and warmer sea surface temperatures [Pike and
76

Kemp, 1997; Mantua and Hare, 2002]. Tems et al. [2015] also found that the PDO drives
variations in the transport of elevated δ
15
NO
3
-
from the ETNP OMZ northward to the California
Borderland.
The 86, 102, and 230-year cycles could be related to solar activity, more specifically, the
Gleissberg and Suess (deVries) cycle. The Gleissberg cycle, traditionally thought to occur at an
80-90 year periodicity and a function of the 11-year Schwabe sunspot cycle, is more complex
having a wider frequency band with a double structure of 50-80 and 90-140 year periodicities,
than the Suess cycle, which shows variation with a period of 170-260 years [Ogurtsov et al.,
2002]. The Earth’s temperature and climate are associated with solar cycles [Friis-Christensen
and Lassen, 1991], suggesting that fluctuations in intensity of the ETNP OMZ are influenced by
changes in climate. Given the coincident periodicities, we hypothesize that ETNP OMZ intensity
is influenced by changes in insolation.
Qualitatively comparing our OMZ intensity record to reconstructed climate indices
including the El Niño Southern Oscillation (ENSO) [Emile-Geay et al., 2013], the Pacific
Decadal Oscillation [MacDonald and Case, 2005], and solar irradiance [Steinhilber et al., 2009]
show some periods of coherence between the records (Figure 3.9). Above average OMZ
intensity appears to correspond with increased solar irradiance and PDO cold phases between
1100 and 1720 AD. However, the relationship reverses from 1720-2000 AD, where above
average OMZ intensity corresponds to decreased solar irradiance and PDO warm phases. The
timing of this reversal corresponds to a divergence in the general trend of declining land carbon
(organic) storage between ~1300 and 1850 and significant increases in atmospheric CO
2
from
anthropogenic processes [Bauska et al., 2015]. Variation in ENSO phase occurs with a time
constant (2-7 years) that would not be easily resolved in our sediment record. General phases of
77

positive and negative ENSO do not appear to directly correlate with a change in OMZ intensity,
however, that does not eliminate the possibility that the OMZ intensity might be influenced by
ENSO strength.
Wavelet analysis is consistent with other studies that also observe 50 -70 year
periodicities that are most prominent between 1250-1450 AD [MacDonald and Case, 2005].
This finding supports the hypothesis that PDO periodicities have exhibited inconsistent strength
during the recent Holocene [MacDonald and Case, 2005].
The correlation between wt % C
org
and δ
15
N
sed
suggests that changes in local biologic
productivity may influence the intensity of the OMZ at the Pescadero Slope (Figure 3.7a).
Elevated δ
15
N
sed
is associated with higher wt % C
org
, suggesting either greater primary
productivity and export of carbon out of the photic zone results in a more intense OMZ or
reduced degradation of organic matter due to more extensive low [O
2
] during expanded OMZs.
A comparison of the two productivity proxies, wt % C
org
and wt % bSi, shows that the proxies
are positively correlated with high significance (Figure 3.7b). The transitive relationship,
however, does not hold since a significant correlation between δ
15
N
sed
and %bSi was not found
(Figure 3.7c). This lack of correlation could be due to: (1) the
15
NO
3
-
signal is entirely due to
transport from lower latitudes of the ETNP OMZ and hence is decoupled from local production
and export, (2) variable amounts of remineralization of bSi in the water column attenuate the
signal in sediments, (3) dissolution/diagenesis of bSi in the sediments occurs after deposition,
and/or (4) bSi accumulation represents diatom bloom events. Others have found that a significant
portion of bSi produced in the Gulf is dissolved within the upper water column (above 200 m)
[Thunell et al., 1994; Pride et al., 1999] which would alter the bSi signal in sediments.
Weight % bSi may better serve as a proxy for diatom/siliceous plankton bloom events rather than
78

overall net primary productivity. This suggests that wt. % C
org
in the sediments is the more
reliable proxy for assessing past changes in primary productivity on Pescadero Slope.
Both δ
15
N
sed
and climate modeling results indicate that warming climate conditions since
the onset of the industrial revolution until ~1990 AD resulted in a weakening of the easterly trade
winds, a deepening of the thermocline, and consequently a contraction of the ETNP OMZ
[Deustch et al., 2014]. The extended history of OMZ fluctuations based on the δ
15
N
sed
record
presented here corroborates this trend. To understand if this trend was consistent during other
anomalously warm and cold events we analyzed Pescadero OMZ intensity trends for the
Medieval Climate Optimum (MCO/950-1200 AD) and the Little Ice Age (LIA/1350-1850 AD)
(Figure 3.10). We find no statistically significant difference in OMZ intensity for the MCO
(mean value of 9.30 ± 0.28‰) and LIA (mean value of 9.33 ± 0.31‰) were not found. This may
suggest that (1) regional multidecadal oscillations were more dominant than centennial climate
events in the recent past or (2) the global climate anomalies were not as significant in the ETNP
as in other regions.
Previous studies assessing long-term productivity changes along the eastern slope of
Pescadero Basin have used diatom species abundance in addition to wt. % biogenic silica, albeit
at a lower resolution. Barron and Bukry [2007] found multiple distinctive peaks in A. nodulifera,
a tropical diatom, abundance and a decrease in the percentage of winter diatoms during the MCO.
This was suggestive of increased sea surface temperatures and reduced diatom productivity,
which the authors attributed to a greater influx of tropical Pacific waters into the Gulf of
California during the MCO. Our higher resolution analysis indicates that the multiple peaks
observed by Barron and Bukry [2007] could be indicative of fluctuations in productivity related
to inherent changes in the Gulf region, such as PDO or ENSO and may not only be due the MCO
79

itself. Our conclusion is that while we do see solar forcing in spectral analysis, the MCO and
LIA do not appear statistically different in this region of the ETNP illustrating the potential for
multidecadal fluctuations to drive regional differences at a time of global climate anomalies.

3.6. Conclusions
The intensity of the eastern tropical north Pacific (ETNP) oxygen minimum zone (OMZ)
during the last 1,200 years could be influenced by solar irradiance, the Pacific Decadal
Oscillation, and local productivity. The most significant finding of this study is that the ETNP
OMZ intensity fluctuates on shorter time scales than previously observed, with the OMZ
expanding quickly and contracting gradually, within the bounds of a regulated system that limits
the degree of δ
15
N
sed
variation to ±2.3‰. The average rate of OMZ intensification (1.2 µM
O
2
/year) is twice as fast as the rate of OMZ re-oxygenation. Intensity fluctuations are associated
with changes in productivity as evidenced by correlations between δ
15
N
sed
and weight percent
C
org
. In the Gulf of California, wt. % C
org
is a more accurate representation of productivity while
wt. % biogenic silica may more accurately serve as a record of bloom events. Productivity
variations and subsequent OMZ intensity fluctuations at centennial frequencies are likely
influenced by changes in wind strength, which is controlled by Earth’s climate. Increased OMZ
intensity also shows coherence with increased solar radiation and Pacific Decadal Oscillation
(PDO) cold phases between ~1100 to 1720 AD, and a reversal of this relationship between
~1720 and 2000 AD.

80

Acknowledgments
This work has been submitted to Paleoceanography (November 2015) and coauthors of
this work include William M. Berelson, Robert Thunell, Eric Tappa, Xiaomei Xu, Deborah
Khider, Steve Lund, and

Oscar González-Yajimovich.
This work was supported by a National Science Foundation grant (OCE-0727123 to
W.B.). We would like to acknowledge how this work was supported and critiqued by Donn
Gorsline who served on my Ph.D. Qualifying Committee, but was not able to see this work come
to fruition. Additionally, we would like to thank Alexander van Geen, Nick Rollins, Jennifer
Lehman, John Fleming, Maria Prokopenko, James McManus, Curtis Deutsch, UC Davis Stable
Isotope Facility, and Captain and Crew of the RV Horizon.

81

References
Altabet, M. A., and R. Francois (1994), Sedimentary nitrogen isotopic ratio as a recorder for
surface ocean nitrate utilization, Global Biogeochem. Cycles, 8(1), 103–116,
doi:10.1029/93GB03396.
Altabet, M.A., R. Francois, D.W. Murray, and W.L. Prell (1995), Climate-related variations in
denitrification in the Arabian Sea from sediment
15
N/
14
N ratios, Nature, 373(6514), 506-509,
doi:10.1038/373506a0.
Altabet, M. A., C. Pilskaln, R. Thunell, C. Pride, D. Sigman, F. Chavez, and R. Francois (1999),
The nitrogen isotope biogeochemistry of sinking particles from the margin of the eastern
North Pacific, Deep Sea Res. Part I, 46, 655–679, doi:10.1016/S0967-0637(98)00084-3.
Babbin, A.R., R.G. Keil, A.H. Devol, and B.B. Ward (2014), Organic matter stoichiometry, flux,
and oxygen control nitrogen loss in the ocean, Science, 344, 406-408,
doi:10.1126/science.1248364.
Barron, J.A. and D. Bukry (2007), Solar forcing of Gulf of California climate during the past
2000 yr suggested by diatoms and silicoflagellates., Marine Micropaleontology 62: 115–139.
Baumgartner, T. R., A. Soutar, and V. Ferreira-Bartrina (1992), Re-construction of the history of
Pacific sardine and northern anchovy populations over the past two millennia from sediments
of the Santa Barbara basin, California. CalCOFI Reports 33, 24–40.
Bauska, T.K., F. Joos, A.C. Mix, R. Roth, J. Ahn, and E.J. Brook (2015), Links between
atmospheric carbon dioxide, the land carbon reservoir and climate over the past millennium,
Nature Geoscience, 8, 383-387, doi:10.1038/NGEO2422.
82

Benton, M.J. and R.J. Twitchett (2003), How to kill (almost all life: the end-Permian extinction
event, Trends in Ecology and Evolution, 18(7), 358-365, doi:10.1016/S0169-5347(03)00093-
4.
Bloemendal, J., B. Lamb, and J. King (1988), Paleoenvironmental implications of rock-magnetic
properties of late Quaternary sediment cores from the eastern equatorial Atlantic.
Paleoceanography, 3, 61-87.
Bograd, S.J., C.G. Castro, E. Di Lorenzo, D.M. Palacios, H. Hailey, W. Gilly, and F.P. Chavez
(2008), Oxygen declines and the shoaling of the hypoxic boundary in the California Current,
Geophysical Research Letters, 35, L12607, doi:10.1029/2008GL034185.
Casciotti, K. L., C. Buchwald, and M. McIlvin (2013), Implications of nitrate and nitrite isotopic
measurements for the mechanisms of nitrogen cycling in the Peru oxygen deficient zone,
Deep-Sea Research I, 80: 78-98.
Chavez, F.P., M. Messié, and J.T. Pennington (2011), Marine Primary Production in Relation to
Climate Variability and Change, Annu. Rev. Mar. Sci, 3, 227-260, doi:
10.1146/annurev.marine.010908.163917.
Chong, L. S., M. G. Prokopenko, W. M. Berelson, A. Townsend-Small, and J. McManus (2012),
Nitrogen cycling within suboxic and anoxic sediments from the continental margin of
western North America, Mar. Chem., 128–129, 13–25, doi:10.1016/j.marchem.2011.10.007.
DeMaster, D.J. (1979), The Marine Budgets of Silica and
32
Si, Yale University, New Haven.
DeMaster, D.J. (1991), Measuring biogenic silica in sediments and suspended matter, American
Geophysical Union.
Deutsch, C. W. Berelson, R. Thunell, T. Webber, C. Tems, J. McManus, J. Crusius, T. Ito, T.
Baumgartner, V. Ferreira, J. Mey, and A. VanGeen (2014), Centennial changes in North
83

Pacific anoxia linked to tropical trade winds, Science, 345 (6197), 665-668,
doi:10.1126/science.1252332.
Douglas, R. O. Gonzalez-Yajimovich, J. Ledesma-Vazquez, and F. Staines-Urias (2007),
Climate forcing, primary production and the distribution of Holocene biogenic sediments in
the Gulf of California, Quaternary Science Reviews, 26, 115-129, doi:
10.1016/j.quascirev.2006.05.003.
Emile-Geay, J., K.M. Cobb, M.E. Mann, and A.T. Wittenberg (2013), Estimating central
equatorial Paciffic SST variability over the past millennium. Part II: reconstructions and
implications, Journal of Climate, 26(7), 2329-2352, doi:10.1175/JCLI-D-11-00511.1.
Feely, R.A., C.L. Sabine, J.M. Hernandez-Ayon, D. Ianson, and B. Hales (2008), Evidence for
upwelling of corrosive "acidified" water onto the continental shelf, Science, 320, 1490-1492.
Fernández-Barajas ME, M.A. Monreal-Gómez, A. Molina-Cruz (1994), Thermohaline structure
and geostrophic flow in the Gulf of California, during 1992, Ciencias Marinas
20, 267–286.
Foster, G. (1996), Wavelets for period analysis of unevenly sampled time series, Astron. J.,
112(4), 1709–1729.
Friis-Christensen, E. and K. Lassen (1991), An indicator of solar activity closely associated with
climate, Science, 254(5032), 698-700.
Ganeshram, R.S., T.F. Pedersen, S.E. Calvert, and J.W. Murray (1995), Large changes in oceanic
nutrient inventories from glacial to interglacial periods, Nature, 376, 755-758.
Ganeshram, R. S., T. F. Pedersen, S. E. Calvert, G. W. McNeil, and M. R. Fontugne (2000),
Glacial-interglacial variability in denitrification in the world’s oceans: Causes and
consequences, Paleoceanography, 15(4), 361–376, doi:10.1029/1999PA000422.
84

Gilly, W.F., J.M. Beman, S.Y. Litvin, and B.H. Robinson (2013), Oceanographic and biological
effects of shoaling of the oxygen minimum zone, Annu. Rev. Mar. Sci, 5, 21.1-21.28,
doi:10.1146/annurev-marine-120710-100849.
Jones, C.E. and H.C. Jenkyns (2001), Seawater strontium isotopes, ocean anoxic events, and
seafloor hydrothermal activity in the Jurassic and Cretaceous, American Journal of Science,
301, 112-149.
Keeling, R.F., A. Körtzinger, and N. Gruber (2010), ocean deoxygenation in a warming world,
Annu. Rev. Mar. Sci. 2, 199-229, doi:101146/annurev.marine.010908.163855.
Khider, D., C. S. Jackson, and L. D. Stott (2014), Assessing millennial-scale variability during
the Holocene: A perspective from the western tropical Pacific, Paleoceanography, 29, 143–
159, doi:10.1002/2013PA002534.
King, J.W., and J.E.T. Channell (1991), Sedimentary magnetism, environmental magnetism, and
magnetostratigraphy. Reviews of Geophysics, supplement, US National Report to IUGG,
358-370.
Lehmann, M. F., S. M. Bernasconi, A. Barbiere, and J. A. McKenzie (2002), Preservation of
organic matter and alteration of its carbon and nitrogen isotope composition during simulated
and in situ early sedimentary diagenesis, Geochim. Cosmochim. Acta, 66(20), 3573–3582,
doi:10.1016/S0016-7037(02)00968-7.
Lluch-Belda, D., D.B. Lluch-Cota, and S.E. Lluch-Cota (2003), Baja California’s biological
transition zones: refuges for the California sardine, Journal of Oceanography, 59, 503-513.
Lomb, N. R. (1976), Least-squares frequency analysis of unequally spaced data, Astrophys.
Space Sci., 39, 447–462.
85

MacDonald, G.M. and R.A. Case (2005), Variations in the Pacific Decadal Oscillation over the
past millennium, Geophysical Research Letters, 32: L08703, doi:10.1029/2005GL022478.
Mantua, N. J., and S. R. Hare (2002), The Pacific Decadal Oscillation, J. Oceanogr., 58, 35–44,
doi:10.1023/A:1015820616384.
McClatchie, S., R. Goericke, R. Cosgrove, and G. Auad, R. Vetter (2010), Oxygen in the
Southern California Bight: Multidecadal trends and implications for demersal fisheries,
Geophysical Research Letters, 37, L19602, doi: 10.1029/2010GL044497.
Meko, D.M. (1992), Spectral properties of tree-ring data in the United States Southwest as
related to El Niño/Southern Oscillation, in El Niño: Historical and Paleoclimatic Aspects of
the Southern Oscillation, edited by H.F Diaz and V. Markgraf, 227-242.
Minobe, S. (1997), A 50-70 year climatic oscillation over the North Pacific and North America,
Geophysical Research Letters, 24(6), 683-686.
Ogurtsov, M.G., Y.A. Nagovitsyn, G.E. Kocharov, and H. Junger (2002), Long-period cycles of
the sun’s activity recorded in direct solar data and proxies, Solar Physics, 211, 371-394.
Olsson, I. (1986), Radiometric Methods. In: Berglund, B. (eds) Handbook of Holocene
palaeoecolgy and palaeohydrology, John Wiley & Sons, Chichester, pp 273-312.
Pichevin, L., R.S. Ganeshram, B.C. Reynolds, F. Prahl, T.F. Pederson, R. Thunell, and E.L.
McClymont (2012), Silicic acid biogeochemistry in the Gulf of California: Insignts from
sedimentary Si isotopes, Paleooceanography, 27, PA2201, doi:10.1029/2011PA002237.
Pierce, S.D., J.A. Barth, R.K. Shearman, and A.Y. Erofeev (2012), Declining oxygen in the
northeast Pacific, Journal of Physical Oceanography, 42(3), 495-501, 10.1175/JPO-D-11-
0170.1
86

Pike, J. and A.E.S. Kemp (1997), Early Holocene decadal-scale ocean variability recorded in
Gulf of California laminated sediments, Paleoceanography, 12(2), 227-238.
Pride, C., R. Thunell, D. Sigman, L. Keigwin, M. Altabet, E. Tappa (1999), Nitrogen isotopic
variations in the Gulf of California since the Last Deglaciation: Response to global climate
change, Paleococeanography, 14(3): 397-409, doi: 1.1029/1999PA900004.
Prokopenko, M., D. E. Hammond, A. J. Spivack, and L. Stott (2006), Impact of long-term
diagenesis on δ
15
N of organic matter in marine sediments: Sites 1227 and 1230, Proceedings
from the Ocean Drilling Program, Scientific Results, 201, 30.
Prokopenko, M., D.M. Sigman, W.M. Berelson, D.E. Hammond, B. Barnett, L. Chong, and A.
Townsend-Small (2011), Denitrification in anoxic sediments supported by biological nitrate
transport, Geochimica et Cosmochimica Acta, 75, 7180-7199, doi:10.1016/j.gca.2011.09.023.
Ragueneau, O., P. Tréguer, A. Leynaert, R.F. Anderson, M.A. Brzezinski, D.J. DeMaster, R.C.
Dugdale, J. Dymond, G. Fischer, R. Francois, C. Heinze, E. Maier-Reimer, V. Martin-
Jézéquel, D.M. Nelson, and B. Quéguiner (2000), A review of the Si cycle in the modern
ocean: recent progress and missing gaps in the application of biogenic opal as a
paleoproductivity proxy, Global and Planetary Change, 26: 317-365, doi:10.1016/S0921-
8181(00)00052-7.
Reimer P.J., E. Bard, A. Bayliss, J.W. Beck, P.G. Blackwell, C. Bronk Ramsey, C.E. Buck, H.
Cheng, R.L. Edwards, M. Friedrich, P.M. Grootes, T.P. Guilderson, H. Haflidason, I. Hajdas,
C. Hatté, T.J. Heaton, A.G. Hogg, K.A. Hughen, K.F. Kaiser, B. Kromer, S.W. Manning, N.
Niu, R.W. Reimer, D.A. Richards, E.M. Scott, J.R. Southon, C.S.M. Turney, J. van der Plicht
(2013), IntCal13 and MARINE13 radiocarbon age calibration curves 0-50000 years calBP,
Radiocarbon 55(4), doi:10.2458/azu_js_rc.55.16947.
87

Riser, S.C. and K.S. Johnson (2008), Net production of oxygen in the subtropical ocean, Nature,
451(7176), 323-325, doi:10.1038/nature06441.
Sañudo-Wilhelmy, S.A., L.S. Cutter, R. Durazo, E.A. Smail, L. Gómez-Consarnau, E.A. Webb,
M.G. Prokopenko, W.M. Berelson, and D.M. Karl (2012), Multiple B-vitamin depletion in
large areas of the coastal ocean, Proceedings of the National Academy of Sciences, 109(35),
14041-14045.
Scargle, J. D. (1982), Studies in astronomical time series analysis. II. Statistical aspects of
spectral analysis of unevenly spaced data, Astrophys J., 263(2), 835–853.
Scargle, J. D. (1989), Studies in astronomical time series analysis. III. Fourier transforms,
autocorrelation functions, and cross-correlation functions of unevenly spaced data, Astrophys.
J., 343(2), 874–887.
Schlesinger M, .E., and N. Ramankutty (1994), An oscillation in the global climate system of
period 65-70 years, Nature, 3(67), 723-726.
Schulz, M., and K. Stattegger (1997), SPECTRUM: Spectral analysis of unevenly spaced time
series, Comput. Geosci., 23(9), 929–945.
Sharp, G.D. (1992), Fishery catch records, El Niño/Southern Oscillation, and longer-term
climate change as inferred from fish remains in marine sediments, in El Niño: Historical and
Paleoclimatic Aspects of the Southern Oscillation, edited by H.F. Diaz and V. Markgraf,
379-418.
Sigman, D. M., J. Granger, P. J. Difiore, M. M. Lehmann, R. Ho, G. Cane, and A. Van Geen
(2005), Coupled nitrogen and oxygen isotope measurements of nitrate along the eastern
North Pacific margin, Global Biogeochem. Cycles, 19, GB4022, doi:10.1029/
2005GB002458.
88

Southon, J.R. and G.M. Santos (2004), Ion source development at KCC AMS, University of
California, Irvine. Radiocarbon 46, 33–39.
Southon, J.R. and G.M. Santos (2007), Life with MC-SNICS. Part II: Further ion source
development at the Keck carbon cycle AMS facility. Nuclear Instruments and
Methods in Physics Research, Section B 259, 88–93.
Staines-Urías, F., R.G. Douglas, D.S. Gorsline (2009), Oceanographic variability in the southern
Gulf of California over the past 400years: Evidence from faunal and isotopic records from
planktic foraminifera, Palaeogeography, Palaeoclimatology, Palaeoecology, 284 (3-4): 337-
354, doi: 10.1016/j.palaeo.2009.10.016.
Steinhilber Steinhilber, F., J. Beer, and C. Fröhlich (2009), Total solar irradiance during the
Holocene, Geophys. Res. Lett., 36, L19704, doi:
10.1029/2009GL040142.
Stramma, L., GC. Johnson, J. Sprintall, V. Mohrholz (2008), Expanding oxygen-minimum
zones in the tropical oceans, Science, 320: 655-658.
Strickland, J.P.T. (1968), A Practical Handbook of Seawater Analysis, Fisheries Research Board
of Canada, Halifax.
Stuiver, M., and A.H. Polach (1977), Discussion: Reporting of 14C data, Radiocarbon, 19:355-
363.
Stuiver, M. and P.J. Reimer (1993), Extended 14C database and revised CALIB radiocarbon
calibration program, Radiocarbon, 35:215-230.
Tems, C. E., W.M. Berelson, and M.G. Prokopenko (2015), Particulate δ
15
N in laminated marine
sediments as a proxy for mixing between the California Undercurrent and the California
Current: A proof of concept, Geophysical Research Letters, 42, doi:10.1002/ 2014GL061993.
89

Thunell, R.C., C.J. Pride, E. Tappa, and F.E. Muller-Karger (1994), Biogenic silica fluxes and
accumulation rates in the Gulf of California, Geology, 22, 303-306, doi: 10.1130/0091-
7613(1994)022<0303:BSFAAR>2.3.CO;2.
Thunell, R. C., and A. B. Kepple (2004), Glacial-Holocene δ
15
N record from the Gulf of
Tehuantepec, Mexico: Implications for denitrification in the eastern equatorial Pacific and
changes in atmospheric N2O, Global Biogeochem. Cycles, 18, GB1001,
doi:10.1029/2002GB002028.
Thunell, R., D.M. Sigman, F.E. Muller-Karger, Y. Astor, R. Varela (2004), Nitrogen Isotopic
Dynamics of the Cariaco Basin, Venezuela, Global Biogeochemical Cycles, 18: GB3001,
doi:10.1029/2003GB002185.
Thurow, J., and M. Schaaf (1995), A comparison of Holocene to late Pleistocene grayvalue-time-
series of sediments from the Santa Barbara Basin (ODP Site 893) and the Guaymas Basin
(DSDP Site 480), in IGCP 374: Palaeoclimatology and Palaeoceanography From Laminated
Sediments, edited by R.J. Behl and J.P. Kennett, 33.
Townsend-Small, A., M. G. Prokopenko, and W. M. Berelson (2014), Nitrous oxide cycling in
the water column and sediments of the oxygen minimum zone, eastern subtropical North
Pacific, Southern California, and Northern Mexico (23°N–34°N), J. Geophys. Res. Oceans,
119, 3158–3170, doi:10.1002/2013JC009580.
Ward, B.B., C.B. Tuit, Am Jayakumar, J.J. Rich, J. Moffett, S. Wajih, and A. Naqvi (2008),
Organic carbon, not copper, controls denitrification in oxygen minimum zones of the ocean,
Deep-Sea Research I, 55: 1672-1683.
90

Ward, B.B, A.H. Devol, J.J. Rich, B.X. Chang, S.E. Bulow, H. Naik, A. Pratihary, and A.
Jayakumar (2009), Denitrification as the dominant nitrogen loss process in the Arabian Sea,
Nature, 461, 78-82, doi:10.1038/nature08276.
Welch, P. D. (1967), The use of Fast Fourier transform for the estimation of power spectra: A
method based on time averaging over short, modified periodograms, IEEE Trans. Audio
Electroacoust., 15(2), 70–73.
White, A.E., R.A. Foster, C.R. Benitez-Nelson, P. Masqué, E. Verdeny, B.N. Popp, K.E. Arthur,
and F.G. Prahl (2013), Nitrogen fixation in the Gulf of California and the Eastern Tropical
North Pacific, Progress in Oceanography, 109, 1-17, doi:10.1016/j.pocean2012.09.002.
Wingall, P.B. and R.J. Twitchett (1996), Ocean anoxia and the end Permian mass extinction,
Science, 272(5265), 1155-1158, doi: 10.1126/science.272.5265.1155.
Xu, X., Trumbore, S.E., Zheng, S.H., Southon, J.R., McDuffee, K.E., Luttgen, M., Liu, J.C.
(2007), Modifying a sealed tube zinc reduction method for preparation of AMS graphite
targets: reducing background and attaining high precision. Nuclear Instruments and Methods
in Physics Research, Section B 259, 320–329.

91

Table 3.1. Original
14
C

Ages (BP) measured at UC Irvine, core depth, the corrected ages (AD)
after applying a local reservoir effect (ΔR = 508 ± 30) and σ
1
error of the corrected ages for the
PESC-MC1 and PESC-GC3.

Average Depth (cm) Uncorrected
14
C

Age (BP) ΔR Corrected Year (AD) σ1 error (± years)
5.25 540 ± 20 not applicable not applicable
18.45 985 ± 20 1860 43.5
41.85 1015 ± 20 1818 40
51.75 1015 ± 20 1818 40
41.85 1090 ± 20 1733 43
46.5 1140 ± 20 1670 37.5
48.45 1180 ± 20 1628 44.5
58.95 1240 ± 20 1576 48
72.45 1235 ± 20 1580 48
79.9 1440 ± 20 1416 32
85.75 1320 ± 20 1497 31.5
96.55 1410 ± 20 1437 25.5
102.85 1440 ± 20 1416 32
104.5 1420 ± 20 1431 27
112.15 1590 ± 20 1303 30
121.35 1530 ± 20 1352 36.5
133.05 1540 ± 20 1345 25.5
142.05 1700 ± 20 1219 40.5
146.8 1725 ± 20 1194 46.5
157.65 1800 ± 20 1113 47.5
167.5 1795 ± 20 1117 48
178.65 1910 ± 20 1001 39.5
188.55 1985 ± 20 927 50.5
193.05 1950 ± 20 962 45
197.65 2035 ± 20 865 54
202.45 2075 ± 20 821 51.5
PESC-MC1
PESC-GC3
14
C Age Model Data
92

Table 3.2. A quantitative assessment of OMZ fluctuations based changes in δ
15
N
sed
. Each data
section is indicated (refer to Figure 3 for exact location) and the slope (absolute value) calculated
between δ
15
N
sed
local minima and maxima (Slope L-H) and δ
15
N
sed
local maxima and minima
(Slope H-L). This is calculated for each section after a 20-year smoothed curve was applied and
for the original data presented in Figure 3. These calculations indicate that there is an average
intensification of 0.025 ± 0.11‰ per year and an average de-intensification of 0.015 ± 0.12‰
per year when the 20 year weighted average is applied and an average of intensification of 0.060
± 0.033‰ per year and de-intensification of 0.030 ± 0.28‰ per year based on the original data
set.

Quantitative Assessment of OMZ Fluctuations
Section Slope (L-H)* Slope (H-L)*
20-year Weighted Average
1 0.0169 0.0086
2 0.0237 0.0037
3 0.0415 0.0458
4 0.0353 0.0203
5 0.0189 0.0121
6 0.0432 0.0169
7 0.0186 0.0091
8 0.0141 0.0091
9 0.0171 0.0157
10 0.0187 0.0069
Average 0.0248 0.0148
Original Data
1 0.0297 0.0170
2 0.0875 0.0080
3 0.1068 0.0545
4 0.1032 0.0996
5 0.0447 0.0207
6 0.0911 0.0196
7 0.0236 0.0133
8 0.0472 0.0302
9 0.0340 0.0225
10 0.0318 0.0105
Average 0.0600 0.0296
* absolute value of the slope between L (low δ
15
N
sed
values) and H (high δ15Nsed values) and vise versa
93

Figure 3.1. A map of the Gulf of California. Pescadero Basin, near the mouth of the Gulf of
California and the Pacific Ocean, is shaded gray. The coring location along the Pescadero Slope
is represented by the black dot.

20˚N
25˚N
30˚N
120˚W 115˚W 110˚W 105˚W
Pescadero Basin
Core Location
Pacific Ocean
Baja California
Mexico
Gulf of California
United States
94

0 10 20 30 40 50 60
0 10 20 30 40 50 60
Depth (cm)
Depth (cm)
ARM
SIRM
ARM
SIRM
PESC-MC1
PESC-GC1
1 2 3 4 5
1 2 3 4 5
0
.5
1.0
1.5
2.0
2.5
ARM (x10-4 SI)
0
.5
1.0
1.5
2.0
2.5
ARM (x10-4 SI)
0
2
4
6
8
10
SIRM (x10-2 SI)
0
2
4
6
8
10
SIRM (x10-2 SI)
a.
b.
95

Figure 3.2. The data utilized to correlate PESC-MC1 with PESC-GC1 and PESC-GC1 with
PESC GC3. PESC-MC1 and PESC-GC1 were correlated based on bulk magnetic susceptibility,
δ
13
C
(total)
, and weight percent total carbon. Bulk magnetic susceptibility (chi) was measured and
artifical remanence (ARM) and a saturation isothermal remanence (SIRM) were applied and
used to assess the concentration of magnetic particles in the sediments. Parameter highs/lows
were correlated between the cores, identifying a simple linear relationship and that PESC-MC1
(a) had 30±2 cm of sediment not present in PESC-GC1 (b). To constrain the correlation between
PESC-MC1 and PESC-GC1 the weight percent of total carbon (c) and δ
13
C
(total)
(d) was
measured on both cores. These measurements indicate that PESC-GC1 is missing 27.9 cm of
sediment when compared to PESC-MC1. Due to a missing section in PESC-GC3, PESC-GC1
was correlated to the core using δ
13
C
(total)
(e) which indicated the PESC-GC3 was missing 0.9 cm
of sediment relative to PESC-GC1 and therefore the offset of between PESC-MC1 and PESC-
GC3 is 28.8 cm.
26 27 28 29 30 31 32 33
3
3.5
4
4.5
26 27 28 29 30 31 32 33
-19
-18.5
-18
-17.5
72 74 76 78 80 82
-22
-20
-18
wt. % Total Carbon
δ
13
C
(total)
(‰) δ
13
C
(total)
(‰)
Depth (cm - relative to PESC-MC1)
Depth (cm - relative to PESC-MC1)
Depth (cm - relative to PESC-MC1)
PESC-MC1 and PESC-GC1
PESC-MC1 and PESC-GC1
PESC-GC1 and PESC-GC3
PESC-MC1
PESC-GC1 (+27.9 cm in relative to PESC-MC1)
PESC-GC3 (+28.8 cm in relative to PESC-MC1)
c.
d.
e.
96

Figure 3.3. The age model constructed for PESC-GC3. a) Radiocarbon measurements were
calibrated using Calib 7.1 with a local reservoir age (corrected for the global reservoir age; ΔR)
of 508 ± 30 years. This age is compared to the integrated mass, which relates to depth in the
core by taking porosity of the sediments into consideration therefore correcting for compaction.
The age of the sediments (y) relates to the integrated mass (x) by the following equation, y=-
12.89*x + 1999 and is statistically significant (r
2
= 0.97, P <0.001). The sedimentation rate is
77.4 mg cm
-2
yr
-1
. b) Varve counts were completed by dividing the core into 1 cm bins. The
number of varves per 1 cm bin is compared to age as calculated from the radiocarbon age model.
If sedimentation rate is linear, as assumed by the age model, 5.2 varves should be present in each
centimeter. While the sedimentation rate does not consistently remain constant, it minimally
fluctuates around 5.2 varves per centimeter.

1200 1400 1600 1800
0
2
4
6
8
10
12

a.
b.
y = -12.89*X + 1999
r
2
= 0.98
P < 0.001
Year (AD)
Integrated Mass (g ● cm
-2
)
Varves (# per cm)
Year (AD)
1000
1500
2000
0 20 40 60 80
500
97

Figure

3.4. δ
15
N
sed
(‰) from the Pescadero Slope multicore and gravity core compared to age as
calculated by
210
Pb and
14
C age models. Individual data points are represented by black dots,
while a 20-year smoothed curve is shown in gray. The record has been divided into 10 sections,
which each include intensification (increase in δ
15
N
sed
) and de-intensification (decrease in
δ
15
N
sed
) that occurred over a minimum of 40 years. This illustrates the ETNP OMZ expands or
intensifies quickly, while it contracts or de-intensifies more gradually. Refer to Table 1 for
calculations of changes for both smoothed and original data. Original data points used in the
calculations and indicated on the figure as white dots (outlined in black). The section between 9
and 10 was not included in this analysis since there is missing data in this section.

1000 1200 1400 1600 1800 2000
8
9
10
1 2 3 4 5 6 7 8 9 10
δ
15
N
sed
(‰)
Year (AD)
98

Figure 3.5. a) Spectral analysis using a Lomb-Scargle Fourier Transform in combination with a
Welch-Overlapped-Segment-Averaging (WOSA) procedure. Peaks representing 65, 86, 102, and
230 years are observed in the record with greater than 95% confidence. b) Results for wavelet
analysis by applying the weighted wavelet Z-transform (WWZ) for unevenly spaced time series
indicate that the strongest periodicities occur between 1200 and 1450 AD in this record. The
white dashed line represents the cone of influence. We only interpreted results below this line
since they are not affected by age effects. Black lines indicate a 95% confidence ensemble of
periodicity in the δ
15
N
sed
.
10 100
0
10
20
30
40
50
60
Power
a.
95%
50%
86
102
230
65
Years (AD)
Period

1000 1200 1400 1600 1800 2000
100
200
300
400
500
Power
Period
(years) b.
0
0.5
1
1.5
3
3.5
2
2.5
99

Figure 3.6. Geochemical measurements from sediment splits from cores PESC-MC1 and PESC-
GC3 compared to the year (AD) the sediments were deposited. a) δ
15
N
sed
(‰) fluctuates between
8.1 and 10.4‰ with inherent cyclicity, however, with no systematic trend and an average value
of 9.28‰. Uncertainty was calculated to be 0.17‰ for each measurement and is represented by
an error bar in the top right corner. b) Weight percent organic carbon varies between 2.8 and
4.3% with an average of 3.8%. Uncertainty is ±0.05% and is encompassed within the size of the
data point. c) Weight percent biogenic silica fluctuates between 3 and 21%, however, most
measurements lie between 3 and 11%. Uncertainty for each measurement was ±0.15%, which is
no larger than the data point itself.

1000 1200 1400 1600 1800 2000
8
9
10
1000 1200 1400 1600 1800 2000
2.5
3.5
4.5
1000 1200 1400 1600 1800 2000
0
10
20
δ
15
N
sed
(‰)
Little Ice Age Medieval Climate Optimum
wt. % C
org
wt. % bSi
a.
b.
c.
Year AD
Year AD
Year AD
100

Figure 3.7. A direct comparison between the three geochemical parameters measured. Error bars
in the top right of each plot represent uncertainty for each measurement. a) Weight percent
organic carbon significantly correlates with δ
15
N
sed
(‰) measured on the same sample (P=0.007).
b) The two productivity proxies, weight percent carbon and weight percent biogenic silica, have
a highly significant correlation (P<0.001). The measurement in parentheses was not included in
the correlation as the measurement was more than three times greater then the average wt. % bSi
measurement. c) The correlation between weight percent biogenic silica and δ
15
N
sed
(‰) is not
particularly strong (P=0.35).

0
10
20
δ
15
N
sed
(‰)
wt. % C
org
2.5
3.5
4.5
8 9 10
wt. % C
org
2.5
3.5
4.5
0 10 20
wt. % bSi
δ
15
N
sed
(‰)
8 9 10
wt. % bSi
( )
( )
y = 0.19 * x +1.8
r
2
= 0.10
P = 0.007
y = 0.077 * x + 3.0
r
2
= 0.33
P = <0.001
y = 0.83 * x - 0.76
r
2
= 0.025
P = 0.35
Present - 124 years BP 123 - 209 years BP 209 - 688 years BP 688 - 1042 years BP
a.
b.
c.
101

Figure 3.8. A comparison between the dissolved δ
15
NO
3
-

(‰) and O
2
content (µM) in the water
column at Magdalena Basin, Pescadero Basin, and Soledad Basin from data collected by
Townsend-Small et al. [2014]. δ
15
NO
3
-

relates to [O
2
] in the following way: δ
15
NO
3
-
= -0.05*[O
2
]
+15 (r
2
= 0.86, P >0.001). The best-fit linear regression of the data is represent by the red line.
Calculation of a least squares fit provided the possible error in the slope and intercept values.
The envelope of uncertainty is represented by the space between the gray (maximum deviation in
slope and intercept values) and green (minimum deviation in slope and intercept values). This
model calculates that a 20 ± 3.1µM change in oxygen content is equivalent to a 1‰ in δ
15
N
sed
.

y = -0.05 * X + 15
r
2
= 0.86
P < 0.001
10
11
12
13
14
15
16
10 0 20 30 40 50 60 70 80 90
9
100
O
2
(µM)
δ
15
NO
3
-

(‰)
Original Data
Best Fit Linear Regression
Minimum Slope/Intercept Value
Maximum Slope/Intercept Value
102

Figure 3.9. Qualitative comparison of (a) δ
15
N
sed
(b) reconstructed El Niño Southern Oscillation
Index (Had SST reconstruction) [Emile-Geay et al., 2013] (c) reconstructed Pacific Decadal
Oscillation Index [MacDonald and Case, 2005] (d) reconstructed solar irradiance (Steinhilber et
al., 2009). The same 20-year smoothing was applied to the original δ
15
N
sed
, ENSO, and PDO
data.

1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
8.5
9
9.5
10
1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
-1
-0.5
0
0.5
1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
-1
0
1
1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
-1
-0.5
0
0.5
δ
15
N
sed
ENSO Index Solar Irradiance PDO Index
Year (AD)
Year (AD)
Year (AD)
Year (AD)
a.
b.
c.
d.
103

Figure 3.10. Comparison between Medieval Climate Optimum (a) and Little Ice age (b) trends
in the intensity of the OMZ in the Gulf of California. The trend of OMZ de-intensification in the
Medieval Climate Optimum and intensification during the Little Ice Age are not statistically
significant. The slopes of the trends are similar (-0.00078 and 0.00032, respectively) and close
to a neutral, zero value.
950 1000 1050 1100 1150 1200
8
8.5
9
9.5
10
10.5
Medieval Climate Optimum

y = - 0.00078*x + 10
r
2
= 0.043
1350 1400 1450 1500 1550 1600 1650 1700 1750 1800 1850
8.5
9
9.5
10
10.5
Little Ice Age

y = 0.00032*x + 8.8
r
2
= 0.022
∆δ
15
N
sed
(
o
/
oo
) ∆δ
15
N
sed
(
o
/
oo
)
Year AD
Year AD
a.
b.
104

Chapter 4
Assessing Regional Variability in the Eastern Tropical North Pacific Oxygen
Minimum Zone through a Geochemical Comparison of Pescadero Slope and
Soledad Basin Sediments

Abstract
The marine nitrogen cycle is dynamic and experiences episodic variations that are not
well constrained. This study investigates changes in the intensity of denitrification and the
oxygen minimum zone (OMZ) in the eastern tropical north Pacific through the measurement of
δ
15
N from bulk laminated sediments between the Pescadero Slope, Gulf of California, and the
Soledad Basin, in the Pacific Ocean 45 km from Baja California Sur, to assess variation in
nitrogen cycling in these environments. Similar trends in OMZ expansion are observed in both
records, however, the expansions are offset in time with events first occurring in Soledad Basin
and then at Pescadero Slope. This temporal offset is not in a direction that would be expected if
changes were propagated with a slow moving current. The discrepancy is hypothesized to be a
product of age model uncertainty, rather than a time lag in the events themselves. This study
utilizes x-ray fluorescence (XRF) data to help constrain the age discrepancies between the two
age models independently from δ
15
N
sed
records. XRF correlation ‘stratigraphy’ revealed a 110-
year offset in the age models, which is likely related to inherent difficulties associated with
reservoir age corrections. After applying an increased reservoir age to the Soledad Basin age
model the δ
15
N
sed
records show greater synchronicity. XRF data analysis also reveals differences
in the geochemistry of the sediments at the two sites with Pescadero Slope showing increased Ti,
K, and Sr values with Soledad Basin exhibited high Ca abundance. The increased Ti, K values
are likely associated with increased detrital input from the volcanic Sierra Madre Occidental
High Sr values are hypothesized to relate to hydrothermal activity in the Gulf of California.
105

Elevated Ca in the Soledad Basin sediment is the result of higher carbonate content, likely due to
coccolithophore blooms.

4.1. Introduction
Nitrogen is an essential nutrient for life and as a result plays a pivotal role in the
biogeochemistry of the ocean. In the ocean, a primary source of nitrogen is microbial nitrogen
fixation, which harnesses abundant dinitrogen (N
2
) gas and converts it into the ammonia with the
use of the enzyme nitrogenase. The dominant mechanisms driving fixed nitrogen loss in the
ocean are denitrification, anaerobic ammonium oxidation and deposition of nitrogen in the
sediments. In the denitrification process, nitrate (NO
3
-
) is used as the primary electron acceptor
to degrade organic matter, releasing N
2
gas. Denitrification occurs in the ocean where oxygen
concentrations are sufficiently low in the water column and the sediments, and consequently
oxygen minimum zones (and anaerobic sediments) globally account for significant portion of
total denitrification. The eastern tropical north Pacific (ETNP) oxygen minimum zone (OMZ)
itself accounts for a third of global water column denitrification [DeVries et al., 2012], and will
be the focus of this study.
Understanding how the marine nitrogen cycle operates has been the subject of many
studies, some of which have suggested that the fixed nitrogen budget may be unbalanced (i.e.
sources and sinks are not equal). As technology has advanced, new methods have been
developed to measure denitrification rates, which has expanded estimated nitrogen loss flux
values. The incorporation of both concentration and isotopic measurements into three
dimensional global ocean circulation models has constrained denitrification loss to a range 120-
240 Tg N yr
-1
[DeVries et al., 2013]. Recent work using direct measurements and model results
predict N
2
fixation rates between 100-145 Tg N yr
-1
[Capone, 2001; Somes et al., 2010; Eugster
106

and Gruber, 2012, Casciotti, 2016], suggesting a balanced nitrogen budget. Both nitrogen inputs
and losses are likely influenced by episodic variations that are not well constrained [Brandes and
Devol, 2002] and need to be further understood. It has also been proposed that changes over the
past 100,000 to 1,000,000 years in the nitrogen budget might be linked to observed changes in
atmospheric carbon dioxide (CO
2
) and nitrous oxide (N
2
O), both of which are greenhouse gases;
N
2
O is an intermediary byproduct of denitrification, [Altabet et al., 1995, 1999; Ganeshram et
al., 1995, 2000; Brandes and Devol, 2002].
The goal of this study is to constrain the regional variability of denitrification in the
ETNP OMZ. Regional for the purpose of this work refers to the northern zones of the ETNP
OMZ near the Mexican margin and Baja California Sur. The previous Chapters discuss local
changes in the ETNP OMZ at the Pescadero Slope, Gulf of California and the Santa Monica
Basin, California Borderland, and this Chapter will delve into a regional assessment of
denitrification and OMZ fluctuations by comparing nitrogen isotopes of bulk organic matter
between sites: the Pescadero Slope and the Soledad Basin, near the Pacific coast of Baja
California. We use x-ray fluorescence to help correlate time-horizons between these sites.
To gain insight into longer-term OMZ intensity, and thus denitrification, variations,
recent studies have utilized δ
15
N

of bulk sedimentary matter (where δ
15
N
sed
‰ =
[(
15
N/
14
Nsample)/(
15
N/
14
Nstandard) - 1] * 1000, and the standard is atmospheric N
2
) from
laminated sediments as a proxy for monitoring changes in water column denitrification to extend
our understanding of fluctuations in OMZ intensity of centennial and decadal scales [Deustch et
al., 2014, Tems et al., 2015, Tems et al., submitted]. Due to preferential fractionation of nitrogen
isotopes during water column denitrification, the δ
15
N
sed
proxy is a useful tracer of OMZ
conditions. It can be used to track OMZ intensity fluctuations at locations where complete nitrate
107

utilization occurs in the photic zone [Altabet and Francios, 1994], sedimentation rates are high
[Thunell et al., 2004], bottom water oxygen content is low [Prokopenko et al., 2006], δ
15
N
sed
is
not diagenetically altered [Altabet et al., 1999; Lehmann et al., 2002; Thunell et al., 2004;
Prokopenko et al., 2006, Deutsch et al., 2014; Tems et al., 2015], and nitrite oxidation is not
significantly influencing the isotopic signature of thermocline nitrate [Casciotti et al., 2013].
These criteria are met and/or assumed to be true along the western coast of North America at
both the Pescadero Slope and Soledad Basin.
Investigations into variations in the intensity of the OMZ in Soledad Basin are not
unprecedented. Previous studies have examined variations between the deposition of laminated
and bioturbated sediments and constrained the transitions between the regimes with radiocarbon
dates. This work showed a gradual transition from bioturbated glacial sediment to laminated
Holocene sediment from 11,300 to 10,000 years and suggests patterns of regional productivity
and ventilation over the past 60,000 years [VanGeen et al., 2003]. While that study provided
valuable information about the most recent glacial interglacial transition, a high-resolution
geochemical analysis of Holocene sediments is used to evaluate OMZ intensity fluctuations at a
much finer (annual to centennial) temporal scale and assess nitrogen dynamics in the basin.
X-ray fluorescence (XRF) analysis provides a method to study geochemical and
sedimentological changes in the sediment characteristics at a very high spatial resolution and in a
non-destructive manner. This well established method provides a means to investigate
geochemical changes in sediment and rocks on decadal, annual, and even subannual scales which
enable insight into paleoenvironmental changes into depositional history, water column
processes, and paleoproductivity. XRF analysis works by exciting the sediment by incident X-
radiation, which results in the ejection of electrons from the inner atomic shells. Electrons from
108

the outer shells fill the vacancies created while also emitting energy in the form of X-radiation
[Jenkins and DeVries, 1970]. The atoms from specific elements emit characteristic energy and
wavelength, spectra, which enables element recognition and estimation of elemental abundance
[Rothwell and Rack, 2006]. In addition to paleoenvironmental reconstruction, correlations of
XRF elemental analyses are commonly used to correlate sediment cores [Hennekam and
DeLang, 2012; Bahr et al., 2005].
Our effort to investigate the regional changes in denitrification in the ETNP and OMZ
intensity and its temporal coherence is constrained by our ability to correlate synchronously
between two sites 470 km apart. We use
210
Pb,
14
C chronology and patterns in δ
15
N, potassium
(K), titanium (Ti), calcium (Ca), strontium (Sr), and barium (Ba) from the Pescadero Slope and
Soledad Basin sediment cores to evaluate the validity of the respective core age models and
identify similar events at both sites.

4.2. Study Sites
Soledad Basin, also known as San Lázaro Baisin, is located on the edge of the continental
shelf, 45 km to west of Baja California Sur, and is a tectonically formed basin (Figure 4.1). The
basin is approximately 85 km long and 35 km wide with the maximum depth near 540 m
[Silverberg et al., 2004]. It is characterized by a flat bottom and restricted water flow due to a
sill at 290 m [vanGeen et al., 2003]. Soledad Basin lies below the California Current System
flowing southward towards the equator and the California Undercurrent, which flows northward
towards the poles [Sañudo-Wilhelmy et al., 2012]. Strong northerly winds induce offshore
Ekman transport and subsequent upwelling in the winter and spring months, with upwelling
dissipating during late summer and fall and the basins proximity to costal area result in relatively
109

continuous high productivity which influences the ecosystem [Silverberg et al., 2004; Lluch-
Belda et al., 2000] and helps maintain a well-developed oxygen minimum zone below which
laminated sediments are deposited. Sediment particulates from the basin are dominated by
marine snow, containing primarily fecal pellets, vacant pteropod shells, and scattered
foraminifera. Sediment trap studies have determined that the range in total mass flux of
sediment is 63-587 mg m
-2
d
-1
and carbon fluxes range between 9 and 40 mgC m
-2
d
-1
. Organic
carbon content ranges from 5.7 to 14% depending on the season and C:N ratios varied between
7.4 and 12.7 [Silverberg et al., 2004]. Pescadero Slope C:N values range between 9-11 [Tems et
al., 2015].
Pescadero Basin is located near the mouth of the Gulf of California and, like Soledad
Basin, provides an excellent study site for reconstructing changes in OMZ fluctuations (Figure
4.1). Strong northwestern winds cause Ekman transport in the Gulf, which intensifies upwelling
and stimulates high productivity on the eastern side of the Gulf from November to March
[Staines-Urías et al., 2009; Douglas et al., 2007]. High productivity coupled with moderate rates
of deep thermohaline ventilation result in the formation of a well-developed oxygen minimum
zone between 100 and 800 meters [Fernández-Barajas et al., 1994]. Where the OMZ impinges
on the basin slope organic-rich (2-5 wt. % C) laminated sediments are deposited. The organic
matter deposited is primarily of marine origin which is evidenced by C:N ratios of 9–11, a zero
intercept plot of weight percent nitrogen versus weight percent carbon [Tems et al., 2015,
Chapter 2], and particulate organic matter (POM) δ
13
C values of -22.1 to -21.2‰ [Tems et al.,
submitted, Chapter 3].

110

4.3. Methods
4.3.1 Core Collection
Two subcores of one Soutar box core (collectively refereed to as SOLE-BC) and a
gravity core (SOLE-GC1) were collected in from Soledad Basin. SOLE-BC was collected in
2007 at 25° 13′ N and 112° 43′ W at 540 m water depth [Deutsch et al., 2014]. SOLE-GC1 was
collected at 25° 12.49′ N and 112° 42.27′ W at 544 m. Bottom water [O
2
] was measured to be 0
µM at this location [Chong et al., 2012]. Visual evaluation of variation in sediment color and the
presence of millimeter scale cocolithophore layers confirmed low bottom water [O
2
] and the
inhibition of bioturbation by benthic macrofauna.
One multicore (PESC-MC1, 64.5) and two gravity cores (PESC-GC1, 137.7 cm and
PESC-GC3, 204.5) were collected on the slope of the Pescadero Basin during the same
oceanographic cruise in 2009. PESC-MC1 was collected at 24°16.88’N and 108°11.79’W in 616
m water depth. PESC-GC1 and PESC-GC3 were collected at 24°16.160’N, 108°11.599’W and
24°16.759’N, 108°11.700’W and at 601 m and 620 m, respectively. Bottom water [O
2
] is ~0.4
µM at these locations [Tems et al., 2015]. Water column and sediment pore water chemistry at
both of these sites is described in Chong et al. [2012], Prokopenko et al. [2011], and Townsend-
Small et al. [2014].
All cores from both sites were split lengthwise and sampled at 3 mm resolution. PESC-
MC1 and PESC-GC1 were digitally scanned at high-resolution and PESC-GC3 was x-
radiographed confirming the presence of laminated sediments throughout all three Pescadero
cores. Soledad Basin cores were visually examined to confirm the presence of laminated
sediments.

111

4.3.2 Age Model
Age models for the Soledad cores were constructed by the same methods used to
construct the Pescadero Slope age models as discussed in Chapters 2 and 3. The age model for
SOLE-BC1 was constructed from 28
210
Pb measurements in the upper 25 cm of the core and 3
measurements between 70 and 72 cm. Measurements were determined on 5-10 gram sediment
splits using a Princeton Gamma-Tech Germanian well detector. The age of the sediment was
determined by dividing the integrated mass of sediment at a give depth by the accumulation rate
(88 ± 11 mg cm
-2
yr
-1
). For comparison, the accumulation rate at Pescadero Slope was 77.4 mg
cm
-2
yr
-1
. The integrated mass for this core was calculated from salt-corrected porosity
measurements while the accumulation rate was estimated from exponential decay regressions of
excess
210
Pb [Deutsch et al., 2014] (Figure 4.2).
The age model for SOLE-GC1 was based on 26 radiocarbon measurements of
sedimentary organic carbon that were graphitized [Xu et al., 2007] and analyzed at the National
Ocean Sciences Accelerator Mass Spectrometry (NOSAMS) facility at Woods Hole
Oceanographic Institution. After analysis, both blank and δ
13
C corrections were applied to the
measurements to calibrate the radiocarbon age following the convention of Stuiver and Polach
[1977] and Stuiver [1980]. Five measurements from SOLE-BC1 were used to constrain the local
reservoir age. This was done in an identical manner as the Pescadero Slope gravity core (PESC-
GC3) discussed in Chapter 3. An age correction was applied using Calib 7.1 [Stuiver and
Reimer, 1993], which included a local reservoir age (relative to the global ocean reservoir age,
ΔR) of 720 ± 68 years. The local reservoir age was estimated by averaging the radiocarbon ages
obtained from SOLE-BC older than 1950 AD (n=5) after each was corrected by the global
112

reservoir age for the year the sediment was deposited as determined by the
210
Pb age model. The
global reservoir age was determined from Marine13 [Reimer et al., 2013] and error in ΔR was
calculated as the standard deviation between the average reservoir corrections from SOLE-BC1.
Upon applying our best estimate of reservoir age correction, a linear regression (r
2
= 0.97) was
fit to the corrected radiocarbon ages versus depth in the core, where depth was the independent
variable, to determine the accumulation rate of 1.22 mm yr
-1
(Table 4.1). This, of course, implies
that the sedimentation rate was constant (an assumption we also make in our interpretation of the
Pescadero age model). Further details of the construction of both the
210
Pb and radiocarbon age
models for the Pescadero Slope are outlined in Chapters 2 and 3 [Deutsch et al., 2014; Tems et
al., 2015; Tems et al., submitted].

4.3.3 Nitrogen Isotope Measurements
SOLE-BC was sampled by extruding the core at 3 to 5 mm intervals up to a depth of 45
cm on board the B/O Francisco de Ulloa. The two subcores were aligned through calcium and
iron X-ray fluorescence (XRF) measurements of freeze-dried sediment. δ
15
N of bulk organic
matter from SOLE-BC and SOLE-GC1 were measured on a Carlo Erba elemental analyzer
interfaced with a VG Optima Stable Isotope Ratio Mass Spectrometer at the University of South
Carolina. Urea (δ
15
N = 0.10‰) was used as a working standard to normalize the data. Replicate
samples indicated analytical error of less than 0.1‰ [Deustch et al., 2014].
The Pescadero Slope cores were sampled at 3 mm intervals and were prepared for bulk
δ
15
N
sed
by drying and grinding to a fine powder. The samples from PESC-MC1 were analyzed
with a PDZ Europa ANCA-GSL elemental analyzer interfaced to a PDZ Europa 20-20
continuous flow isotope ratio mass spectrometer (IRMS) at the UC Davis Stable Isotope Facility.
113

The references used to normalize the data were G-11 (nylon, δ
15
N of -9.77‰), G-12 (glutamic
acid enriched, δ
15
N of 45.31‰), G-13 (bovine liver, δ
15
N of 7.72‰), and G-9 (glutamic acid,
δ
15
N of-4.26‰). PESC-GC3 samples were measured for δ
15
N
sed
at the University of South
Carolina on a Euro Elemental Analyzer interfaced to a GV Isoprime continuous flow IRMS. The
reference standards used to normalize the data were N-1 (δ
15
N = 0.4‰), N-2 (δ
15
N = 20.41‰),
N-3 (δ
15
N = 4.7‰), and USGS-40 (δ
15
N = -4.52‰). An inter-laboratory calibration showed that
the measurements at each facility were within analytical error of each other. Uncertainty was
determined by averaging the standard deviation between replicate samples (n=15) and was
calculated to be 0.17‰.

4.3.4 X-ray Florescence Analysis
Soledad Basin
XRF analysis of SOLE-GC1 was completed with an InnovX Delta Premium X-ray
Fluorescence Analyzer at Lamont Doherty Earth Observatory at Columbia University in 2010.
The scanner was calibrated against National Institute of Standards and Technology (NIST)
standards. The scanner provided elemental data for P, S, Cl, K, Ca, Ti, Cr, Mn, Fe, Co, Ni, Cu,
Zn, As, Se, Rb, Sr, Zr, Mo, Ag, Cd, Sn, Sb, I, Ba, Hg, and Pb.

Pescadero Slope
X-ray florescence (XRF) of the PESC-GC3 was measured using an AvaaTech XRF Core-
Scanner at ETH Zurich (Swiss Federal Institute of Technology). The core-scanner X-ray source
was a Rhodium anode and was calibrated using an internal standard, SARM4. Scans were
completed at 10kV, 30kV, and 50 kV to measure Al, Si, P, S, Cl, K, Ca, Rh, K, Ca, Ti, Cr, Mn,
114

Fe, V, Ba, Ni, Cu, Zn Ga, Br, Rb, Sr, Y, Zr, Nb, Mo, Pb, Bi, Ag, Cd, Sn, Te, Ba. Analysis of
XRF data for this study will be constrained to titanium (Ti), potassium (K), calcium (Ca),
strontium (Sr), and barium (Ba) at each site. These elements were measured in both the
Pescadero and Soledad gravity cores and their counts per sediment interval were well enough
above background to ensure the validity of the measurements and potential. Counts were
conducted over 20 second time intervals at 0.5 mm intervals. The data presented here is a six-
point weighted average, which was used to reduce noise and integrate the measurements to a
similar sampling interval as the δ
15
N
sed
measurements.

4.4 Results
4.4.1 Soledad Basin δ
15
N
sed
and Periodicity
The δ
15
N
sed
record from the Soledad Basin fluctuates between 9.1 and 10.9‰ over the
past 1600 years (Figure 4.3). In SOLE-BC1, values vary between 9.1 and 10.5‰ exhibiting a
similar range of fluctuations as seen in the Pescadero Slope record (8.1-10.5‰). The record
from SOLE-GC1 deposited before 1800 AD, however, exhibits a much smaller range of values
with minimum values of 10.0‰ and maximum values of 10.9‰. Transitions between low
δ
15
N
sed
values, which represent a less intense OMZ, and high δ
15
N
sed
, which represent a more
intense OMZ, occur rapidly and are succeeded by gradual contractions of the OMZ. This pattern
(rapid intensification of OMZ, gradual oxygenation) was also found in the Pescadero Slope data.
Spectral analysis of the Soledad Basin δ
15
N
sed
reveals decadal, multi-decadal, and
centennial periodicities. Peaks at or above the 95% confidence level (Figure 4.4a) are found at
40, 194, and 324 years. Spectral analysis of Pescadero Slope δ
15
N
sed
(Chapter 3) exhibits
periodicities at 65, 86, 102, and 232 years with 95% confidence (Figure 4.4b). Wavelet analysis
115

indicates Soledad Basin periodicities are strongest between 230 and 350 AD, 650 and 700 AD,
850 and 1100 D, 1300 and 1400 AD, and 1500-1600 AD (Figure 4.4c). Pescadero periodicities
appear strongest between 1200 and 1450 AD (Figure 4.4d).

4.4.2 X-Ray Fluorescence Analysis
Soledad Basin
In the SOLE-GC1, Ti fluctuates between 1,300 and 2,300 counts with an average value
of 1,800 counts and an average error of ± 40 counts for each measurement (Figure 4.5a).
Significant shifts in Ti trends occur near 1630 AD, 1560 AD, 1450 AD, 1230 AD, 950 AD, and
540 AD (when ΔR = 720 years). Counts of K vary between 11,800 and 15,600 counts, with an
average of 14,000 counts and an average error of ±150 counts per measurement and show
similar timing of trend shifts as Ti abundance (Figure 4.5b). Ca counts range from 102,200 and
223,200 counts, with an average value of 171,500 counts and an average measurement error of
±1,050 counts (Figure 4.5c). Ca and Sr abundance exhibits significant trend changes near 1600
AD, 1550 AD, 1400 AD, 1250 AD, 950 AD, 600 AD, and 450 AD. Sr varies between 460 and
750 counts with an average value of 680 counts and an average error of ±5 counts for each
measurement (Figure 4.5d). The range of Ba is between 250 and 480 counts and an average of
330 counts (Figure 4.5e). The average error on each measurement was ±14 counts. Ba
abundance trends show significant shifts at 1600 AD, 1500 AD, 1250 AD, 900 AD, 775 AD, and
500 AD.

116

Pescadero Slope
XRF values for Ti at the Pescadero Slope site has a much greater and dynamic range;
varies between 170 and 5,600 counts with an average of value of 3500 counts, an average error
of 290 counts (Figure 4.6a). Significant shifts in Ti trends occur near 1600 AD, 1330 AD, 1200
AD, 1090 AD, and 1000 AD. K varies between 250 and 15,400 counts with an average value of
8680 counts and an average error of 180 counts with significant shifts in abundance occurring
synchronously with Ti shifts (Figure 4.6b). Ca counts ranged between 2,300 and 56,300, with an
average value of 9,200 counts and an average error of 170 counts (Figure 4.6c). Periods of
elevated Ca are found between 1590 and 1610 AD, 1525-1550 AD, 1470-1480 AD, 1360-1400
AD, 1180-1300 AD, and 1050-1130 AD. Sr counts fluctuate between 70 and 2,360 counts, with
a mean count of 890 and average error of 40 counts (Figure 4.6d). Significant changes in Sr
trends occur at 1600 AD, 1580 AD, 1300 AD, 1200 AD, 1080, and 1000 AD. Ba varies between
80 and 660 counts, with an average of 410 counts and an average error of 50 counts (Figure
4.6e). Major shifts in Ba abundance trends are also found to occur at 1600 AD, 1560 AD, 1200
AD, 1100 AD, and 1000 AD.

A Comparison Between Study Sites
In order to compare the variation in counts of each element between the Gulf of
California and the Baja California Sur Pacific margin, terrigenous and biogenic elements are
directly compared. Comparing Ti and K, shows that in the Soledad Basin these elements show
far less variation than at the Pescadero Slope. Ti counts are low and K counts have more
moderate values but neither shows a wide range of variability. The relationship between Ti and
K is defined by the linear regression y = 258.8*x -4543 (r
2
= 0.70), where y refers to Ti counts
117

and x refers to K counts. Pescadero Slope shows a much wider range of variation in both
elements with significantly higher Ti counts. The relationship between Ti (y) and K (x) counts is
defined by the following equation: y = 22.55*x -10690 (r
2
= 0.58) (Figure 4.7). A plot of Ca and
Sr, the biogenic components, shows distinctly different trends and elemental compositions of
both elements. In Soledad Basin, which is characterized by high Ca counts, the relationship
between the elements is y = 0.1765*x – 613.1 (r
2
= 0.64) where y refers to Ca counts and x to Sr
counts. Pescadero Slope exhibits lower Ca counts and significantly higher Sr counts with a wide
range of variability, which exhibit the following linear relationship: y = 0.2834*x + 1088 (r
2
=
0.78), when y represents Ti counts and x represents Sr counts (Figure 4.8). A cluster of data
points exhibit exceeding high Ti counts and low Sr counts occur is circled on Figure 4.8. These
points correspond to 1332-34 AD in the record which is not an anonymously period in the
δ
15
N
sed
record. The ratio of Ti to K based on a generalized composition of crustal rocks ranges
between 0.23–0.38 and the average ratios at Soledad Basin and Pescadero Slopes are 0.13 and
0.40, respectively. The average ratio of Sr to Ca found in modern carbonates is 0.00170-0.0151
[Taft and Harbaugh, 1964] which is similar the Soledad Basin ratio of 0.004 but an order of
magnitude different to the Pescadero Slope ration of 0.10.

4.5 Discussion
One major question we seek to address is whether the Pescadero Slope and Soledad Basin
records show synchronous similarities in changes in the OMZ (Figure 4.9). The side-by-side
records shows that there are multiple peaks in δ
15
N
sed
values, which may be cross-correlated
between the cores. Yet, given the time-scale adopted for each core, these features are generally
offset; occurring earlier in the Soledad record and later in the Pescadero record. This could be
118

due to either (1) a time lag between OMZ expansion at the Pescadero Slope and Soledad Basin
or (2) a product of uncertainty and assumptions chosen to create the respective age models of
each core. δ
15
N
sed
values from the most recently deposited sediment (1850 to 2007) from both
locations appear to show similar trends in OMZ contraction and expansion, suggesting that the
variation in OMZ expansion between the gravity core records is a product of uncertainties and
assumptions inherent in the construction of radiocarbon age models as well as possible coring
artifacts. The Pescadero Slope age model, as discussed in detail in Chapter 3, was constructed by
applying a marine correction with a ΔR of 508 years. This local reservoir age (ΔR) was
independently constrained through radiocarbon dates in the multicore and gravity core and with
this inter-comparison, the offset was < ± 30-year. Additionally, the PESC-GC3 was
geochemically correlated to PESC-MC1 to reduce uncertainty and varve counts further
confirmed the validity of the age model. Uncertainty is an inherent characteristic of sediment
core age models, and due to the extensive work to verify the Pescadero Slope radiocarbon age
model and its consistency with the
210
Pb age model, for the purpose of this study we will assume
the Pescadero core represents the ‘correct’ age and will evaluate potential variation in the SOLE-
GC1 age model.
One significant assumption that has the ability to greatly influence the age assignments is
the local
14
C reservoir age. This age is difficult to constrain and it is standard practice to
estimate the average local reservoir age by comparing radiocarbon and
210
Pb calculated ages.
However, applying one value down-core assumes that the local reservoir age is constant. As can
be seen with reconstructions of the global reservoir age from Marine13 [Reimer et al., 2013] and
surface water reservoir age evaluation [Butzin et al., 2012], and sub-surface reservoir ages [De
La Fuente et al., 2015], the reservoir age has very likely fluctuated during the Holocene. As a
119

result, it was not surprising to discover the offset between increased OMZ strength in the records
is similar but not identical for each interval (Figure 4.9).
The ΔR calculated based on both radiocarbon and
210
Pb calculated ages for the Soledad
Basin was 720 ± 68 years. Previous studies have found the ΔR of surface water in the region
from planktonic foraminifera to be 225 ± 60 years [Stuiver and Brazunias, 1993], 233 ± 60 years
[Ingram and Southon, 1996], and 200 ± 100 years [Marchitto et al., 2010] however upwelling
regimens can bring significantly older carbon to the surface which has been observed in the
eastern Pacific during the Holocene [Fontugne et al., 2004].
In order to constrain the potential offset between in age models between the two sites, we
investigated correlations in elemental abundance from XRF analysis independent from
correlations between δ
15
N
sed
. Correlating sediment cores with elemental data is not
unprecedented and is commonly used as a technique to correlate sedimentological changes in
cores and boreholes for petroleum exploration.
Correlations were first explored between the terrigenous components (Ti and K). While
the XRF counts of elements between the two sites have different ranges due to differences in the
weight % detrital content, one might assume that significant shifts in sedimentation from
continental material should be observed regionally. Peaks in Ti and K were visually correlated
between the two records based on the peaks themselves and the trends preceding and succeeding
the peaks. Using this method, five correlation peaks were observed (Table 4.2) with an average
offset of 113 years; Soledad Basin sediments appear older than Pescadero Slope sediments.
The calibration of the Soledad Basin radiocarbon age model was re-analyzed to account
for this offset between the two records by increasing the ΔR to 830 years and increasing
uncertainty of this value to 110 years, in order for it to encompass the original calculated ΔR
120

value derived from radiocarbon and
210
Pb measurements. This improves the synchronicity
between the Pescadero Slope and Soledad Basin δ
15
N
sed
records (Figure 4.9).
The spectral analysis data presented in the results section would not be significantly
altered by this age change, since the analysis was run only on SOLE-GC and not SOLE-BC, thus
while the age was altered the relationship between δ
15
N
sed
fluctuations was not. The overlap
between similar values in from SOL-BC and SOL-GC1 further supports the adjustment of local
reservoir age (Figure 4.9). Small discrepancies between intensification in OMZ strength is still
observed with the adjusted reservoir age radiocarbon age model, however, these are likely due to
deviations from an average reservoir age to a more dynamic reservoir age.
Coherence between OMZ fluctuations in the two records is additionally supported by
spectral analyses, which reveal similar periodicities in OMZ intensity variations (Pescadero
Slope: 68, 86, 102, and 232 year; Soledad Basin: 40, 62, 93, and 194 year periodicities; Figure
4.4a/c). The 40 and 62 year periodicities are likely related to the Pacific Decadal Oscillation as
with the 68 year periodicity found in the Pescadero Slope OMZ intensity record [Baumgartner et
al., 1992; Meko et al., 1992; Sharp, 1992; Schlesinger and Ramankutt, 1994; Thurow and
Schaaf, 1995; Minobe, 1997; Pike and Kemp, 1997; MacDonald and Case, 2005]. The 93-year
periodicity is likely related to the Gleissberg cycle, similar to the 86 and 102-year periodicities at
Pescadero Slope [Ogurtsov et al., 2002].
The range of δ
15
N
sed
values at the Pescadero Slope and Soledad Basin is similar between
1850 and 2007 (8.5-10.0‰ and 9.0-10.6‰, respectively), suggesting coherence between the
records and age models over the past 150 years. While the range of values remains similar (8.1-
10.5‰) at the Pescadero Slope from 900-1850 AD, the range of δ
15
N
sed
values at Soledad Basin
is significantly diminished by more than 0.6‰ to values of 10.0-11.0‰ (Figure 4.9). Despite the
121

change in δ
15
N
sed
, similar peaks can be correlated between the Pescadero Slope and Soledad
Basin records suggesting that local variations at each site are overprinted by regional events.
The dampening of the δ
15
N
sed
range at the Soledad Basin may be the product of
significant variations in local biological productivity. A key difference between the Pacific
Soledad Basin and the Gulf of California Pescadero Slope is the documentation of major
coccolithophore blooms in the Soledad Basin. Coccolithophore blooms, phytoplankton that
precipitates calcium carbonate skeletons, form white laminae in sediment cores and are observed
in the Soledad Basin cores [VanGeen et al., 2003] but are not present in the cores collected at
Pescadero Slope. These blooms occur most often in the summer and El Niño conditions when
warm water is transported northward from the ETNP OMZ [Silverberg et al., 2004] and could
contribute to the near continuous productivity found in the basin [Lluch-Belda et al., 2000;
Silverberg et al., 2004]. If increased carbon export from the photic zone contributes to increased
OMZ intensity as suggested by geochemical comparisons at the Pescadero Slope (Chapter 3),
additional carbon export from coccolithophore blooms may dampen the range of variation
observed in OMZ intensity, and thus δ
15
N
sed
variability, at Soledad Basin. Differences in
circulation, driven by the bathymetry of each site, may also contribute to the elevated δ
15
N
sed
values observed at Soledad Basin. Here, the premise is that the water residence time on the
Soledad shelf/slope may be influenced by sluggish mixing.
As noted in the results section, significantly different total counts of Ti and K were
measured at the two different sites. This variation is likely due to difference in the fraction of
detrital input deposited into the basins. Pescadero Slope, on the eastern side of the Gulf of
California, exhibits significantly elevated counts of Ti, which is like a result of proximity to the
volcanic Sierra Madre Occidental (mountain range) on the adjacent Mexican mainland. Five
122

rivers, the Sonora, Matape, Yaqui, Mayo and Fuerte, drain the mountain range into the Gulf of
California. The Fuerte River is the most southern, and although it still lies significantly north of
the Pescadero Slope at entering the Gulf near 25.5°N and 108.6°W, the input of detrital sediment
is clearly diluting the biogenic sediment accumulation at this location [Dean et al., 2004].
Terrigenous influx from Baja California to into Pacific basins is minimal due the much smaller
volume of landmass from which the material would originate [Baba et al., 1991]. Additionally,
the geology in western Baja California Sur near the Soledad Basin is primarily compassed of
Pliocene and Quaternary sedimentary rocks [Hausback, 1984], which will have a different
geochemical composition compared to the volcanic rocks of the Sierra Madre Occidental near
Pescadero Slope. Consequently, the ratio of Ti to K at Pescadero Slope closely resembles that
of average crustal rocks (0.4 (PS) compared to 0.23-0.38 (CR)). A significant deviation from the
Ti versus K trend (Figure 4.7) at Pescadero Slope occurs between 1332 and 1335 AD (circled
data in Figure 7). During this interval, high Ti abundance corresponds to low K abundance. The
values correspond to elevated δ
15
N
sed
, however, this does not appear to be a significant deviation
in the OMZ intensity record (Figure 4.9, peak 4).
Calcium and Sr also exhibit significant differences in abundance between Pescadero
Slope and Soledad Basin. As discussed in relation to the dampened variability in OMZ strength
in Soledad Basin, increased abundance of Ca in the basin is not unexpected due to the presence
of calcium carbonate coccolithophore blooms at the site, and the generally high CaCO
3
content.
The high levels of Sr found in the sediments at Pescadero Slope, however, were surprising. Sr
abundance is typically associated with Ca abundance, due to the similar bioreactivity. Mostly, it
is thought that Sr is incorporated into carbonates. One hypothesis is that this elevated Sr could
be introduced through hydrothermal activity in the Gulf of California. Hydrothermal activity is
123

present in the Gulf due to active spreading presently occurring as part of the East Pacific Rise
tectonic system [Curray et al., 1979; Weber and Jorgensen, 2002]. Hydrothermal fluids diffuse
through the seafloor when temperature are below 200°C [Von Damm et al., 1985] and from
mounds/chimneys rising above the seafloor [Webber and Jorgensen, 2002] and hydrothermal
activity is a source of Sr to the ocean [Palmer and Edmond, 1989]. Studies examining the ratio
of Sr to Ca in coccolith Emiliania huxleyi indicate that the ratio is linearly related to Sr/Ca of the
seawater [Langer et al., 2006], thus suggesting the ratio of these elements in the sediments would
be influenced by increased Sr concentrations in seawater from hydrothermal sources.
Alternatively, protozoans, known as Acantharians, are unique organisms that precipitate a
skeleton made of the mineral celestite (SrSO
4
). The skeletons are thought to dissolve readily due
to under saturation of this mineral in the ocean [Beers et al., 1975; Deckker, 2004]. Sr
availability in the Gulf of California may locally reduce Sr under-saturation leading to reduced
dissolution and thus greater deposition in the sediments. Studies have shown that some Ca and
Sr input is not associated with calcium carbonate and the increase could be partially related to
detrital input [Dean et al., 2004]. We have yet to distinguish which is the most likely source of Sr
enrichment at the Pescadero site.

4.6 Conclusions
Fluctuations in the intensity of the eastern tropical north Pacific (ETNP) oxygen
minimum zone (OMZ) at Soledad Basin and the Pescadero Slope show temporal coherence
between the two sites through the past 1,500 years, yet also illustrating shorter time-scale local
variations. The OMZ at both locations could be self-regulating, as evidenced by fluctuations that
remain bracketed between fixed maximum and minimum δ
15
N
sed
values. The severity of OMZ
124

strength fluctuations varies between the two locations, with higher δ
15
N
sed
values in Soledad
Basin suggesting increased denitrification and OMZ intensity along the Baja California Pacific
margin. This could be the result of near continuous productivity in this region due to year-round
upwelling [Silverberg et al., 2004], opposed to high productivity seasonality of the Gulf of
California [Dean et al., 2004]. The higher productivity and coccolithophore blooms, unique to
Soledad Basin, supports greater deposition of calcium carbonate to the sediments, which is
evidenced by significantly higher Ca content in the sediment as shown in XRF analysis. The
Gulf of California exhibits higher abundances of Ti and Sr, which we suggest are related to
greater terrigenous input and potentially due to hydrothermal activity. Finally, we hypothesize
that standard assumptions made in order to determine the local reservoir age (ΔR) lead to an
incorrect offset between regional variations in δ
15
N
sed
and elemental abundance trends.
Correlation of XRF elemental abundance counts between cores allow for re-evaluation of the age
models and a better understanding of regional events in the eastern tropical north Pacific and
suggests an 110 year increase in the local reservoir age of Soledad Basin (ΔR
corr
= 830 ± 110
years). This study brings to light the importance of developing new methods in order to
quantitatively account for changes in reservoir ages in sediment cores in order for accurate age
model and therefor paleoenvironmental interpretations.

125

Acknowledgements
This study was possible through collaborations between University of Southern
California, University of South Carolina, Lamont-Doherty Earth Observatory of Columbia
University, and ETH Zurich. I would like to first and foremost thank Alexander van Geen
(Lamont-Doherty Earth Observatory) for collaborating on this final chapter and supplying the
δ
15
N
sed
and x-ray fluorescence data from Soledad Basin for evaluation. This analysis would not
have been possible without this contribution. I would also like to thank Bob Thunell and Eric
Tappa at University of South Carolina for assisting with the δ
15
N
sed
measurements of both cores
and Gerald Haug (ETH Zurich) for collaboration on x-ray florescence analysis of the Pescadero
Slope gravity core. Additionally, I would like to thank Deborah Khider and Nick Rollins at
University of Southern California for their input and helpful discussions on this final chapter.

126

References
Altabet, M. A., and R. Francois (1994), Sedimentary nitrogen isotopic ratio as a recorder for
surface ocean nitrate utilization, Global Biogeochem. Cycles, 8(1), 103–116,
doi:10.1029/93GB03396.
Altabet, M.A., R. Francois, D.W. Murray, and W.L. Prell (1995), Climate-related variations in
denitrification in the Arabian Sea from sediment
15
N/
14
N ratios, Nature, 373(6514), 506-509,
doi:10.1038/373506a0.
Altabet, M. A., C. Pilskaln, R. Thunell, C. Pride, D. Sigman, F. Chavez, and R. Francois (1999),
The nitrogen isotope biogeochemistry of sinking particles from the margin of the eastern
North Pacific, Deep Sea Res. Part I, 46, 655–679, doi:10.1016/S0967-0637(98)00084-3.
Baba, J., C.D. Peterson, H.J. Schrader (1991), Modern fine-grained sediments in the Gulf of
California, In: The Gulf and Peninsular Province of the Californias (eds. Dauphin, J.P., and
B.R.T. Simoneit), American Association of Petroleum Geologists Memoir, 47, 569–587.
Bahr, A., F. Lamy, H. Arz, H. Kuhlmann, and G. Wefer (2005), Late glacial to Holocene climate
and sedimentation history in the NW Black Sea, Marine Geology, 214(4), 309-322.
Baumgartner,T., V. Ferreira, J. Mey, and A. VanGeen (2014), Centennial changes in North
Pacific anoxia linked to tropical trade winds, Science, 345 (6197), 665-668,
doi:10.1126/science.1252332.
Beers, J. R., F. M. H. Reid and G. L. Stewart (1975), Microplankton of the north Pacific central
gyre: population structure and abundance, Internationale Revue Gesamtes Hydrobiologia 60,
607–638.
127

Brandes, J. A., and A. H. Devol (2002), A global marine-fixed nitrogen isotopic budget:
Implications for Holocene nitrogen cycling, Global Biogeochemical Cycles, 16(4), 1120,
doi:10.1029/2001GB001856.
Butzin, M., M. Prange and G. Lohmann (2012), Readjustment of glacial radiocarbon
chronologies by self-consistent three-dimensional ocean circulation modeling, Earth
Planetary Science Letters, 317-318, 177–184.
Capone, D. G. (2001), Marine nitrogen fixation: what's the fuss? Current opinion in
microbiology, 4(3), 341-348.
Casciotti, K. L., C. Buchwald, and M. McIlvin (2013), Implications of nitrate and nitrite isotopic
measurements for the mechanisms of nitrogen cycling in the Peru oxygen deficient zone,
Deep-Sea Research I, 80: 78-98.
Casciotti, K. L. (2016), Nitrogen and oxygen isotopic studies of the marine nitrogen cycle,
Annual Reviews of Marine Science, 8, 379-407, doi:
10.1146/annurev-marine-010213-
135052.
Chong, L. S., M. G. Prokopenko, W. M. Berelson, A. Townsend-Small, and J. McManus (2012),
Nitrogen cycling within suboxic and anoxic sediments from the continental margin of
western North America, Marine Chemistry, 128–129, 13–25,
doi:10.1016/j.marchem.2011.10.007.
Curray, J.R., D.G. Moore, J.E. Aguayo, M.P. Aubry, G. Einsele, D.J. Fornari, J. Gieskes, J.C
Guerrero, M. Kastner, K. Kelts, M. Lyle, Y. Matoba, A. Molina-Cruz, J. Niemitz, J. Rueda,
A.D. Sauders, J. Schrader, B.R.T. Simoneit, and V. Vacquier (1979), Leg 64 seeds evidence
on development of basins, Geotimes 24, 18–20.
128

Dean, W., C. Pride, and R. Thunell (2004), Geochemical cycles in sediments deposited on the
slopes of the Guaymas and Carmen Basins of the Gulf of California over the last 180 years,
Quaternary Science Reviews, 23(16), 1817-1833.
Deckker, P.D. (2004), On the celestite-secreting Acantharia and their effect on seawater
strontium to calcium rations, Hydrobiologia, 517, 1-13.
De La Fuente, M., L. Skinner, E. Valvo, C. Pelejero, and I. Cacho (2015), Increased reservoir
ages and poorly ventilated deep waters inferred in the glacial eastern equatorial Pacific,
Nature Communications, 6, 7420, doi:10.1038/ncomms8420.
Deutsch, C. W. Berelson, R. Thunell, T. Webber, C. Tems, J. McManus, J. Crusius, T. Ito, T.
Baumgartner, V. Ferreira, J. Mey, and A. VanGeen (2014), Centennial changes in North
Pacific anoxia linked to tropical trade winds, Science, 345 (6197), 665-668,
doi:10.1126/science.1252332.
DeVries, T., C. Deutsch, F. Primeau, B. Chang, and A. Devol (2012), Global rates of water-
column denitrification derived from nitrogen gas measurements, Nature Geoscience, 5(8),
547–550, doi:10.1038/ngeo1515.
DeVries, T., C. Deutsch, P. A. Rafter, and F. Primeau (2013), Marine denitrification rates
determined from a global 3-D inverse model, Biogeosciences, 10(4).
Douglas, R. O. Gonzalez-Yajimovich, J. Ledesma-Vazquez, and F. Staines-Urias (2007),
Climate forcing, primary production and the distribution of Holocene biogenic sediments in
the Gulf of California, Quaternary Science Reviews, 26, 115-129, doi:
10.1016/j.quascirev.2006.05.003.

Eugster O., and N. Gruber (2012), A probabilistic estimate of global marine N-fixation and
denitrification, Global Biogeochemical Cycles, 26, GB4013.
129

Fernández-Barajas ME, M.A. Monreal-Gómez, A. Molina-Cruz (1994), Thermohaline structure
and geostrophic flow in the Gulf of California, during 1992, Ciencias Marinas 20, 267–286.
Fontugne, M., M. Carre, I. Bentaleb, M. Julie, and D. Lavallee (2004), Radiocarbon reservoir
age variations in the south Peruvian upwelling during the Holocene, Radiocarbon, 46(2), 51-
537.
Ganeshram, R.S., T.F. Pedersen, S.E. Calvert, and J.W. Murray (1995), Large changes in oceanic
nutrient inventories from glacial to interglacial periods, Nature, 376, 755-758.
Ganeshram, R. S., T. F. Pedersen, S. E. Calvert, G. W. McNeil, and M. R. Fontugne (2000),
Glacial-interglacial variability in denitrification in the world’s oceans: Causes and
consequences, Paleoceanography, 15(4), 361–376, doi:10.1029/1999PA000422.
Hausback, B.P (1984), Cenozoic volcanic and tectonic evolution of Baja California Sur, Mexico,
In: Geology of the Baja California Peninsula (ed. F. Virgil), Pacific Section Society for
Sedimentary Geology, 39, 219-236.
Hennekam, R., and G. De Lange (2012), X-ray fluorescence core scanning of wet marine
sediments: methods to improve quality and reproducibility of high-resolution
paleoenvironmental records, Limnology and Oceanography: Methods, 10, 991-1003.
Ingram, B. L., and J. R. Southon (1996), Reservoir ages in eastern Pacific coastal and estuarine
waters, Radiocarbon, 38, 573–582.
Jenkins, R., and J.L. De Vries (1970), Practical X-ray Spectrometry, MacMillan, London, pp. 1–
190.
Langer, G., N. Gussone, G. Nehrke, U. Riebesell, A. Eisenhauer, H. Kuhnert, B. Rost, S.
Trimborn, and S. Thoms (2006), Coccolith strontium to calcium ratios in Emiliania huxleyi:
130

the dependence on seawater strontium and calcium concentrations, Limnology and
Oceanography, 51(1), 310-320.
Lehmann, M. F., S. M. Bernasconi, A. Barbiere, and J. A. McKenzie (2002), Preservation of
organic matter and alteration of its carbon and nitrogen isotope composition during simulated
and in situ early sedimentary diagenesis, Geochim. Cosmochim. Acta, 66(20), 3573–3582,
doi:10.1016/S0016-7037(02)00968-7.
Lluch-Belda, D., M.E. Hernández-Rivas, R. Saldierna-Martínez, and R. Guerrero-Caballero
(2000), Variabilidad de la temperatura superficial del mar en Bahía Magdalena, BCS.
Oceánides, 15(1), 1-23.
MacDonald, G.M. and R.A. Case (2005), Variations in the Pacific Decadal Oscillation over the
past millennium, Geophysical Research Letters, 32: L08703, doi:10.1029/2005GL022478.
Marchitto, T.M., R. Muscheler, J.D. Ortiz, J.D. Carriquiry, and A. VanGeen (2010), Dynamical
response of the tropical pacific ocean to solar forcing during he early Holocene, Science,
330, 1378-1381.
Meko, D.M. (1992), Spectral properties of tree-ring data in the United States Southwest as
related to El Niño/Southern Oscillation, in El Niño: Historical and Paleoclimatic Aspects of
the Southern Oscillation (eds. H.F Diaz and V. Markgraf), 227-242.
Minobe, S. (1997), A 50-70 year climatic oscillation over the North Pacific and North America,
Geophysical Research Letters, 24(6), 683-686.
Ogurtsov, M.G., Y.A. Nagovitsyn, G.E. Kocharov, and H. Junger (2002), Long-period cycles of
the sun’s activity recorded in direct solar data and proxies, Solar Physics, 211, 371-394.
Palmer, M.R. and J.M. Edmond (1989), The strontium isotope budget of the modern ocean,
Earth and Planetary Science Letters, 92, 11-26.
131

Pike, J. and A.E.S. Kemp (1997), Early Holocene decadal-scale ocean variability recorded in
Gulf of California laminated sediments, Paleoceanography, 12(2), 227-238.
Prokopenko, M., D. E. Hammond, A. J. Spivack, and L. Stott (2006), Impact of long-term
diagenesis on δ
15
N of organic matter in marine sediments: Sites 1227 and 1230, Proceedings
from the Ocean Drilling Program, Scientific Results, 201, 30.
Prokopenko, M., D.M. Sigman, W.M. Berelson, D.E. Hammond, B. Barnett, L. Chong, and A.
Townsend-Small (2011), Denitrification in anoxic sediments supported by biological nitrate
transport, Geochimica et Cosmochimica Acta, 75, 7180-7199,
doi:10.1016/j.gca.2011.09.023.
Reimer P.J., E. Bard, A. Bayliss, J.W. Beck, P.G. Blackwell, C. Bronk Ramsey, C.E. Buck, H.
Cheng, R.L. Edwards, M. Friedrich, P.M. Grootes, T.P. Guilderson, H. Haflidason, I. Hajdas,
C. Hatté, T.J. Heaton, A.G. Hogg, K.A. Hughen, K.F. Kaiser, B. Kromer, S.W. Manning, N.
Niu, R.W. Reimer, D.A. Richards, E.M. Scott, J.R. Southon, C.S.M. Turney, J. van der Plicht
(2013), IntCal13 and MARINE13 radiocarbon age calibration curves 0-50000 years calBP,
Radiocarbon 55(4), doi:10.2458/azu_js_rc.55.16947.
Rothwell, R. G., and F.R. Rack (2006), New techniques in sediment core analysis: an
introduction, Geological Society, London, Special Publications, 267(1), 1-29.
Sañudo-Wilhelmy, S. A., L.S. Cutter, R. Durazo, E.A. Smail, L. Gómez-Consarnau, L., E.A.
Webb, M.G. Prokopenko, W.M. Berelson, and D.M. Karl (2012), Multiple B-vitamin
depletion in large areas of the coastal ocean, Proceedings of the National Academy of
Sciences, 109(35), 14041-14045.
Schlesinger M, .E., and N. Ramankutty (1994), An oscillation in the global climate system of
period 65-70 years, Nature, 3(67), 723-726.
132

Silverberg, N., A. Martínez, S. Aguíña, J.D. Carriquiry, N. Romero, E. Shumilin, and S. Cota
(2004), Contrasts in sedimentation flux below the southern California Current in late 1996
and during the El Niño event of 1997-1998, Estuarine, Coastal and Shelf Science 59, 575-
587.
Sharp, G.D. (1992), Fishery catch records, El Niño/Southern Oscillation, and longer-term
climate change as inferred from fish remains in marine sediments, in El Niño: Historical and
Paleoclimatic Aspects of the Southern Oscillation, edited by H.F. Diaz and V. Markgraf,
379-418.
Somes CJ, A. Schmittner, and M.A. Altabet (2010), Nitrogen isotope simulations show the
importance of atmospheric iron deposition for nitrogen fixation across the Pacific Ocean,
Geophysical Research Letters, 37, L23605.
Staines-Urías, F., R.G. Douglas, D.S. Gorsline (2009), Oceanographic variability in the southern
Gulf of California over the past 400years: Evidence from faunal and isotopic records from
planktic foraminifera, Palaeogeography, Palaeoclimatology, Palaeoecology, 284 (3-4): 337-
354, doi: 10.1016/j.palaeo.2009.10.016.
Stuiver, M., and H.A. Polach (1977), Discussion: Reporting of
14
C data, Radiocarbon, 19, 355-
363.
Stuiver, M. (1980), Workshop on
14
C data reporting, Radiocarbon, 22, 964-966.
Stuiver, M., and T. F. Brazunias (1993), Modeling atmospheric 14C influences and 14C ages of
marine samples to 10,000 BC, Radiocarbon, 35, 137–189.
Stuiver, M. and P.J. Reimer (1993), Extended 14C database and revised CALIB radiocarbon
calibration program, Radiocarbon, 35:215-230.
133

Taft, W.H, and J.W. Harbaugh (1964), Modern carbonate sediments of southern Florida,
Bahamas, and Espiritu Santo Island, Baja California: a comparison of their mineralogy and
chemistry, Stanford University Press.
Tems, C. E., W.M. Berelson, and M.G. Prokopenko (2015), Particulate δ
15
N in laminated marine
sediments as a proxy for mixing between the California Undercurrent and the California
Current: A proof of concept, Geophysical Research Letters, 42, doi:10.1002/
2014GL061993.
Tems, C.E., W.M. Berelson, R. Thunell, E. Tappa, X. Xu, D. Khider, S. Lund and O. González-
Yajimovich, Sedimentary δ
15
N reveal decadal fluctuations in the intensity of the eastern
tropical north Pacific oxygen minimum zone during the last 1200 years, submitted to
Paleoceanography November 2015.
Thunell, R., D.M. Sigman, F.E. Muller-Karger, Y. Astor, R. Varela (2004), Nitrogen Isotopic
Dynamics of the Cariaco Basin, Venezuela, Global Biogeochemical Cycles, 18: GB3001,
doi:10.1029/2003GB002185.
Thurow, J., and M. Schaaf (1995), A comparison of Holocene to late Pleistocene grayvalue-time-
series of sediments from the Santa Barbara Basin (ODP Site 893) and the Guaymas Basin
(DSDP Site 480), in IGCP 374: Palaeoclimatology and Palaeoceanography From Laminated
Sediments, edited by R.J. Behl and J.P. Kennett, 33.
Townsend-Small, A., M. G. Prokopenko, and W. M. Berelson (2014), Nitrous oxide cycling in
the water column and sediments of the oxygen minimum zone, eastern subtropical North
Pacific, Southern California, and Northern Mexico (23°N–34°N), Journal of Geophysical
Research Oceans, 119, 3158–3170, doi:10.1002/2013JC009580.
134

Van Geen, A., Y. Zheng, J. M. Bernhard, K. G. Cannariato, J. Carriquiry, W. E. Dean, B. W.
Eakins, J. D. Ortiz, and J. Pike (2003), On the preservation of laminated sediments along the
western margin of North America, Paleoceanography, 18(4), 1098,
doi:10.1029/2003PA000911.
Von Damm, K. L., J.M. Edmond, B. Grant, C.I. Measures, B. Walden, and R.F. Weiss (1985),
Chemistry of submarine hydrothermal solutions at 21° N, East Pacific Rise, Geochimica et
Cosmochimica Acta, 49(11), 2197-2220.
Webber, A., and B.B. Jorgensen (2002), Bacterial sulfate reduction in hydrothermal sediments of
the Guaymas Basin, Gulf of California, Mexico, Deep-Sea Research I 49(5),827-841.

135
20˚N
25˚N
30˚N
120˚W 115˚W 110˚W 105˚W
Pescadero
Slope
Pacic Ocean
Baja California
Mexico
Gulf of California
United States
Soledad Basin
Figure 4.1. A map showing the study sites, Pescadero Slope and Soledad Basin, the eastern
tropical north Pacific. Coring locations are designated by the black dots.
136
Figure 4.2. Total
210
Pb profiles plotted against integrated mass for Soledad Basin (black) and Pescadero Slope
(white). The profiles were fitted with an exponential decay equation in order to calculate to mass accumulation rates
at each site. This figure has been published in Deutsch et al., 2014.
Integrated Mass (g cm
-2
)
210
Pb (Bq kg
-1
)
137
Average
Depth
(cm) Uncorrected

14
C

Age
(BP) Corrected
Year
(ΔR=
720
±
68) Corrected
Year
(ΔR
=
830
±
110)
15.15 1180

±
15 1849
±
71.5 N/A
25.35 1310

±
20 1733
±
87.5 1812
±
74.5
34.35 1500

±
15 1538
±
63.5 1642
±
113
46.05 1470

±
20 1563
±
67.5 1675
±
112.5
59.55 1630

±
20 1426
±
49.5 1529
±
96
70.65 1830

±
20 1279
±
64 1362
±
81
80.55 1880

±
20 1234
±
64 1325
±
88.5
83.55 2090

±
20 1034
±
84 1139
±
105
95.55 2110

±
20 1011
±
83 1122
±
107.5
105.45 2100

±
25 1022
±
86.5 1130
±
106.5
119.25 2210

±
20 904
±
82.5 1021
±
124.5
125.85 2310

±
30 803
±
82 914
±
119.5
131.25 2390

±
20 724
±
67.5 838
±
115.5
139.65 2400

±
20 715
±
67 829
±
115.5
154.35 2460

±
15 655
±
68 771
±
111.5
164.55 2450

±
25 665
±
70 781
±
113
173.85 2690

±
25 422
±
93 531
±
114
182.55 2610

±
20 510
±
79 612
±
116.5
190.05 2870

±
20 211
±
90 337
±
135
199.65 2780

±
20 314
±
86.5 441
±
129
Soledad
Basin

14
C
Age
Model
Data
Table 4.1. Data in this table includes the
14
C uncalibrated ages from their respective depths and the calibrated
ages and errors when local reservoir ages (ΔR) of 720 ± 68 years and 830 ± 110 years are applied to the
radiocarbon ages.
138
400 600 800 1000 1800 2000
9
9.5
10
10.5
11
11.5
Soledad Basin
δ
15
N
sed
(‰)
1200 1400 1600
Year AD
Figure 4.3. The δ
15
N
sed
record from the Soledad Basin compared to the radiocarbon age model constructed with a
ΔR of 720 ± 68 years (black dots), as calculated from the original data. δ
15
N
sed
varies between 9.0 and 10.9‰ over
the past 1600 years. Gray dots represent data points compared the the
210
Pb age model.
139
Soledad Basin
a.
Power
194
40
Years A.D.
Period

400 600 800 1000 1200 1400 1600 1800
50
100
200
300
400
500
Power
0.4
0.6
0.8
1
1.2
1.4
1.6
1.8
2
c.
Figure 4.4. a) Spectral analysis of δ
15
N
sed
from Soledad Basin using a Lomb-Scargle Fourier Transform in combination
with a Welch-Overlapped-Segment-Averaging (WOSA) procedure. Peaks representing 40, 194, and 324 with greater
than 95% confidence and 62 and 93 years with greater than 50% confidence. b) Spectral analysis of Pescadero Slope δ
15
N
sed
using the same methods. Periodicities with greater than 95% confidence were found at 65, 86, 102, and 232 years.
c) Results of wavelet analysis from Soledad Basin using a weighted wavelet Z-transform (WWZ) for unevenly spaced
time series indicate that the strongest periodicities occur between 230 and 350 AD, 650 and 700 AD, 850 and 1100 D,
1300 and 1400 AD, and 1500-1600 AD in this record. The white dashed line represents the cone of influence. We only
interpreted results below this line since they are not affected by age effects. Black lines indicate a 95% confidence
ensemble of periodicity in the δ15Nsed. d) Wavelet analsis (using the same methods) at Pescadero Slope strongest
periodicities between 1200 and 1450 AD.
10 100
0
5
10
15
20
25
30
35
40
45
50
95%
50%
Period
(years)
324
93
62
10 100
0
10
20
30
40
50
60
Power
b.
95%
50%
86
102
230
65
Years (AD)
Period

1000 1200 1400 1600 1800 2000
100
200
300
400
500
Power
Period
(years)
d.
0
0.5
1
1.5
3
3.5
2
2.5
Pescadero Slope
140
Ti
(counts)
1500 2000 2500
Year AD
400
600
800
1000
1200
1400
1600
1800
Ti
K
(counts x 10
4
)
1.2 1.4 1.6
Year (AD)
400
600
800
1000
1200
1400
1600
1800
K
Ca
(counts x 10
5
)
1 1.5 2
Year (AD)
400
600
800
1000
1200
1400
1600
1800
Ca
Sr
(counts)
500 600 700 800
Year (AD)
400
600
800
1000
1200
1400
1600
1800
Sr
Ba
(counts)
300 400
Year (AD)
400
600
800
1000
1200
1400
1600
1800
Ba a. b. c. d. e.
Figure 4.5. XRF elemental analysis of SOLE-GC1. The elemental abundance is reported in terms of counts.
a) titanium (Ti), b) potassium (K), c) calcium (Ca), d) strontium (Sr), and e) barium (Ba).
141
Ti
(counts)
0 5000
Year(AD)
1000
1100
1200
1300
1400
1500
1600
Ti
K
(counts x 10
4
)
0 1 2
Year (AD)
1000
1100
1200
1300
1400
1500
1600
K
Ca
(counts x 10
4
)
1 2
Year (AD)
1000
1100
1200
1300
1400
1500
1600
Ca
Sr
(counts)
500 1000 1500
Year (AD)
1000
1100
1200
1300
1400
1500
1600
Sr
Ba
(counts)
200 400 600
Year (AD)
1000
1100
1200
1300
1400
1500
1600
Ba
a. b. c. d. e.
Figure 4.6. XRF elemental analysis of PESC-GC3. The elemental abundance is reported in terms of counts.
a) titanium (Ti), b) potassium (K), c) calcium (Ca), d) strontium (Sr), and e) barium (Ba).
142
6000
5000
4000
3000
2000
1000
0
6000 2000 4000 8000 10000 12000 14000 16000 0
Ti (counts)
K (counts)
Pescadero Slope
Soledad Basin
y = 0.1765*X - 613.1
r
2
= 0.6375
y = 0.2834*X + 1088
r
2
= 0.7768
Figure 4.7. A comparison of XRF elemental abundance of terrigenous components titanium (Ti)
and potassium (K) from Pescadero Slope (black) and Soledad Basin (gray). Linear regressions
are depicted in their respective colors. Anomalous data points with high Ti and low K from
the Pescadero Slope are circled.
143
200000
150000
100000
50000
0
0 200 400 600 800 1000 1200 1400 1600 1800
Ca (counts)
Sr (counts)
Pescadero Slope
Soledad Basin
y = 22.55*X - 10690
r
2
= 0.5807
y = 258.8*X - 4543
r
2
= 0.7015
Figure 4.8. A comparison of XRF elemental abundance of biogenic components calcium (Ca) and
strontium (Sr) from Pescadero Slope (black) and Soledad Basin (gray). Linear regressions are
depicted in their respective colors.
144
1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
8.5
9
9.5
10
10.5
Pescadero Slope ΔR 508 ± 30
1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
9.5
10
10.5
11
Soledad Basin ΔR 720 ± 68
1000 1100 1200 1300 1400 1500 1600 1700 1800 1900 2000
9
9.5
10
10.5
11
Soledad Basin ΔR 830 ± 110
δ
15
N
sed
(‰)
Year AD
Year AD
Year AD
δ
15
N
sed
(‰) δ
15
N
sed
(‰)
Figure 4.9. a) The record of OMZ intensity fluctuations from Soledad Basin (represented by δ
15
N
sed
) compared to radiocarbon age calibrated with a
ΔR of 720 ± 68 years. The dark gray color represents measurements from SOLE-GC1 while the lighter gray color are measurements from
SOLE-BC b) Soledad Basin δ15Nsed compared to radio carbon age calibrated with a ΔR of 820 ± 110 years, like part a, the dark gray color
represents measurements from SOLE-GC1 while the lighter gray color are measurements from SOLE-B. c) δ
15
N
sed
at Pescadero Slope compared to
radiocarbon age calibrated with a ΔR of 508 ± 30 years. Peaks indicating increased OMZ strength are numbered and can be correlated between
the records.
1
1
1
2
2
2
3
3
3
4
4
4
5
5
5
6
6
6
7
7
7
a.
b.
c.
145
Ti (counts)
0 5000
Year AD
1000
1100
1200
1300
1400
1500
1600
1700
Pescadero T
i
K (counts x 10
4
)
0 1 2
Year AD
1000
1100
1200
1300
1400
1500
1600
1700
Pescadero K
Ti (counts)
1500 2000 2500
Year (AD)
1000
1100
1200
1300
1400
1500
1600
1700
Soledad Ti
K (counts x 10
4
)
1.2 1.4 1.6
Year AD
1000
1100
1200
1300
1400
1500
1600
1700
Soledad K
Figure 4.10. Variations in terrigenously sourced elements from 1700 to 950 AD. a. Ti counts compared to the original radiocar-
bon age models of the Soledad and Pescadero cores. b. K coutns compared to the original radiocarbon age models on the Sole-
dad and Pescadero cores. Both data sets indicated an average offset of 113 years between sites.
a. b.
146
Correlation Peaks Pescadero Slope Age (AD) Soledad Basin Age (AD) Age Offset (years)
1 1583 1480 103
2 1515 1384 131
3 1211 1107 104
4 1349 1237 112
5 1149 1036 113
Average Offset (years) 112.6
Standard Deviation Offest (years) 11.2
XRF Analysis Correlation
Table 4.2 The values of correlation peaks between the Pescadero Slope and Soledad Basin based on the original age models. There is
an average offsed of 112.6 years with a standard deviation of 11.2 years.
147

Chapter 5
Summary and Conclusions
The work presented in this dissertation has examined the geochemistry of marine
sediments to investigate variations in oxygen minimum zone (OMZ) intensity and changes in
nitrogen cycling in the eastern tropical north Pacific (ETNP). The δ
15
N

of bulk laminated
sediments from the Santa Monica Basin, Soledad Basin and Pescadero Slope have been
evaluated to identify changes unique to each site and investigate similar trends in OMZ
fluctuations across the northern portion of the ETNP OMZ. The first portion of this study
focused on investigating the advection of the OMZ signal from the Mexican waters into the
California Borderland through the California Undercurrent (Chapter 2). The second portion,
centered on developing high-resolution (near annual) records of fluctuations in OMZ intensity
during the late Holocene and investigates the factors influencing variation in the strength of the
OMZ and frequency patterns of fluctuation (Chapter 3). The final portion of this dissertation
examined regional variations in the intensity of the ETNP OMZ, by comparing proxy signals
between the Pescadero Slope and the Soledad Basin, which are both on the periphery of the
ETNP OMZ but separated by ~470 km (Chapter 4).
Studies examining dissolved oxygen (DO) concentration in the ocean have revealed
significant declines in DO and the expansion of OMZs globally in the past 30 years [Stramma et
al., 2008]. This work prompted the investigation of how the intensity of the ETNP OMZ has
varied over the past 150 years at the Pescadero Slope, at the mouth of the Gulf of California, and
further northward in the Santa Monica Basin in the Southern California Borderland. Using δ
15
N

of bulk sedimentary matter from laminated sediments, which has been used to reconstruct large-
scale glacial-interglacial changes in denitrification [Ganeshram et al., 1995, 2000], high-
148

resolution (near annual) records of denitrification/OMZ intensity were reconstructed at both
sites. The construction of a
210
Pb age model through evaluations of integrated mass and
210
Pb
activity constrained the age of the sediments downcore. This revealed that the recent 30-year
intensification of the OMZ (corroborating the observations made from O
2
measurements in the
water column) was preceded by a gradual contraction of the OMZ from the 1850 to 1990. This
latter finding was novel and unexpected. A paper that describes modeling of OMZ contraction
and expansion consistent with the sediment records was published by Deutsch et al. [2014]. I am
an contributing author to this study. This paper appears in Appendix A and served as a launching
point for the work I present in my dissertation.
Denitrification in the OMZ (C
org
respiration with nitrate as electron acceptor) is a process
that increases the isotopic value of the residual nitrate. Significant active denitrification is not
occurring below the thermocline in the waters off southern California, indicating that the
elevated δ
15
N
sed
recorded in the sediments there must be the product of previous denitrification in
the ETNP and its advection northward into the California Borderland. Hydrographic studies
indicate water with low N* (indicating denitrified waters) and elevated
15
NO
3
-
are transported
from the eastern tropical north Pacific northward into the Southern California Bight by the
California Undercurrent [Liu and Kaplan, 1989; Castro et al., 2001] between 150 and 250 m in
the water column [Hickey, 1992]. As water from these depths upwells, the
15
N signature is
preserved in the laminated sediment in the Santa Monica Basin. Conversely, the California
Current transports water with characteristically high N* (indicating minimal denitrification) with
reduced
15
NO
3
-
from the north into the Southern California Bight [Castro et al., 2001]. These
water masses mix in the Southern California Bight and the composite is recorded in the δ
15
N
sediment record. This premise was verified in this study through investigations between
149

salinity, a common ocean chemistry tracer, in the core of the California Undercurrent (150-250
m) in Santa Monica water column and the δ
15
N
sed
from Santa Monica Basin. We introduce
Δδ
15
N
sed
, a parameter defined as the difference in δ
15
N
sed
values between the Pescadero Slope
and Santa Monica Basin, to account for variation in
15
NO
3
-
in the southern end member.
Significant positive correlations between salinity and δ
15
N
sed
in Santa Monica Basin between
1952 and 1990 suggest that elevated δ
15
N
sed
corresponds to a stronger California Undercurrent.
The comparison of Δδ
15
N
sed
versus salinity showed a stronger, significant correlation between
1952 and 1990, suggesting that down core analysis of the Δδ
15
N
sed
between the two sites
provides a robust metric to investigate variations in the strength of the California Undercurrent.
Trends in Δδ
15
N
sed
through the 20
th
century suggest a relationship between the strength of
the California Undercurrent and the Pacific Decadal Oscillation. From 1900 to 1993, the PDO
index and Δδ
15
N
sed
are anticorrelated, suggesting that when the California Current is stronger (as
indicated by negative PDO Index values) the California Undercurrent is weaker as evidenced by
an increased offset between the δ
15
N
sed
values at the Pescadero Slope and the Santa Monica
Basin. This anticorrelation breaks down, however, and the relationship is reversed after 1993
when significant intensification of the OMZ and a convergence of Pescadero Slope and Santa
Monica Basin δ
15
N
sed
begin.
The change in the relationship between PDO and Δδ
15
N
sed
is synchronous with the
disintegration of trends between δ
15
N
sed
/Δδ
15
N
sed
and salinity, which is hypothesized to be
caused by a change in water chemistry of the California Current. Modeling studies suggest that
the source of water to the California Current has changed and water parcels with higher nitrate
concentrations sourced from a more northerly location are being advected into the California
Borderland [Rykaczewski and Dunne, 2010]. Northern Pacific water generally exhibits elevated
150

15
NO
3
-
[Casciotti et al., 2008]. The combination of increased nitrate concentrations with
potentially elevated
15
NO
3
-
would significantly disrupt the mixing relationship between the
California Undercurrent and the California Current.
The second portion of this dissertation focused on developing a high-resolution record of
OMZ fluctuations over the past 1200 years to evaluate if the trends observed in the past 150
years in at the Pescadero Slope and the Santa Monica Basin are anomalous. This record was
reconstructed by measuring δ
15
N
sed
in the laminated sediment core and determining the age of
the sediment through radiocarbon measurements corrected for reservoir effects and varve counts.
This record revealed that the expansion of the OMZ seen in both δ
15
N
sed
records and O
2

measurements in the last 30 years is not unprecedented in the late Holocene. This 1200-year
record indicates that the OMZ at Pescadero Slope is characterized by rapid intensifications that
are followed by more gradual relaxations within the bounds of maximum and minimum δ
15
N
sed
values. Trends in OMZ intensity changes were assessed quantitatively by comparing the slopes
between low to high δ
15
N
sed
values (OMZ expansion) and high to low δ
15
N
sed
(OMZ contraction)
for ten intervals in which each interval was determined by exhibiting an OMZ expansion and
contraction over a minimum of 40 years. This analysis showed that OMZ intensification occurs
twice as fast as OMZ re-oxygenation.
In order to evaluate how these variations in nitrogen isotopes reflect changes in O
2

content, the relationship between δ
15
NO
3
-
and O
2
in the water column was evaluated from prior
studies at the Pescadero Basin and nearby Soledad and Magdalena Basins. This indicated that a
20 ± 3.1µM fluctuation in O
2
would be represented as a 1‰ change in δ
15
NO
3
-
, and thus δ
15
N
sed
.
Based on variation in δ
15
N
sed
measured at the Pescadero Slope, fluctuations of 45 µM O
2
can
occur in as quickly as 40 years which is similar to the rate of change today.
151

OMZ fluctuations at the Pescadero Slope are capped by minimum (8.0‰) and maximum
(10.5‰) δ
15
N
sed
values, with variations between these values. This shows that the OMZ is
dynamic and suggests that the OMZ could be self-regulating within these limits. This study used
weight percent organic carbon (wt. % C
org
) as a proxy for organic carbon delivery to the
sediments. A significant correlation between wt. % C
org
and δ
15
N
sed
suggests that increased
carbon export from the photic zone is related to elevated δ
15
N
sed
values, suggesting that changes
in carbon export could be one factor involved in OMZ self-regulation. Diatoms and radiolarians,
which make their skeletons from silica, are abundant in the Gulf of California, and thus it was
hypothesized that wt. % bSi, like wt. % C
org
, would have a similar relationship to δ
15
N
sed.
This
was not found to be the case; a significant relationship between wt. % bSi and δ
15
N
sed
was not
found. We hypothesize that this is due to significant remineralization of biogenic silica in the
upper water column as has been observed by other studies [Thunell et al., 1994; Pride et al.,
1999].

This suggests that wt. % C
org
in the sediments is a more reliable proxy for assessing past
changes in primary productivity at the Pescadero Slope.
To investigate what factors could be driving changes in carbon export and OMZ intensity
spectral analysis using a Lomb-Scargle Fourier Transform [Lomb, 1976; Scargle, 1982, 1989] in
combination with a Welch-Overlapped-Segment-Averaging (WOSA) for unevenly spaced data
[Schulz and Stattegger, 1997; Welch, 1967; Khider et al., 2014] was preformed. This revealed
periodicities at 65, 86, 102, and 230 years above 95% confidence. These periodicities may relate
to the PDO (65 years), the Gleissberg Cycle (86 and 102 years), and Suess (DeVries) Cycle (230
years). The Gleissberg Cycle and the Suess Cycle are solar influenced, related to the 11-year
sunspot cycle suggesting variation in solar irradiance may play a role in OMZ intensity
fluctuations and regulation. Solar irradiance and PDO phase change variations were also
152

compared to the OMZ intensity records qualitatively, exhibiting a some correspondence in
trends. Wavelet analysis of these periodicities indicate that the strongest periodicities occur
between 1200 and 1450 AD in this record.
The final section of this work aimed to expand the understanding of OMZ fluctuations
beyond local scales in the Gulf of California and investigate if variations in OMZ intensity are
synchronous in the ETNP region. This study is a proof of concept insofar as both Chapter 2 and
Deutsch et al. [2014] (Appendix A) assume that synchronous changes occur throughout this
region. This study investigates, in detail, if these are valid assumptions. This chapter focuses on
evaluating assumptions of traditional radiocarbon age model construction. This study found
similar peaks in OMZ expansion in both records suggesting that regional events overprint local
variations, however, these peaks appeared to be offset in time with OMZ intensification first
occurring in the Soledad Basin. It is hypothesized that this variation in timing is a product of age
model construction rather than these events occurring at different times in each location. Due to
the extensive work to verify the Pescadero Slope radiocarbon age model and its consistency with
the
210
Pb age model, for the purpose of this study we assume the Pescadero core represents the
‘correct’ age and evaluate potential variation in Soledad Basin radiocarbon age model, although
it is recognized that the Pescadero Slope age model also contain a degree of inherent uncertainty.
Offset between the records was constrained independently from the δ
15
N
sed
records by using
XRF analysis of titanium and potassium elemental abundance. Titanium and potassium, both
elements of primarily terrigenous origin, should show regional coherence in large-scale events,
which would lead to variations in deposition in the marine realm. Correlation points in both the
titanium and potassium records in both the Pescadero Slope and Soledad Basins suggest an
average offset of 110 years between the two records. Uncertainty in radiocarbon age models is
153

largely due to the estimated local reservoir effect, ΔR, which is the offset between a specific
location reservoir age and the global reservoir age. The recalibrated radiocarbon age model for
Soledad Basin sediments showed greater synchronicity to the Pescadero Slope δ
15
N
sed
records,
illustrating the significant impact that estimated local reservoir ages can have on
paleoceanographic and paleoenvironmental reconstructions.
Spectral analysis was also used to investigate what factors were driving variations in the
intensity of the OMZ at Soledad Basin and if the overall periodicities were similar to those found
in the Pescadero Slope δ
15
N
sed
record. Spectral analysis of the δ
15
N
sed
record revealed
periodicities at 40, 62, 93, and 194 year periodicities (40 and 194 year with greater than 95%
confidence; 62 and 93 year with greater than 50% confidence). The 40 and 62-year periodicities
are likely related to the Pacific Decadal Oscillation as with the 65-year periodicity found in the
Pescadero Slope OMZ intensity record. The 93-year periodicity could be related to the
Gleissberg cycle, similar to the 86 and 102 year periodicities at Pescadero Slope. The similar
periodicities also suggest regional variations in ETNP OMZ intensity.
XRF analysis was also used to evaluate the local geochemical signatures of the sediments
at both Pescadero Slope and Soledad Basin. Analysis of terrigenously sourced titanium and
potassium revealed more variability and higher abundances of both elements, likely due to
increased detrital input at the Pescadero Slope due to the proximity of the location to the volcanic
Sierra Madre Occidental on the Mexican mainland. Soledad Basin receives much less detrital
input due to its location near Baja California, a much smaller landmass with less relief. The
primarily biogenic elements, calcium and strontium, also show significant geochemical
variations between the two sites. The higher calcium abundances at Soledad Basin are the
product of coccolithophore blooms that occur in the summer months at this location, which do
154

not occur at Pescadero Slope. Strontium, which is typically associated with calcium carbonate,
is significantly elevated at Pescadero Slope compared to Soledad Basin. We propose that
strontium in the Gulf of California is delivered by an alternative source, potentially related to
hydrothermal activity in the Gulf.
The high-resolution δ
15
N
sed
records from Pescadero Slope, Santa Monica Basin, and
Soledad Basin, provide insight into the dynamics of fluctuations in the intensity of the ETNP
OMZ over the past 1200 years (150 years in the case of Santa Monica Basin). The near annual
records, with carefully evaluated age constraints, greatly expand the scientific communities
paleoceanographic archives. These data sets can be used to quantitatively test the mechanisms
linking climate and anoxia in General Climate Models (GCMs), enabling further investigation of
the mechanisms influencing OMZ expansion and contraction and changes in ocean circulation,
while also providing a means to constrain predictions of future OMZ fluctuations, which will
provide valuable information for ocean management and climate policies.
In a study conducted in collaboration with Dr. Curtis Deutsch at the University of
Washington, we used the of the δ
15
N
sed
multicore records from Pescadero Slope, Soledad Basin,
and Santa Monica Basin to test hypothesized mechanisms underlying recent changes in the OMZ
[Deutsch et al., 2014]. The use of a GCM that includes an explicit nitrogen and nitrogen isotope
cycle predicts decadal differences in the volume of anoxic water in the ETNP are driven by
respiration [Deutsch et al., 2001; Ito and Deutsch, 2013], which varies in relation to the depth of
the thermocline. Reconstructed variations in thermocline depth (along the 13°C isotherm)
strongly correlate to the δ
15
N
sed
data, insofar as the slopes between the data and model
predictions are indistinguishable. In the tropical Pacific, respiration and consequently the depth
of the thermocline is closely linked to atmospheric Walker circulation, suggesting a long-term
155

weakening of the equatorial trade winds led to the contraction of OMZ in the ETNP [Deutsch et
al., 2014] (Appendix A). This study only brushes the surface of potential investigations that can
utilize the late Holocene δ
15
N
sed
records to quantitative assess the validity of GCMs in the
reconstruction of past oceanographic conditions.
The work in this dissertation has expanded the current understanding of the dynamics of
the ETNP OMZ over the past 1200 years, suggesting that the ETNP OMZ could be self-
regulating in part governed by variations in carbon export and productivity which is influenced
by changes in thermocline depth, PDO, solar irradiance, the Gleissberg and Suess cycles, and the
intensity of the easterly trade winds. While this body of work expands scientific knowledge
about the eastern tropical north Pacific, it also promotes more questions about these critical
ecosystems and provides insight into the direction of future studies. For instance, is the ETNP
OMZ self-regulation solely related to carbon export or are there other mechanisms involved that
keep the amount of nitrogen loss from the OMZ in check? With unprecedented concentrations of
greenhouse gases in the atmosphere and a rapidly warming climate will the mechanisms
regulating the size of the OMZ be significantly altered influencing regional and global climate
further? Can incorporation of δ
15
N
sed
data into general circulation models provide insight into
other potential regulation mechanisms and predict future variations? Along similar lines, would
O
2
variations on the order of 45 µM (the maximum calculated O
2
content fluctuation)
significantly impact the amount of organic carbon preserved in the sediments beneath the OMZ?
Have changes in California Current water chemistry, as suggested by the evaluations of δ
15
N
sed
at
the Pescadero Slope and Santa Monica Basin and salinity data, significantly impacted
productivity and biogeochemical cycling along the western coast of North America in the past or
are the changes occurring today an anomalous product of the warming climate? This work
156

additionally highlights the importance of developing methods to better constrain local reservoir
ages and potential variation in local reservoir ages quantitatively. It also reinforces the utility of
developing paleoceanographic records from multiple sites to disentangle regional and local
signals and the origin of the signals.
A starting point for future studies that encompasses many of the themes discussed above
involves developing reconstructions of the ETNP OMZ signal of similar lengths at sites in which
elevated
15
NO
3
-
has been advected northward by the California Undercurrent. An ideal study site
for this is Santa Barbara Basin in the Southern California Borderland. The Santa Barbara Basin
is characterized by high rates of primary productivity, rapid accumulation of marine organic
matter, and a regional OMZ, which result in the deposition of varved sediments that serve as
high resolution paleoclimate and paleoceanographic archives [Barron et al., 2015; Emmer and
Thunell, 2000; Behl and Kennett, 1996]. An 18-year δ
15
N record of particulate matter in
sediment traps from Santa Barbara Basin [Deutsch et al., 2014] exhibits similar trends to those
observed in the Pescadero Slope, Soledad Basin, and Santa Monica Basin records. This suggests
that developing a high-resolution 1200-year record of OMZ intensity will provide a significant
comparison to the Pescadero Slope and Soledad Basin records, expanding the evaluation of
regional ETNP OMZ dynamics and providing a means to assess late Holocene California
Undercurrent dynamics and investigate potential variations in California Current water chemistry
during the late Holocene.
To further understand the dynamics driving variation in OMZ strength in the ETNP it is
critical to evaluate how the OMZ in the eastern tropical south Pacific (ETSP) has varied over
similar time frames and at similar resolution. For instance, if variation of OMZ intensity is
related to the intensity of the eastern trade winds [Deutsch et al., 2014] similar trends in δ
15
N
sed

157

of bulk laminated sediments should be found in ETSP records. Collecting cores from the ETSP
at a location that has complete nitrate utilization in the photic zone, high sediment accumulation
and preservation rates, and no diagenetic alteration and measuring δ
15
N
sed
at near annual
resolution would provide insight into the dynamics regulating the OMZ in both regions of the
Pacific.

158

References
Barron, J.A., D. Bukry, and I.L. Hendy (2015), High-resolution paleoclimatology of the Santa
Barbara Basin during the Medieval Climate Anomaly and early Little Ice Age based on
diatom and silicoflagellate assemblages in Kasten core SPR0901-02KC, Quaternary
International, 387, 13-22.
Behl, R. J., and J. P. Kennett (1996), Brief interstadial events in the Santa Barbara basin, NE
Pacific, during the past 60 kyr, Nature, 379(6562), 243–246, doi:10.1038/379243a0.
Butzin, M., M. Prange and G. Lohmann (2012), Readjustment of glacial radiocarbon
chronologies by self-consistent three-dimensional ocean circulation modeling, Earth
Planetary Science Letters, 317-318, 177–184.
Casciotti, K. L., T. W. Trull, D. M. Glover, and D. Davies (2008), Constraints on nitrogen
cycling at the subtropical North Pacific Station ALOHA from isotopic measurements of
nitrate and particulate nitrogen, Deep Sea Res. Part II, 55, 1661–1672,
doi:10.1016/j.dsr2.2008.04.017.
Castro, C. G., F. P. Chavez, and C. A. Collins (2001), Role of the California Undercurrent in the
export of denitrified waters from the eastern tropical North Pacific, Global Biogeochem.
Cycles, 15(4), 819–830, doi:10.1029/2000GB001324.
De La Fuente, M., L. Skinner, E. Valvo, C. Pelejero, and I. Cacho (2015), Increased reservoir
ages and poorly ventilated deep waters inferred in the glacial eastern equatorial Pacific,
Nature Communications, 6, 7420, doi:10.1038/ncomms8420.
Deutsch, C., H. Brix, T. Ito, H. Frenzel, and L. Thompson (2011), Climate-forced variability of
ocean hypoxia, Science, 333(6040), 336–339, doi:10.1126/science.1202422.
159

Deutsch, C., W. Berelson, R. Thunell, T. Webber, C. Tems, J. McManus, J. Crusius, T. Ito, T.
Baumgartner, V. Ferreira, J. Mey, and A. VanGeen (2014), Centennial changes in North
Pacific anoxia linked to tropical trade winds, Science, 345 (6197), 665-668,
doi:10.1126/science.1252332.
Emile-Geay, J., K.M. Cobb, M.E. Mann, and A.T. Wittenberg (2013), Estimating central
equatorial Pacific SST variability over the past millennium. Part II: reconstructions and
implications, Journal of Climate, 26(7), 2329-2352, doi:10.1175/JCLI-D-11-00511.1.
Emmer, E. and R.C. Thunell (2000), Nitrogen isotope variations in Santa Barbara Basin
sediments: implications for denitrification in the eastern tropical north Pacific during the last
50,000 years, Paleoceanography, 15(4), 377-387.
Hickey, B. M. (1992), Circulation over the Santa Monica-San Pedro Basin and shelf, Progress in
Oceanography, 30, 37–115, doi:10.1016/0079-6611(92) 90009-O.
Ganeshram, R.S., T.F. Pedersen, S.E. Calvert, and J.W. Murray (1995), Large changes in oceanic
nutrient inventories from glacial to interglacial periods, Nature, 376, 755-758.
Ganeshram, R. S., T. F. Pedersen, S. E. Calvert, G. W. McNeil, and M. R. Fontugne (2000),
Glacial-interglacial variability in denitrification in the world’s oceans: Causes and
consequences, Paleoceanography, 15(4), 361–376, doi:10.1029/1999PA000422.
Ito, T., and C. Deutsch (2013), Variability of the oxygen minimum zone in the tropical North
Pacific during the late twentieth century, Global Biogeochemical Cycles, 27(4), 1119-1128.
Khider, D., C. S. Jackson, and L. D. Stott (2014), Assessing millennial-scale variability during
the Holocene: A perspective from the western tropical Pacific, Paleoceanography, 29, 143–
159, doi:10.1002/2013PA002534.
160

Liu, K.-K., and I. F. Kaplan (1989), The Eastern Tropical Pacific as a source of 15 N-enriched
nitrate in seawater off Southern California, Limnology and Oceanography, 34(5), 820–830.
MacDonald, G.M. and R.A. Case (2005), Variations in the Pacific Decadal Oscillation over the
past millennium, Geophysical Research Letters, 32: L08703, doi:10.1029/2005GL022478.
Pride, C., R. Thunell, D. Sigman, L. Keigwin, M. Altabet, and E. Tappa (1999), Nitrogen
isotopic variations in the Gulf of California since the last deglaciation: Response to global
climate change, Paleoceanography, 14(3), 397–409, doi:10.1029/1999PA900004.
Rykaczewski, R. R., and J. P. Dunne (2010), Enhanced nutrient supply to the California Current
Ecosystem with global warming and increased stratification in an Earth system model,
Geophysical Research Letters, 37, L21606, doi:10.1029/2010GL045019.
Scargle, J. D. (1982), Studies in astronomical time series analysis. II. Statistical aspects of
spectral analysis of unevenly spaced data, Astrophys J., 263(2), 835–853.
Scargle, J. D. (1989), Studies in astronomical time series analysis. III. Fourier transforms,
autocorrelation functions, and cross-correlation functions of unevenly spaced data,
Astrophys. J., 343(2), 874–887.
Schulz, M., and K. Stattegger (1997), SPECTRUM: Spectral analysis of unevenly spaced time
series, Comput. Geosci., 23(9), 929–945.
Steinhilber, F., J. Beer, and C. Fröhlich (2009), Total solar irradiance during the Holocene,
Geophysical Research Letters, 36, L19704, doi: 10.1029/2009GL040142.
Stramma, L., GC. Johnson, J. Sprintall, and V. Mohrholz (2008), Expanding oxygen minimum
zones in the tropical oceans, Science, 320: 655-658.
161

Thunell, R.C., C.J. Pride, E. Tappa, and F.E. Muller-Karger (1994), Biogenic silica fluxes and
accumulation rates in the Gulf of California, Geology, 22, 303-306, doi: 10.1130/0091-
7613(1994)022<0303:BSFAAR>2.3.CO;2.
Welch, P. D. (1967), The use of Fast Fourier transform for the estimation of power spectra: A
method based on time averaging over short, modified periodograms, IEEE Trans. Audio
Electroacoust., 15(2), 70–73.

162
Bibliography
Altabet, M. A., and R. Francois (1994), Sedimentary nitrogen isotopic ratio as a recorder for
surface ocean nitrate utilization, Global Biogeochemical Cycles, 8(1), 103–116,
doi:10.1029/93GB03396.
Altabet, M.A., R. Francois, D.W. Murray, and W.L. Prell (1995), Climate-related variations in
denitrification in the Arabian Sea from sediment
15
N/
14
N ratios, Nature, 373(6514), 506-509,
doi:10.1038/373506a0.
Altabet, M. A., C. Pilskaln, R. Thunell, C. Pride, D. Sigman, F. Chavez, and R. Francois (1999),
The nitrogen isotope biogeochemistry of sinking particles from the margin of the eastern
North Pacific, Deep Sea Research Part I, 46, 655–679, doi:10.1016/S0967-0637(98)00084-3.
Anderson, .H., C. Woulds, M. Schwartz, G.L. Cowie, L.A. Levin, K. Soetaert, J.J. Middelburg
(2008), Short-term fate of phytodetrtus in sediments across the Arabian Sea Oxygen
Minimum Zone, Biogeosciences, 5, 43-53.
Baba, J., C.D. Peterson, H.J. Schrader (1991), Modern fine-grained sediments in the Gulf of
California, In: The Gulf and Peninsular Province of the Californias (eds. Dauphin, J.P., and
B.R.T. Simoneit), American Association of Petroleum Geologists Memoir, 47, 569–587.
Babbin, A.R., R.G. Keil, A.H. Devol, and B.B. Ward (2014), Organic matter stoichiometry, flux,
and oxygen control nitrogen loss in the ocean, Science, 344, 406-408,
doi:10.1126/science.1248364.
Bahr, A., F. Lamy, H. Arz, H. Kuhlmann, and G. Wefer (2005), Late glacial to Holocene climate
and sedimentation history in the NW Black Sea, Marine Geology, 214(4), 309-322.
163
Barnett, P.R.O., J. Watson, and D. Connelly (1984), A multiple corer for taking virtually
undisturbed samples from shelf, bathyal and abyssal sediments, Oceanologica Acta,
7(4), 399-408.
Barron, J.A., D. Bukry, and I.L. Hendy (2015), High-resolution paleoclimatology of the Santa
Barbara Basin during the Medieval Climate Anomaly and early Little Ice Age based on
diatom and silicoflagellate assemblages in Kasten core SPR0901-02KC, Quaternary
International, 387, 13-22.
Barron, J.A., and D. Bukry (2007), Solar forcing of Gulf of California climate during the past
2000 yr suggested by diatoms and silicoflagellates, Marine Micropaleontology, 62, 115-139,
doi:10.1016/j.marmicro.2006.08.003.
Baumgartner, T. R., A. Soutar, and V. Ferreira-Bartrina (1992), Re-construction of the history of
Pacific sardine and northern anchovy populations over the past two millennia from sediments
of the Santa Barbara basin, California, CalCOFI Reports 33, 24–40.
Baumgartner,T., V. Ferreira, J. Mey, and A. VanGeen (2014), Centennial changes in North
Pacific anoxia linked to tropical trade winds, Science, 345 (6197), 665-668,
doi:10.1126/science.1252332.
Bauska, T.K., F. Joos, A.C. Mix, R. Roth, J. Ahn, and E.J. Brook (2015), Links between
atmospheric carbon dioxide, the land carbon reservoir and climate over the past millennium,
Nature Geoscience, 8, 383-387, doi:10.1038/NGEO2422.
Beers, J. R., F. M. H. Reid and G. L. Stewart (1975), Microplankton of the north Pacific central
gyre: population structure and abundance, Internationale Revue Gesamtes Hydrobiologia 60,
607–638.
164
Behl, R.J. and J.P. Kennett (1996), Brief interstadial events in the Santa Barbara basin, NE
Pacific, during the past 60 kyr, Nature, 379(6562), 243-246.
Bender, M., R. Jahnke, F. Weiss, W. Martin, D.T. Heggie, J. Orchardo, T. Sowers (1989),
Organic-carbon oxidation and benthic nitrogen and silica dynamics in San-Clemente Basin, a
Continental Borderland site, Geochimica et Cosmochimica Acta, 53(3), 685-697.
Benton, M.J. and R.J. Twitchett (2003), How to kill (almost all life: the end-Permian extinction
event, Trends in Ecology and Evolution, 18(7), 358-365, doi:10.1016/S0169-5347(03)00093-
4.
Berelson, W.M., M. Prokopenko, F.J. Sansone, A.W. Graham, J. McManus, J.M. Bernhard
(2005), Anaerobic diagenesis of silica and carbon in continental margin sediments: discrete
zones of TCO2 production, Geochimica et Cosmochimica Acta, 69(19), 4611-4629.
Bindoff, N. L. and T. J. McDougall (2000), Decadal changes along an Indian Ocean section at
32 degrees S and their interpretation, Journal of Physical Oceanography, 30, 1207– 1222.
Bloemendal, J., B. Lamb, and J. King (1988), Paleoenvironmental implications of rock-magnetic
properties of late Quaternary sediment cores from the eastern equatorial Atlantic.
Paleoceanography, 3, 61-87.
Bograd, S.J., C.G. Castro, E. Di Lorenzo, D.M. Palacios, H. Hailey, W. Gilly, and F.P. Chavez
(2008), Oxygen declines and the shoaling of the hypoxic boundary in the California Current,
Geophysical Research Letters, 35, L12607, doi:10.1029/2008GL034185.
Booth, J.A.T., C.B. Woodson, M. Sutula, F. Micheli, S.B. Weisberg, S.J. Bograd, A. Steele, J.
Schoen,7 and L. B. Crowder (2014), Patterns and potential drivers of declining oxygen
content along the southern California coast, Limnology Oceanography, 59(4), 1127-1138,
doi: 10.4319/lo.2014.59.4.1127.
165
Brandes, J. A., and A. H. Devol (2002), A global marine-fixed nitrogen isotopic budget:
Implications for Holocene nitrogen cycling, Global Biogeochemical Cycles, 16(4), 1120,
doi:10.1029/2001GB001856.
Butzin, M., M. Prange and G. Lohmann (2012), Readjustment of glacial radiocarbon
chronologies by self-consistent three-dimensional ocean circulation modeling, Earth
Planetary Science Letters, 317-318, 177–184.
Capone, D. G. (2001), Marine nitrogen fixation: what's the fuss? Current opinion in
microbiology, 4(3), 341-348.
Casciotti, K.L, T.W. Trull, D.M. Glover, and D. Davies (2008), Constraints on nitrogen cycling
at the subtropical North Pacific Station ALOHA from isotopic measurements of nitrate and
particulate nitrogen, Deep-Sea Research II, 55, 1661-1672, doi: 10.1016/j.dsr2.2008.04.017.
Casciotti, K. L., C. Buchwald, and M. McIlvin (2013), Implications of nitrate and nitrite isotopic
measurements for the mechanisms of nitrogen cycling in the Peru oxygen deficient zone,
Deep-Sea Research I, 80: 78-98.
Casciotti, K. L. (2016), Nitrogen and oxygen isotopic studies of the marine nitrogen cycle,
Annual Reviews of Marine Science, 8, 379-407, doi: 10.1146/annurev-marine-010213-
135052.
Castro, C. G., F. P. Chavez, and C. A. Collins (2001), Role of the California Undercurrent in the
export of denitrified waters from the eastern tropical North Pacific, Global Biogeochemical
Cycles, 15(4), 819–830, doi:10.1029/2000GB001324.
Chamchoi, N., S. Nitisoravut, and J.E. Schmidt (2008), Inactivation of ANAMMOX
communities under concurrent operation of anaerobic ammoium oxidation (ANAMMOX)
and denitrification, Bioresource Technology, 9 (99), 3331-6.
166
Chan, F., J.A. Barth, J. Lubchenco, A. Kirincic, H. Weeks, W.T. Peterson, and B.A. Menge
(2008), Emergence of anoxia in the California Current large marine ecosystem, Science, 319,
920.
Chavez, F.P., M. Messié, and J.T. Pennington (2011), Marine Primary Production in Relation to
Climate Variability and Change, Annu. Rev. Mar. Sci, 3, 227-260, doi:
10.1146/annurev.marine.010908.163917.
Chong, L. S., M. G. Prokopenko, W. M. Berelson, A. Townsend-Small, and J. McManus (2012),
Nitrogen cycling within suboxic and anoxic sediments from the continental margin of
western North America, Marine Chemistry, 128–129, 13–25,
doi:10.1016/j.marchem.2011.10.007.
Chong-jian, T., P. Zheng, C-h Wang, and Q. Mahmood (2010), Suppression of anaerobic
ammonium oxidizers under high organic content in high-rate Anammox UASB reactor,
Bioresource Technology, 101(6), 1762-1768.
Christensen, C.J., D.S. Gorsline, D.E. Hammond, S.P. Lund (1994), Non-annual laminations and
expansion of anoxic basin-floor conditions in Santa Monica Basin, California Borderland,
over the past four centuries, Marine Geology, 116, 399-418.
Cline, J.D. and F.A. Richards (1972), Low oxygen concentrations and nitrate reduction in the
eastern tropical North Pacific ocean, Limnology and Oceanography, 17, 885-900.
Codispoti, L.A. (1973), Denitrification in the Eastern Tropical North Pacific ocean, Ph.D.
Dissertation, University of Washington, Seattle.
Collins, L.E., W. Berelson, D.E. Hammond, A. Knapp, R. Schwartz, D. Capone (2011), Particle
fluxes in San Pedro Basin, California: A four-year record of sedimentation and physical
forcing, Deep-Sea Research I, 58(8), 898-914, doi:10.1016/j.dsr.2011.06.008.
167
Curray, J.R., D.G. Moore, J.E. Aguayo, M.P. Aubry, G. Einsele, D.J. Fornari, J. Gieskes, J.C
Guerrero, M. Kastner, K. Kelts, M. Lyle, Y. Matoba, A. Molina-Cruz, J. Niemitz, J. Rueda,
A.D. Sauders, J. Schrader, B.R.T. Simoneit, and V. Vacquier (1979), Leg 64 seeds evidence
on development of basins, Geotimes 24, 18–20.
Dean, W., C. Pride, and R. Thunell (2004), Geochemical cycles in sediments deposited on the
slopes of the Guaymas and Carmen Basins of the Gulf of California over the last 180 years,
Quaternary Science Reviews, 23(16), 1817-1833.
Deckker, P.D. (2004), On the celestite-secreting Acantharia and their effect on seawater
strontium to calcium rations, Hydrobiologia, 517, 1-13.
De La Fuente, M., L. Skinner, E. Valvo, C. Pelejero, and I. Cacho (2015), Increased reservoir
ages and poorly ventilated deep waters inferred in the glacial eastern equatorial Pacific,
Nature Communications, 6, 7420, doi:10.1038/ncomms8420.
DeMaster, D.J. (1979), The Marine Budgets of Silica and
32
Si, Yale University, New Haven.
DeMaster, D.J. (1991), Measuring biogenic silica in sediments and suspended matter, American
Geophysical Union.
De Pol-Holz, R. O. Ulloa, L. Dezilau, J. Kaiser, F. Lamy, D. Hebbeln (2006), Melting of the
Patagonian Ice Sheet and deglacial perturbations of the nitrogen cycle in the eastern South
Pacific, Geophysical Research Letters, 33(4), L04704, doi:10.1029/2005GL024477.
Deutsch, C., S. Emerson, L. Thompson (2005), Fingerprints of climate change in North Pacific
oxygen, Geophysical Research Letters, 32, L16604, doi:10.1029/2005GL023190.
Deutsch, C., H. Brix, T. Ito, H. Frenzel, and L. Thompson (2011), Climate-Forced Variability of
Ocean Hypoxia, Science, 333(6040), 336-339, doi:10.1126/science.1202422.
168
Deutsch, C. W. Berelson, R. Thunell, T. Webber, C. Tems, J. McManus, J. Crusius, T. Ito, T.
Baumgartner, V. Ferreira, J. Mey, and A. VanGeen (2014), Centennial changes in North
Pacific anoxia linked to tropical trade winds, Science, 345 (6197), 665-668,
doi:10.1126/science.1252332.
DeVries, T., C. Deutsch, F. Primeau, B. Chang, and A. Devol (2012), Global rates of water-
column denitrification derived from nitrogen gas measurements, Nature Geoscience, 5(8),
547–550, doi:10.1038/ngeo1515.
DeVries, T., C. Deutsch, P. A. Rafter, and F. Primeau (2013), Marine denitrification rates
determined from a global 3-D inverse model, Biogeosciences, 10(4).
Devol, A.H. (2003), Solution to a marine mystery, Nature, 422, 575-576.
Devol, A.H. (2008), Denitrification including anammox, in Nitrogen in the Marine Environment
(ed. Capone, D.), 263-301, doi: 10.1016/B978-0-12-372522-6.00006-2.
Diaz, R.J., R. Rosenberg (2008), Spreading dead zones and consequences for marine
ecosystems, Science, 321, 926-929.
Di Lorenzo, E. (2003), Seasonal dynamics of the surface circulation in the Southern California
Current Systems, Deep-Sea Research II, 50, 2371-2388, doi:10.1016/S0967-0645(03)00125-
5.
Douglas, R. O. Gonzalez-Yajimovich, J. Ledesma-Vazquez, and F. Staines-Urias (2007),
Climate forcing, primary production and the distribution of Holocene biogenic sediments in
the Gulf of California, Quaternary Science Reviews, 26, 115-129, doi:
10.1016/j.quascirev.2006.05.003.
169
Emerson, S.R., R. Jahnke, M. Bender, P. Froelich, G. Klinkhammer, C. Bowser, F. Setlock
(1980), Early diagenesis in sediments from the eastern equatorial Pacific. I. Pore water
nutrient and carbonate results, Earth and Planetary Science Letters, 49, 57-80.
Emerson, S., S. Mecking, and J. Abell (2001), The biological pump in the subtropical North
Pacific Ocean: Nutrient sources, Redfield ratios, and recent changes. Global Biogeochemical
Cycles, 15, 535–554.
Emile-Geay, J., K.M. Cobb, M.E. Mann, and A.T. Wittenberg (2013), Estimating central
equatorial Paciffic SST variability over the past millennium. Part II: reconstructions and
implications, Journal of Climate, 26(7), 2329-2352, doi:10.1175/JCLI-D-11-00511.1.
Emmer, E. and R. Thunell (2000), Nitrogen isotope variations in Santa Barbara Basin sediments:
implications for denitrification in the eastern tropical North Pacific during the last 50,000
years, Paleoceanography, 15, 377-387.
Eugster O., and N. Gruber (2012), A probabilistic estimate of global marine N-fixation and
denitrification, Global Biogeochemical Cycles, 26, GB4013.
Feely, R.A., C.L. Sabine, J.M. Hernandez-Ayon, D. Ianson, and B. Hales (2008), Evidence for
upwelling of corrosive "acidified" water onto the continental shelf, Science, 320, 1490-1492.
Fernández-Barajas ME, M.A. Monreal-Gómez, A. Molina-Cruz (1994), Thermohaline structure
and geostrophic flow in the Gulf of California, during 1992, Ciencias Marinas 20, 267–286.
Fontugne, M., M. Carre, I. Bentaleb, M. Julie, and D. Lavallee (2004), Radiocarbon reservoir
age variations in the south Peruvian upwelling during the Holocene, Radiocarbon, 46(2), 51-
537.
Foster, G. (1996), Wavelets for period analysis of unevenly sampled time series, Astron. J.,
112(4), 1709–1729.
170
Friis-Christensen, E. and K. Lassen (1991), An indicator of solar activity closely associated with
climate, Science, 254(5032), 698-700.
Galbraith, E.D., M. Kienast and the NICOPP working group members (2013), The acceleration
of oceanic denitrification during deglacial warming, Nature Geoscience, 6, 579-584,
doi:10.1038/NGEO1832.
Ganeshram, R.S., T.F. Pedersen, S.E. Calvert, and J.W. Murray (1995), Large changes in oceanic
nutrient inventories from glacial to interglacial periods, Nature, 376, 755-758.
Ganeshram, R. S., T. F. Pedersen, S. E. Calvert, G. W. McNeil, and M. R. Fontugne (2000),
Glacial-interglacial variability in denitrification in the world’s oceans: Causes and
consequences, Paleoceanography, 15(4), 361–376, doi:10.1029/1999PA000422.
Garcia, H., A. Cruzado, L. Gordon, and J. Escanez (1998), Decadal-scale chemical variability in
the subtropical North Atlantic deduced from nutrient and oxygen data, Journal of
Geophysical Research, 103, 2817– 2830.
Getzlaff, J. and H. Dietze (2013), Effects of increased isopycnal diffusivity mimicking the
unresolved equatorial intermediate current system in an earth system climate model,
Geophysical Research Letters, 40, 2166-2170, doi:10.1002/grl.50419.
Gilly, W.F., J.M. Beman, S.Y. Litvin, and B.H. Robinson (2013), Oceanographic and biological
effects of shoaling of the oxygen minimum zone, Annual Reviews Marine Science, 5, 21.1-
21.28, doi:10.1146/annurev-marine-120710-100849.
Gonzalez-Yajimovich, O., R. G. Douglas, and D. S. Gorsline (2005), The preserved carbonate
record in Holocene sediments of the Alfonso and Pescadero basins, Gulf of California,
Mexico, Proceedings of the Geologist Association, 116, 315–330, doi:10.1016/S0016-
7878(05)80050-1.
171
Gonzalez-Yajimovich, O., R.G. Douglas, and D. Gorsline (2007), Holocene record of ocean-
climate variation in the Gulf of California: rainfall and productivity, Proceedings of the
Twenty-first annual Pacific Climate Workshop, Asilomar Conference Grounds, California,
Technical Report 77 of the Interagency Ecological program for the San Francisco Estuary,
59-68.
Granger, J., D.M. Sigman, M.M. Rohde, M.T. Maldonado, P.D. Tortell (2010), N and O isotope
effects during nitrate assimilation by unicellular prokaryotic and eukarotic plankton cultures,
Geochimica et Cosmochimica Acta, 74(3), 1030-1040, doi:10.1016/j.gca.2009.10.044.
Gruber, N. (2004), The dynamics of the marine nitrogen cycle and its influence on the
atmospheric CO
2
variations, In: Carbon-Climate Interactions NATO ASI (eds., Follows, M.
and Oguz, T.), John Wiley and Sons, New York: 1-39.
Hagadorn, J.W. (1996), Laminated sediments of Santa Monica Basin, California continental
borderland, Geological Society, London, Special Publications, 116, 111-120,
doi:10.1144/GSL.SP.1996.116.01.11.
Hamersley, M.R., G. Lavik, D. Woebken, J.E. Fattray, P. Lam, E.C Hopmans, J.S.S. Damste, S.
Kruger, M. Graco, D. Gutierrez, and M.M.M. Kuypers (2007), Anaerobic ammonium
oxidation in the Peruvian oxygen minimum zone, Limnology and Oceanography, 52(3), 923-
933.
Haug, G.H., T.F.Pedersen, D.M. Sigman, S.E. Calvert, B. Nielsen, and L.C. Peterson (1998),
Glacial/interglacial variations in production and nitrogen fixation in the Cariaco Basin during
the last 580 kyr, Paleoceanography, 13(5), 427-432.
172
Hausback, B.P (1984), Cenozoic volcanic and tectonic evolution of Baja California Sur, Mexico,
In: Geology of the Baja California Peninsula (ed. F. Virgil), Pacific Section Society for
Sedimentary Geology, 39, 219-236.
Hendy, I.L., and T.F. Pedersen (2006), Oxygen minimum zone expansion in the eastern tropical
North Pacific during deglaciation, Geophysical Research Letters, 33, L20602,
doi:10.1029/2006GL025975.
Hennekam, R., and G. De Lange (2012), X-ray fluorescence core scanning of wet marine
sediments: methods to improve quality and reproducibility of high-resolution
paleoenvironmental records, Limnology and Oceanography: Methods, 10, 991-1003.
Hickey, B.M. (1992), Circulation over the Santa Monica-San Pedro Basin and shelf, Progress in
Oceanography, 30, 37-115, doi:10.1016/0079-6611(92)90009-O.
Ingram, B. L., and J. R. Southon (1996), Reservoir ages in eastern Pacific coastal and estuarine
waters, Radiocarbon, 38, 573–582.
Ito, T., and C. Deutsch (2013), Variability of the oxygen minimum zone in the tropical North
Pacific during the late twentieth century, Global Biogeochemical Cycles, 27(4), 1119-1128.
Jenkins, R., and J.L. De Vries (1970), Practical X-ray Spectrometry, MacMillan, London, pp. 1–
190.
Jones, C.E. and H.C. Jenkyns (2001), Seawater strontium isotopes, ocean anoxic events, and
seafloor hydrothermal activity in the Jurassic and Cretaceous, American Journal of Science,
301, 112-149.
Kalansky, J.F., R.S. Robinson, B.N. Popp (2011), Insights into nitrogen cycling in the western
Gulf of California from the nitrogen isotopic composition of diatom‐bound organic matter,
Geochemistry, Geophysics, Geosystems, 12(6), Q06015, doi:10.1029/2010GC003437.
173
Keeling, R.F. and H.E. Garcia (2002), The change in oceanic O
2
inventory associated with recent
global warming, Proceedings of the National Academy of Sciences, 99 (12), 7848-7853.
Keeling, R.F., A. Körtzinger, and N. Gruber (2010), ocean deoxygenation in a warming world,
Annu. Rev. Mar. Sci. 2, 199-229, doi:101146/annurev.marine.010908.163855.
Kemp, A.E.S. (1996), Laminated sediments as palaeo-indicators, Geological Society, London,
Special Publications, 116, vii-xii, doi:10.1144/GSL.SP.1996.116.01.01.
Kienast, S. S., S. E. Calvert, and T. F. Pederson (2002), Nitrogen isotope and productivity
variations along the northeast Pacific margin over the last 120 kyr: Surface and subsurface
paleoceanography, Paleoceanography, 17(4), 1055, doi:10.1029/2001PA000650.
Khider, D., C. S. Jackson, and L. D. Stott (2014), Assessing millennial-scale variability during
the Holocene: A perspective from the western tropical Pacific, Paleoceanography, 29, 143–
159, doi:10.1002/2013PA002534.
King, J.W., and J.E.T. Channell (1991), Sedimentary magnetism, environmental magnetism, and
magnetostratigraphy. Reviews of Geophysics, supplement, US National Report to IUGG,
358-370.
Kuypers, M.M.M, G. Lavik, D. Woebken, M. Schmid, B.M. Fuchs, R. Amann, BB Jorgensen,
and m.S.M. Jetten (2005), Massive nitrogen loss from the Benguela upwelling system
through anaerobic ammonium oxidation, Proceedings of the National Academy of Sciences,
102(18): 6478-6483, doi︎10.1073︎pnas.0502088102.
Lam, P., G. Lavik, M.M. Jensen, J. van de Vossenberg, M. Schmid, D. Woebken, D. Gutiérrez,
R. Amann, M.S.M. Jetten, and M.M.M. Kuypers (2009), Revising the nitrogen cycle in the
Peruvian oxygen minimum zone, Proceedings of the National Academy of Sciences, 106(12),
4752-4757.
174
Langer, G., N. Gussone, G. Nehrke, U. Riebesell, A. Eisenhauer, H. Kuhnert, B. Rost, S.
Trimborn, and S. Thoms (2006), Coccolith strontium to calcium ratios in Emiliania huxleyi:
the dependence on seawater strontium and calcium concentrations, Limnology and
Oceanography, 51(1), 310-320.
Lehmann, M. F., S. M. Bernasconi, A. Barbiere, and J. A. McKenzie (2002), Preservation of
organic matter and alteration of its carbon and nitrogen isotope composition during simulated
and in situ early sedimentary diagenesis, Geochim. Cosmochim. Acta, 66(20), 3573–3582,
doi:10.1016/S0016-7037(02)00968-7.
Levin, L.A., and N. LeBris (2015), The deep ocean under climate change, Science, 350(6262),
766-768, doi:10.1126/science.aad0126.
Liu, K-K, and I.F. Kaplan (1989), The Eastern Tropical Pacific as a source of 15N-enriched
nitrate in seawater off Southern California, Limnology and Oceanography, 34(5), 820-830.
Lluch-Belda, D., M.E. Hernández-Rivas, R. Saldierna-Martínez, and R. Guerrero-Caballero
(2000), Variabilidad de la temperatura superficial del mar en Bahía Magdalena, BCS.
Oceánides, 15(1), 1-23.
Lluch-Belda, D., D.B. Lluch-Cota, and S.E. Lluch-Cota (2003), Baja California’s biological
transition zones: refuges for the California sardine, Journal of Oceanography, 59, 503-513.
Lomb, N. R. (1976), Least-squares frequency analysis of unequally spaced data, Astrophys.
Space Sci., 39, 447–462.
Lynn, R.J., and J.J. Simpson (1987), The California Current System: The seasonal variability of
its physical characteristics, Journal of Geophysical Research, 92(C12), 12,947-12,966.
MacDonald, G.M. and R.A. Case (2005), Variations in the Pacific Decadal Oscillation over the
past millennium, Geophysical Research Letters, 32: L08703, doi:10.1029/2005GL022478.
175
Mantua, N. J., and S. R. Hare (2002), The Pacific Decadal Oscillation, J. Oceanogr., 58, 35–44,
doi:10.1023/A:1015820616384.
Marchitto, T.M., R. Muscheler, J.D. Ortiz, J.D. Carriquiry, and A. VanGeen (2010), Dynamical
response of the tropical pacific ocean to solar forcing during he early Holocene, Science,
330, 1378-1381.
Martinez, P.M., and R.S. Robinson (2010), Increase in water column denitrification during the
last deglaciation: the influence of oxygen demand in the eastern equatorial Pacific,
Biogeosciences, 7, 1-9, doi:10.5194/bg-7-1-2010.
Matear, .J., A.C. Hirst, and B.I. McNeil (2000), Changes in dissolved oxygen in the Southern
Ocean with climate change, Geochemistry, Geo. Geophysics, Geosystems, 12000GC000086.
McClatchie, S., R. Goericke, R. Cosgrove, and G. Auad, R. Vetter (2010), Oxygen in the
Southern California Bight: Multidecadal trends and implications for demersal fisheries,
Geophysical Research Letters, 37, L19602, doi: 10.1029/2010GL044497.
Meko, D.M. (1992), Spectral properties of tree-ring data in the United States Southwest as
related to El Niño/Southern Oscillation, in El Niño: Historical and Paleoclimatic Aspects of
the Southern Oscillation, edited by H.F Diaz and V. Markgraf, 227-242.
Minobe, S. (1997), A 50-70 year climatic oscillation over the North Pacific and North America,
Geophysical Research Letters, 24(6), 683-686.
Montoya, J.P. (2007), Natural abundance of
15
N in marine planktonic ecosystems, Stable
isotopes in ecology and environmental science, 176.
Murray, J.W., L.A. Codispoiti, G.E. Friederich (1995), Oxidation-reduction environments: the
suboxic zone in the Black Sea, In: Aquatic Chemistry: Interfacial and Interspecies Processes
176
(eds. Huang, C.P., C.R. O’Melia, and J.J. Morgan), American Chemical Society,
Washington, DC, 157-176.
Ogurtsov, M.G., Y.A. Nagovitsyn, G.E. Kocharov, and H. Junger (2002), Long-period cycles of
the sun’s activity recorded in direct solar data and proxies, Solar Physics, 211, 371-394.
Olsson, I. (1986), Radiometric Methods. In: Berglund, B. (eds) Handbook of Holocene
palaeoecolgy and palaeohydrology, John Wiley & Sons, Chichester, pp 273-312.
Palmer, M.R. and J.M. Edmond (1989), The strontium isotope budget of the modern ocean,
Earth and Planetary Science Letters, 92, 11-26.
Pichevin, L., R.S. Ganeshram, B.C. Reynolds, F. Prahl, T.F. Pederson, R. Thunell, and E.L.
McClymont (2012), Silicic acid biogeochemistry in the Gulf of California: Insignts from
sedimentary Si isotopes, Paleooceanography, 27, PA2201, doi:10.1029/2011PA002237.
Pierce, S.D., J.A. Barth, R.K. Shearman, and A.Y. Erofeev (2012), Declining oxygen in the
northeast Pacific, Journal of Physical Oceanography, 42(3), 495-501, 10.1175/JPO-D-11-
0170.1
Pike, J. and A.E.S. Kemp (1997), Early Holocene decadal-scale ocean variability recorded in
Gulf of California laminated sediments, Paleoceanography, 12(2), 227-238.
Pride, C., R. Thunell, D. Sigman, L. Keigwin, M. Altabet, E. Tappa (1999), Nitrogen isotopic
variations in the Gulf of California since the Last Deglaciation: Response to global climate
change, Paleococeanography, 14(3): 397-409, doi: 1.1029/1999PA900004.
Prince, E.D., C.P. Goodyear (2006), Hypoxia-based habitat compression of tropical pelagic
fishes, Fisheries Oceanography, 25: 451-464.
Prokopenko, M. (2004), Fractionation of nitrogen isotopes during early diagenesis, Ph.D
Dissertation, University of Southern California, Los Angeles.
177
Prokopenko, M., D. E. Hammond, A. J. Spivack, and L. Stott (2006a), Impact of long-term
diagenesis on δ
15
N of organic matter in marine sediments: Sites 1227 and 1230, Proceedings
from the Ocean Drilling Program, Scientific Results, 201, 30.
Prokopenko, M., D. E. Hammond, and L. Stott (2006b), Lack of isotopic fractionation of d15N
of organic matter during long-term diagenesis in marine sediments, ODP Leg 202, Sites 1234
and 1235, Proceedings from the Ocean Drilling Program, Scientific Results, 202, 22.
Prokopenko, M.G., D.E. Hammond, W.M. Berelson, J.M. Bernhard, L. Stott, R. Douglas
(2006c), Nitrogen cycling in the sediments of Santa Barbara basin and the Eastern
Subtropical North Pacific: Nitrogen isotopes, diagenesis and possible chemosymbiosis
between two lithotrophs (Thioploca and Anammox)—“riding on a glider.” Earth and
Planetary Science Letters, 242, 186-204.
Prokopenko, M., D.M. Sigman, W.M. Berelson, D.E. Hammond, B. Barnett, L. Chong, and A.
Townsend-Small (2011), Denitrification in anoxic sediments supported by biological nitrate
transport, Geochimica et Cosmochimica Acta, 75, 7180-7199,
doi:10.1016/j.gca.2011.09.023.
Prokopenko, M.G., M.B. Hirst, L. De Brabandere, D.J.P. Lawrence, W.M. Berelson, J. Granger,
B.X. Chang, S. Dawson, E.J. Crane III, L. Chong, B. Thamdrup, A. Townsend-Small, and
D.M. Sigman (2013), Nitrogen losses in anoxic marine sediments driven by Thioploca-
anammox bacterial consortia, Nature, 500, 194-200, doi:10.1038/nature12365.
Purcell, J.E., D.L. Breitburg, M.B Decker, W.M Graham, M.J. Youngbluth, K.A. Raskoff
(2001), Pelagic cnidarians and ctenophores in low dissolved oxygen environments. In:
Costal Hypoxia: Consequences for Living Resources and Ecosystems (eds. Rabalais, N.B.
and R.E. Turner), American Geophysical Union, Washington, DC, 77-100.
178
Purcell, J.E., S. Uye, W.T. Lo (2007), Anthropogenic causes of jellyfish blooms and their direct
consequences for humans: a review, Marine Ecology Progress Series, 350, 153-174.
Ragueneau, O., P. Tréguer, A. Leynaert, R.F. Anderson, M.A. Brzezinski, D.J. DeMaster, R.C.
Dugdale, J. Dymond, G. Fischer, R. Francois, C. Heinze, E. Maier-Reimer, V. Martin-
Jézéquel, D.M. Nelson, and B. Quéguiner (2000), A review of the Si cycle in the modern
ocean: recent progress and missing gaps in the application of biogenic opal as a
paleoproductivity proxy, Global and Planetary Change, 26: 317-365, doi:10.1016/S0921-
8181(00)00052-7.
Reimer P.J., E. Bard, A. Bayliss, J.W. Beck, P.G. Blackwell, C. Bronk Ramsey, C.E. Buck, H.
Cheng, R.L. Edwards, M. Friedrich, P.M. Grootes, T.P. Guilderson, H. Haflidason, I. Hajdas,
C. Hatté, T.J. Heaton, A.G. Hogg, K.A. Hughen, K.F. Kaiser, B. Kromer, S.W. Manning, N.
Niu, R.W. Reimer, D.A. Richards, E.M. Scott, J.R. Southon, C.S.M. Turney, J. van der Plicht
(2013), IntCal13 and MARINE13 radiocarbon age calibration curves 0-50000 years calBP,
Radiocarbon 55(4), doi:10.2458/azu_js_rc.55.16947.
Riser, S.C. and K.S. Johnson (2008), Net production of oxygen in the subtropical ocean, Nature,
451(7176), 323-325, doi:10.1038/nature06441.
Robinson, R.S., D.M. Sigman, P.J. DiFiore, M.M. Rohde, T.A. Mashiotta, and D.W. Lea (2005),
Diatom-bound
15
N/
14
N: new support for enhanced nutrient consumption in the ice age
subantarctic, Paleoceanography, 20, PA3003, doi:10.1029/2004PA001114.
Robinson, R.S., M. Kienast, A.L. Albuquerque, M. Altabet, S. Contreras, R. De Pol Holz, N.
Dubois, R. Francois, E. Galbraith, T-C. Hsu, T. Ivanochko S. Jaccard, S-J. Kao, T. Kiefer, S.
Kienast, M. Lehmann, P. Martinez, M. McCarthy, J. Möbius, T. Pedersen, T.M. Quan, E.
Ryabenko, A. Schmittner, R. Schneider, A. Schneider-Mor, M. Shigemitsu, D. Sinclair, C.
179
Somes, A. Studer, R. Thunell, and J-Y. (2012), A review of nitrogen isotopic alteration in
marine sediments, Paleoceanography, 27, PA2403, doi:10.1029/2012PA002321.
Rothwell, R. G., and F.R. Rack (2006), New techniques in sediment core analysis: an
introduction, Geological Society, London, Special Publications, 267(1), 1-29.
Rykaczewski, R.R., and J.P. Dunne (2010), Enhanced nutrient supply to the California Current
Ecosystem with global warming and increased stratification in an earth system model,
Geophysical Research Letters, 37(21), L21606, doi:10.1029/2010GL045019.
Sañudo-Wilhelmy, S.A., L.S. Cutter, R. Durazo, E.A. Smail, L. Gómez-Consarnau, E.A. Webb,
M.G. Prokopenko, W.M. Berelson, and D.M. Karl (2012), Multiple B-vitamin depletion in
large areas of the coastal ocean, Proceedings of the National Academy of Sciences, 109(35),
14041-14045.
Sarmiento, J. L., T.M.C. Hughes, R.J. Stouffer, S. Manabe (1998), Simulated response of the
ocean carbon cycle to anthropogenic climate warming, Nature, 393; 245– 249.
Scargle, J. D. (1982), Studies in astronomical time series analysis. II. Statistical aspects of
spectral analysis of unevenly spaced data, Astrophys J., 263(2), 835–853.
Scargle, J. D. (1989), Studies in astronomical time series analysis. III. Fourier transforms,
autocorrelation functions, and cross-correlation functions of unevenly spaced data,
Astrophys. J., 343(2), 874–887.
Schimmelmann, A. and M.J. Tegner (1991), Historical oceanographic events reflected in
13
C/
12
C
ratio of total organic carbon in laminated Santa Barbara Basin sediment, Global
Biogeochemical Cycles, 5(2), 173-188.
180
Schimmelmann A. and C. Lange (1996), Tales of 1001 varves: a review of Santa Barbara Basin
sediment studies, Geological Society, London, Special Publications, 116, 121-141,
doi:10.1144/GSL.SP.1996.116.01.12.
Schimmelmann, A., I.L. Hendy, L. Dunn, D.K. Pak, and C.B. Lange (2013), Revised ∼2000-year
chronostratigraphy of partially varved marine sediment in Santa Barbara Basin, California,
GFF Special Issue: Varve Genesis, Chronology and Paleoclimate, 135(3-4),
doi:10.1080/11035897.2013.773066.
Schlesinger M, .E., and N. Ramankutty (1994), An oscillation in the global climate system of
period 65-70 years, Nature, 3(67), 723-726.
Schulz, M., and K. Stattegger (1997), SPECTRUM: Spectral analysis of unevenly spaced time
series, Comput. Geosci., 23(9), 929–945.
Sharp, G.D. (1992), Fishery catch records, El Niño/Southern Oscillation, and longer-term
climate change as inferred from fish remains in marine sediments, in El Niño: Historical and
Paleoclimatic Aspects of the Southern Oscillation, edited by H.F. Diaz and V. Markgraf,
379-418.
Sigman, D. M., R. Robinson, A. N. Knapp, A. van Geen, D. C. McCorkle, J. A. Brandes, and R.
C. Thunell (2003), Distinguishing between water column and sedimentary denitrification in
the Santa Barbara Basin using the stable isotopes of nitrate, Geochemistry, Geophysics,
Geosystems, 4(5), 1040, doi:10.1029/2002GC000384.
Sigman, D. M., J. Granger, P. J. Difiore, M. M. Lehmann, R. Ho, G. Cane, and A. Van Geen
(2005), Coupled nitrogen and oxygen isotope measurements of nitrate along the eastern
North Pacific margin, Global Biogeochem. Cycles, 19, GB4022, doi:10.1029/
2005GB002458.
181
Silverberg, N., A. Martıınez, S. Aguíñiga, J. D. Carriquiry, N. Romero, E. Shumilin, and S. Cota
(2004), Contrasts in sedimentation flux below the southern California Current in late 1996
and during the El Niño event of 1997–1998, Estuarine, Coastal and Shelf Science, 59(4),
575–587, doi:10.1016/j.ecss.2003.11.003.
Smith, K. L., H.A. Ruhl, M. Kahru, C. L. Huffard, and A.D. Sherman (2013), Deep ocean
communities impacted by changing climate over 24 y in the abyssal northeast Pacific Ocean,
Proceedings of the National Academy of Sciences, 110(49), 19838-19841,
doi:10.1073/pnas.1315447110.
Somes CJ, A. Schmittner, and M.A. Altabet (2010), Nitrogen isotope simulations show the
importance of atmospheric iron deposition for nitrogen fixation across the Pacific Ocean,
Geophysical Research Letters, 37, L23605.
Southon, J.R. and G.M. Santos (2004), Ion source development at KCC AMS, University of
California, Irvine. Radiocarbon 46, 33–39.
Southon, J.R. and G.M. Santos (2007), Life with MC-SNICS. Part II: Further ion source
development at the Keck carbon cycle AMS facility. Nuclear Instruments and
Methods in Physics Research, Section B 259, 88–93.
Staines-Urías, F., R.G. Douglas, D.S. Gorsline (2009), Oceanographic variability in the southern
Gulf of California over the past 400years: Evidence from faunal and isotopic records from
planktic foraminifera, Palaeogeography, Palaeoclimatology, Palaeoecology, 284 (3-4): 337-
354, doi: 10.1016/j.palaeo.2009.10.016.
Steinhilber, F., J. Beer, and C. Fröhlich (2009), Total solar irradiance during the Holocene,
Geophys. Res. Lett., 36, L19704, doi:
10.1029/2009GL040142.
182
Stramma, L., GC. Johnson, J. Sprintall, V. Mohrholz (2008), Expanding oxygen-minimum
zones in the tropical oceans, Science, 320: 655-658.
Stramma, L., S. Schmidtko, L.A. Levin, G.C. Johnson (2010), Ocean oxygen minima expansions
and their biological impacts. Deep-Sea Research I, 57, 587-595.
Straub, M., D.M. Sigman, H. Ren, A. Martinez-Garcia, A.N. Meckler, M.P. Hain, and G.H.
Haug (2013), Changes in North Atlantic nitrogen fixation controlled by ocean circulation,
Nature, 501(7466), 200-203, doi:10.1038/nature12397.
Strickland, J.P.T. (1968), A Practical Handbook of Seawater Analysis, Fisheries Research Board
of Canada, Halifax.
Strous, M., J.A. Fuerst, E.H.M. Kramer, S. Logemann, G. Muyxer, K.T. Van de Pas-schoonen,
R. Webb, G.C. Kuenen, M.S.M. Jetten (1999), Missing lithotroph identified as new
planctomycetes, Nature, 400, 446-449.
Stuiver, M. (1980), Workshop on
14
C data reporting, Radiocarbon, 22, 964-966.
Stuiver, M., and A.H. Polach (1977), Discussion: Reporting of 14C data, Radiocarbon, 19:355-
363.
Stuiver, M. and P.J. Reimer (1993), Extended 14C database and revised CALIB radiocarbon
calibration program, Radiocarbon, 35:215-230.
Sweeney, R.E. and I.R. Kaplan (1980), Natural abundances of N-15 as a source indicator for
near-shore marine sedimentary dissolved nitrogen, Marine Chemistry, 9, 81-94.
Taft, W.H, and J.W. Harbaugh (1964), Modern carbonate sediments of southern Florida,
Bahamas, and Espiritu Santo Island, Baja California: a comparison of their mineralogy and
chemistry, Stanford University Press.
183
Tems, C. E., W.M. Berelson, and M.G. Prokopenko (2015), Particulate δ
15
N in laminated marine
sediments as a proxy for mixing between the California Undercurrent and the California
Current: A proof of concept, Geophysical Research Letters, 42, doi:10.1002/
2014GL061993.
Tems, C.E., W.M. Berelson, R. Thunell, E. Tappa, X. Xu, D. Khider, S. Lund and O. González-
Yajimovich, Sedimentary δ
15
N reveal decadal fluctuations in the intensity of the eastern
tropical north Pacific oxygen minimum zone during the last 1200 years, submitted to
Paleoceanography November 2015.
Thunell, R.C., C.J. Pride, E. Tappa, and F.E. Muller-Karger (1994), Biogenic silica fluxes and
accumulation rates in the Gulf of California, Geology, 22, 303-306, doi: 10.1130/0091-
7613(1994)022<0303:BSFAAR>2.3.CO;2.
Thunell, R. C., and A. B. Kepple (2004), Glacial-Holocene δ
15
N record from the Gulf of
Tehuantepec, Mexico: Implications for denitrification in the eastern equatorial Pacific and
changes in atmospheric N2O, Global Biogeochem. Cycles, 18, GB1001,
doi:10.1029/2002GB002028.
Thunell, R., D.M. Sigman, F.E. Muller-Karger, Y. Astor, R. Varela (2004), Nitrogen Isotopic
Dynamics of the Cariaco Basin, Venezuela, Global Biogeochemical Cycles, 18: GB3001,
doi:10.1029/2003GB002185.
Thurow, J., and M. Schaaf (1995), A comparison of Holocene to late Pleistocene grayvalue-time-
series of sediments from the Santa Barbara Basin (ODP Site 893) and the Guaymas Basin
(DSDP Site 480), in IGCP 374: Palaeoclimatology and Palaeoceanography From Laminated
Sediments, edited by R.J. Behl and J.P. Kennett, 33.
184
Townsend-Small, A., M. G. Prokopenko, and W. M. Berelson (2014), Nitrous oxide cycling in
the water column and sediments of the oxygen minimum zone, eastern subtropical North
Pacific, Southern California, and Northern Mexico (23°N–34°N), J. Geophys. Res. Oceans,
119, 3158–3170, doi:10.1002/2013JC009580.
VanGeen, A., Y. Zheng, J.M. Bernhard, K.G. Cannariato, J. Carriquiry, W.E. Dean, B.W.
Eakins, J.D. Ortiz, and J. Pike (2003), On the preservation of laminated sediments along the
western margin of North America, Paleoceanography, 18(4), 1098,
doi:10.1029/2003OA000911.
Von Damm, K. L., J.M. Edmond, B. Grant, C.I. Measures, B. Walden, and R.F. Weiss (1985),
Chemistry of submarine hydrothermal solutions at 21° N, East Pacific Rise, Geochimica et
Cosmochimica Acta, 49(11), 2197-2220.
Watanabe, Y. W., T. Ono, A. Shimamoto, T. Sugimoto, M. Wakita, and S. Watanabe (2001),
Probability of a reduction in the formation rate of the subsurface water in the North Pacific
during the 1980s and 1990s, Geophysical Research Letters, 28, 3289– 3292.
Ward, B.B., C.B. Tuit, Am Jayakumar, J.J. Rich, J. Moffett, S. Wajih, and A. Naqvi (2008),
Organic carbon, not copper, controls denitrification in oxygen minimum zones of the ocean,
Deep-Sea Research I, 55: 1672-1683.
Ward, B.B, A.H. Devol, J.J. Rich, B.X. Chang, S.E. Bulow, H. Naik, A. Pratihary, and A.
Jayakumar (2009), Denitrification as the dominant nitrogen loss process in the Arabian Sea,
Nature, 461, 78-82, doi:10.1038/nature08276.
Wasser, N.A.D., P.J. Harrison, B. Nielsen, and S.E. Calvert (1998), Nitrogen isotope
fractionation during uptake and assimilation of nitrate, nitrite, ammonium, and urea by a
185
marine diatom, Limnology and Oceanography, 43(2), 215-224,
doi:0.4319/lo.1998.43.2.0215.
Webber, A., and B.B. Jorgensen (2002), Bacterial sulfate reduction in hydrothermal sediments of
the Guaymas Basin, Gulf of California, Mexico, Deep-Sea Research I 49(5),827-841.
Welch, P. D. (1967), The use of Fast Fourier transform for the estimation of power spectra: A
method based on time averaging over short, modified periodograms, IEEE Trans. Audio
Electroacoust., 15(2), 70–73.
White, A.E., R.A. Foster, C.R. Benitez-Nelson, P. Masqué, E. Verdeny, B.N. Popp, K.E. Arthur,
and F.G. Prahl (2013), Nitrogen fixation in the Gulf of California and the Eastern Tropical
North Pacific, Progress in Oceanography, 109, 1-17, doi:10.1016/j.pocean2012.09.002.
Wingall, P.B. and R.J. Twitchett (1996), Ocean anoxia and the end Permian mass extinction,
Science, 272(5265), 1155-1158, doi: 10.1126/science.272.5265.1155.
Wolter, K., and M. S. Timlin (2011), El Niño/Southern Oscillation behaviour since 1871 as
diagnosed in an extended multivariate ENSO index (MEI.ext). International Journal of
Climatology, 31(7), 1074-1087, doi:10.1002/joc.2336.
Xu, X., Trumbore, S.E., Zheng, S.H., Southon, J.R., McDuffee, K.E., Luttgen, M., Liu, J.C.
(2007), Modifying a sealed tube zinc reduction method for preparation of AMS graphite
targets: reducing background and attaining high precision. Nuclear Instruments and Methods
in Physics Research, Section B 259, 320–329.

186
Appendix A.

The published article included in this appendix, “Centennial changes in North Pacific anoxia
linked to tropical trade winds,” is the result of collaborative effort in which I contributed. The
authors of this work include Curtis Deutsch, William Berelson, Robert Thunell, Thomas Weber,
myself, James McManus, John Crusius, Taka Ito, Timothy Baumgartner, Vicente Ferreira, Jacob
Mey, and Alexander van Geen.

Citation:
Deutsch, C., W. Berelson, R. Thunell, T. Weber, C. Tems, J. McManus, J. Crusius, T. Ito, T.
Baumgartner, V. Ferreira, J. Mey, and A. van Geen (2014), Centennial changes in North Pacific
anoxia linked to tropical trade winds, Science, 345, 665-668, doi: 10.1126/science.1252332

187
9. T. J. Buschman, E. K. Miller, Science 315, 1860–1862 (2007).
10. G. G. Gregoriou, S. J. Gotts, H. Zhou, R. Desimone, Science
324, 1207–1210 (2009).
11. T. Moore, K. M. Armstrong, Nature 421, 370–373 (2003).
12. B. Noudoost, T. Moore, Nature 474, 372–375 (2011).
13. L. B. Ekstrom, P. R. Roelfsema, J. T. Arsenault, G. Bonmassar,
W. Vanduffel, Science 321,414–417 (2008).
14. Y. Chen et al., Nat. Neurosci. 11, 974–982 (2008).
15. K. A. Sundberg, J. F. Mitchell, J. H. Reynolds, Neuron 61,
952–963 (2009).
16. K. McAlonan, J. Cavanaugh, R. H. Wurtz, Nature 456, 391–394
(2008).
17. G. Purushothaman, R. Marion, K. Li, V. A. Casagrande,
Nat. Neurosci. 15, 905–912 (2012).
18. M. Sarter, M. E. Hasselmo, J. P. Bruno, B. Givens, Brain Res.
Brain Res. Rev. 48,98–111 (2005).
19. Y. B. Saalmann, M. A. Pinsk, L. Wang, X. Li, S. Kastner, Science
337, 753–756 (2012).
20. M. A. Segraves, M. E. Goldberg, J. Neurophysiol. 58, 1387–1419
(1987).
21. The projection to superior colliculus arises only from deep
layers (Fig. 1D), similar to that in primate FEF.
22. S. Treue, J. C. Martínez Trujillo, Nature 399, 575–579 (1999).
23. T. Moore, M. Fallah, J. Neurophysiol. 91, 152–162 (2004).
24. X. Xu, K. D. Roby, E. M. Callaway, J. Comp. Neurol. 518,
389–404 (2010).
25. G. Fishell, B. Rudy,Annu.Rev.Neurosci. 34,535–567 (2011).
26. H. Adesnik, W. Bruns, H. Taniguchi, Z. J. Huang, M. Scanziani,
Nature 490, 226–231 (2012).
27. S. H. Lee et al., Nature 488, 379–383 (2012).
28. N. R. Wilson, C. A. Runyan, F. L. Wang, M. Sur, Nature 488,
343–348 (2012).
29. L. Petreanu, T. Mao, S. M. Sternson, K. Svoboda, Nature 457,
1142–1145 (2009).
30. I. R. Wickersham et al., Neuron 53, 639–647 (2007).
31. S. Lee, I. Kruglikov, Z. J. Huang, G. Fishell, B. Rudy, Nat. Neurosci.
16, 1662–1670 (2013).
32. L. Madisen et al., Nat. Neurosci. 15, 793–802 (2012).
33. Y. Fu et al., Cell 156, 1139–1152 (2014).
34. C. K. Pfeffer, M. Xue, M. He, Z. J. Huang, M. Scanziani,
Nat. Neurosci. 16, 1068–1076 (2013).
35. H. J. Pi et al., Nature 503, 521–524 (2013).
36. J. F. Mitchell, K. A. Sundberg, J. H. Reynolds, Neuron 55,
131–141 (2007).
37. Based on a magnification factor of 10 mm per degree in mouse
V1 (42), 200 mm of cortical distance corresponds to 20° of
visual angle. A previous study (26) showed that for most
neurons, the preferred size of visual stimulus was <15° in
radius, and stimuli beyond this radius suppressed neuronal
responses. This suggests that surround suppression for
top-down modulation and bottom-up processing occur on
similar spatial scales. The same inhibitory circuits could also
contribute to decreased receptive field similarity and signal
correlation between V1 neurons over ~200 mm(42).
38. J. J. Letzkus et al., Nature 480, 331–335 (2011).
39. X. Jiang, G. Wang, A. J. Lee, R. L. Stornetta, J. J. Zhu,
Nat. Neurosci. 16, 210–218 (2013).
40. S. P. Mysore, E. I. Knudsen, Nat. Neurosci. 16, 473–478 (2013).
41. S. Ardid, X. J. Wang, A. Compte, J. Neurosci. 27, 8486–8495
(2007).
42. V. Bonin, M. H. Histed, S. Yurgenson, R. Clay Reid, J. Neurosci.
31, 18506 (2011).
ACKNOWLEDGMENTS
WethankL.PintoandY.Zhuforhelpwithdataanalysis;S.H.Leeand
M. Zhao for technical assistance; Standford Neuroscience Gene
Vector and Virus Core for AAV-DJ supply; K. Deisseroth, E. Callaway,
B.Lim,andB.C.Weissbourdforvirusandconstructs;andR.Desimone,
L. Wang, and M. A. Segraves for helpful discussions.This workwas
supportedbyNIHgrantR01EY018861,NSFgrant22250400-42533,a
Uehara Memorial Foundation fellowship, andtheHuman Frontier
Science Program. All primary histological, electrophysiological, and
behavioral data are archived in the Department of Molecular andCell
Biology, University of California, Berkeley.
SUPPLEMENTARY MATERIALS
www.sciencemag.org/content/345/6197/660/suppl/DC1
Materials and Methods
Figs. S1 to S8
References (43–51)
31 March 2014; accepted 27 June 2014
10.1126/science.1254126
OCEANOGRAPHY
CentennialchangesinNorthPacific
anoxialinkedtotropicaltradewinds
CurtisDeutsch,
1
*WilliamBerelson,
2
RobertThunell,
3
ThomasWeber,
1
CaitlinTems,
2
JamesMcManus,
4
†JohnCrusius,
5
TakaIto,
6
TimothyBaumgartner,
7
VicenteFerreira,
7
JacobMey,
8,9
AlexandervanGeen
8
Climate warming is expected to reduce oxygen (O
2
) supply to the ocean and expand its
oxygen minimum zones (OMZs).We reconstructed variations in the extent of North Pacific
anoxia since 1850 using a geochemical proxy for denitrification (d
15
N) from multiple
sediment cores. Increasing d
15
N since ~1990 records an expansion of anoxia, consistent
with observed O
2
trends. However, this was preceded by a longer declining d
15
N trend
that implies that the anoxic zone was shrinking for most of the 20th century. Both
periodscanbeexplainedbychanges inwindsover thetropicalPacific thatdriveupwelling,
biological productivity, and O
2
demand within the OMZ. If equatorial Pacific winds
resume their predicted weakening trend, the ocean’s largest anoxic zone will contract
despite a global O
2
decline.
B
elow the ocean’s surface, the decomposi-
tion of sinking detritus creates a layer of
low-O
2
waterinhospitabletomanymarine
species (1). These oxygen minimum zones
(OMZs) are predicted to expand with cli-
matewarming(2,3),causingamajordisruption
to ecosystems, especially in areas where OMZ
waters impinge on coastal environments already
underlow-O
2
stressfromnaturalorhumancauses
(4–6). This putative expansion stems from two
direct consequences of climate change: As the
surface ocean warms, its gas solubility and den-
sity both decrease, reducing the concentration
of O
2
in surface water and the rate at which
that water is transported downward against a
more stable stratification. The resulting decline
in O
2
supply to the ocean interior is generally
supported by observed trends toward lower O
2
over the past few decades, over many parts of
the world’soceans(7), including the strong OMZs
in the tropics (8).
TheclimaticresponseoftheOMZalsodepends
on O
2
demand, although the factors governing
those biological rates are less well understood.
Earth system model simulations project large
future decreases in the sinking flux of organic
matter throughout the tropics (9), which should
reduce O
2
demand in the OMZ, counteracting
the loss of O
2
supply. In the eastern tropical
andsubtropicalPacificOcean,theOMZvariabil-
ity over the last 50 years appears to have been
driven primarily by O
2
demand, which is
strongly modulated by decadal climate varia-
bility (10, 11). In light of that variability, the
instrumental O
2
record is still sparse and short,
makinglong-termtrendsintheOMZdifficultto
detectorattributeespeciallyinthemostintense
tropical OMZs in the Pacific and Indian oceans
(8,12).Therelativestrengthoffuturechangesin
O
2
supply versus demand in the tropical OMZ
would be clearer if their long-term response to
theclimate warmingsincethe industrialrevolu-
tionwere known.
We reconstructed changes in the OMZ of the
easterntropicalnorthernPacific(ETNP)overthe
past 150 years using a geochemical proxy for
watercolumnanoxiathatisrecordedinsediments
(Fig.1).AttheeasternterminusoftheOMZ,where
thermocline waters shoal toward the produc-
tive surface layer, respiration depletes O
2
and
anaerobicbacteriabeginreducingnitrate(NO
3
–
)
to oxidize organic matter (13)(Fig.1).Thisde-
nitrification process preferentially removes the
lighter
14
N isotope of N, leaving a residual NO
3
–
pool enriched in heavier
15
N(14, 15). The result-
ingnitratewithahighisotoperatio,d
15
N[d
15
N=
(
15
N/
14
N)/R
air
– 1) × 1000, where R
air
is the N
isotope ratio in air], is upwelled to the surface,
where it can be transferred to plankton com-
munities and then, via sinking particles, into
sediments(16,17).Inregionswithcompletecon-
sumptionofupwellednitrate,thed
15
Nofpartic-
ulateorganicnitrogen(PON)accumulatingonthe
seafloorcloselyresemblesthatofnitrateinwaters
at~100m(18).Intheabsenceofpostdepositional
alteration, downcore variations in d
15
Npro-
vide a history of the integrated rates of de-
nitrificationandthesizeoftheOMZtowhich
it is confined (19,20).
To ensure a representative history of changes
in OMZ intensity, we analyzed sediment cores
fromthreesitesalongtheNorthAmericanmargin
SCIENCE sciencemag.org 8 AUGUST 2014 VOL 345 ISSUE 6197 665
1
School of Oceanography, University of Washington, Seattle,
WA, USA.
2
Department of Earth Sciences, University of
Southern California, Los Angeles, CA, USA.
3
Department of
Earth and Ocean Sciences, University of South Carolina,
Columbia, SC, USA.
4
College of Earth, Ocean, and
Atmospheric Sciences, Oregon State University, Corvallis,
OR, USA.
5
U.S. Geological Survey, University of Washington
School of Oceanography, Seattle, WA, USA.
6
School of Earth
and Atmospheric Sciences, Georgia Institute of Technology,
Atlanta, GA, USA.
7
Departamento de Oceanografía Biológica,
Centro de Investigación Científica y de Educación Superior
de Ensenada, Baja California, México.
8
Lamont-Doherty Earth
Observatory of Columbia University, Palisades, NY, USA.
9
Department of Physical Sciences, Kingsborough Community
College, City University of New York, New York, NY, USA.
*Corresponding author. E-mail: cdeutsch@uw.edu †Present
address: Department of Geosciences, University of Akron, Akron,
OH, USA.
RESEARCH | REPORTS
188
athightemporalresolution(21).Twositeslaywithin
the anoxic zone off Baja California (Pescadero
Slope and Soledad Basin). A third is from the
SouthernCaliforniaBight(SantaMonicaBasin),
whichis>1000kmnorthoftheanoxiczonebut
physicallyconnectedviathecoastalundercurrent
(22,23)androutinelysampledbytheworld’slon-
gest regional hydrographic time series, the Cali-
forniaCooperativeOceanicFisheriesInvestigations
(CalCOFI) (6). Sediment samples were taken at
0.8-to3-mmintervals,yieldinganearlyannual
(1-to2-yearaverage)temporalresolution.Allthree
sitescontaindistinctlightanddarklaminaein-
dicatingrelativelyundisturbedsedimentsanda
lack of bioturbation. Measurements of d
15
Nasa
function of sediment age were determined by
meansof
210
Pbchronology(fig.S1).Tominimize
thepotentialforartifactsatthecoretop(24),we
examinedsedimentintervalswithanagegreater
than1 year (21).
During the last 50 years, all three sediment
cores reveal d
15
N values declining gradually by
nearly 1‰between ~1960 andthemid-1980s, fol-
lowed by a more rapid increase of ~1‰ since
~1990 (Fig. 2A). Given the large differences in
sedimentaccumulationrateandcompositionbe-
tween sites (25), the coherence in d
15
N suggests
thattheyhaverecordedprimarilylarge-scaleocean-
ographicchangesoriginatinginthewatercolumn
ratherthaninthesediments.Thisiscorroborated
by an independent 18-year time series of sinking
particulatematterd
15
Nfromsedimenttrapsin
Santa Barbara Basin, which exhibits an increas-
ing trend similar to that of the sediment cores
(Fig. 2A) (21). The timing of these decadal varia-
tionsalsocoincideswithobservedchangesinoxy-
genconcentrationsandnitratedeficitsinhypoxic
waterssampledbyCalCOFI,andwithchangesin
thevolumeofanoxicwaterintheETNPpredicted
by a hindcast model for this period (10). These
sedimentaryd
15
Nrecordsprovideauniquemeans
to quantitatively test the mechanisms linking
climate and anoxia in those model simulations.
Toestimatethemagnitudeofchangesinanoxia
and denitrification implied by the sediment
666 8 AUGUST 2014 VOL 345 ISSUE 6197 sciencemag.org SCIENCE
Fig. 1. Oceanographic context of study sites.
Time-mean map of O
2
(colors) at a density of
Central Mode Water (s
q
= 26.2) whose mean depth
(thin lines) shoals to the east,where anoxia is found
off the equator in both hemispheres. Under anoxic
conditions, oxidation of organic matter proceeds via
denitrification, resulting in large regions (white lines)
in which the primary N nutrient, nitrate (NO
3
–
), falls
below its expected covariation with phosphate
(PO
4
3–
) by more than 10 mM, as measured by the
tracer N* (N* = [NO
3
–
]–16 × [PO
4
3–
]). Locations of
coresitesareshownwithcircles(blue:SantaMonica
Basin; red: Pescadero; green: Soledad). Gray box
defines ETNP region used for analysis of modeled
and observed quantities.
Fig. 2. Changes in sedimentary isotope ratio
(d
15
N)anditsrelationshiptothermoclinedepth
(Z
tc
) in the ETNP, from data and model simula-
tions. (A)Bulkorganic d
15
N anomalies from three
sedimentcoresitesalongtheNorthAmericanmar-
gin (Soledad, Pescadero, Santa Monica Basin) and
from a particle flux time series in Santa Barbara
Basin. Each time series is plotted as a deviation
from its own time-averaged value. (B) Variations of
sinkingparticleflux d
15
Ninanoceanhindcastmod-
el with an explicit N cycle, forced by reconstructed
atmospheric variability.The model d
15
Nfluxisav-
eraged over the anoxic zone (125°W to coast, 5° to
25°N). A range of values is based on uncertainty in
the isotopic fractionation effect (e
WC
=15to25‰)
associated with bacterial denitrification (15, 21).Thin
lines are 5-year running averages of the sediment
core data. (C) Relationship between thermocline
depth (defined as the 13° isotherm) from World
Ocean Database and sediment d
15
NinPescadero
sediment core, and (D) thermocline depth and par-
ticleflux d
15
N inmodelsimulations.Changesin ther-
mocline depth are averaged across the broader
ETNPregion(160°Wtocoast,0°to30°N).Theslope
of the relationship between d
15
N and thermocline
depth in observations (–0.044 T 0.01‰ m
−1
;90%
confidenceinterval)andmodelsimulations(–0.049T
0.011‰ m
−1
) are indistinguishable.
RESEARCH | REPORTS
189
records, we directly simulate changes in sedi-
mentary d
15
N in recent decades using an ocean
generalcirculationmodel(GCM).Themodelis
expanded from previous work to include an
explicit cycle of N and its isotopes (21). It captures
the observed structure of the tropical Pacific
anoxic zone, data-based rates of denitrification
(26), and the basin-scale distribution of d
15
NO
3
(fig. S3). In response to imposed wind and
buoyancy fluxes from atmospheric reanalyses
from1959to2005,theNcyclesimulationspredict
large changes in anoxic water volume and
integrated denitrification rates, roughly consistent
withsimplerstoichiometriccalculations(10)(fig.
S4). The decadal variations in the model-
predicted d
15
N of particulate organic N sinking
into the anoxic zone are driven by the changes
in denitrification and closely match the sedi-
ment proxy data in both amplitude and timing
(Fig. 2B). Additional isotopic variability arising
fromincompletesurfacenutrientutilizationismuch
smaller in magnitude and confined to interan-
nual time scales (fig.S4).
The sedimentary d
15
N data can be used to
test the hypothesized mechanisms underlying
recent changes in the OMZ. According to model
predictions, the decadal differences in the vol-
umeoftheanoxiczonearedrivenbyrespiration
rates (10, 11), which vary concurrently with the
depth of the thermocline waters in which the
OMZisembedded.Ashallowerthermoclinecan
increasetherespirationrateinOMZwatersboth
by increasing surface photosyntheticproduction
andsinkingoforganicmatter,andbypositioning
the low-O
2
waters at depths where decomposi-
tion of the sinking particles is relatively fast. We
tested this relationship using historical observa-
tions from the World Ocean Database (27)to
estimatechangesinthedepthofthe13°Cisotherm
(fig. S5), which lies near the core of the OMZ in
the eastern tropical North Pacific. Regional var-
iations in the thermocline depth are corrected
for heat uptake, and therefore attributable to
changesinoceandynamics(21).Theyarestrongly
correlated to sedimentary d
15
N(Fig.2C),andthe
slope of the relationship in the data is indistin-
guishablefromthatinmodelsimulations(Fig.
2D).Thecoherencebetweenmodeledparticulate
d
15
N and sediment core data suggests that the
d
15
Nrecordsareindeedareliableproxyforlarge-
scale changes in the intensity of the OMZ. The
strong relationship to thermocline depth inde-
pendently observed in model simulations and
data supports the hypothesized link to climate
variability, at least ondecadal time scales.
Thefull150-yearrecordofsedimentaryd
15
N
measurementsrevealsadecliningtrendthrough-
out most of the 20th century (Fig. 3A). During
mostoftheperiodofclimatewarming,theanoxic
OMZ was evidently shrinking. To determine
whether the apparent long-term contraction of
anoxic volume is consistent with declining O
2
demand,weextendedthethermoclinedepthproxy
back to 1850 using Intergovernmental Panel on
ClimateChange(IPCC)climatemodelreconstruc-
tions(fig.S6).Thesemodelsdonotdirectlyincor-
porate oceanic observations, and thus do not
reproducethephasingormagnitudeofinternally
generated climate variability. However, they pro-
vide a self-consistent estimate of the oceanic re-
sponsetoobservedradiativeforcing,primarilyby
greenhouse gases. The ~10-m deepening of the
thermoclineover 100yearsinconjunctionwitha
~1‰ reduction in sediment d
15
Nisconsistent
with the relationship observed over decadal time
scales (Fig.3B).TherelationshipbetweenNorth
Pacific anoxia and thermocline depth diagnosed
overthelastseveraldecadesthusappearstochar-
acterize the dominant climatic mechanism gov-
erninganoxiczonevariabilityfor overacentury.
In the tropical Pacific, high rates of surface
biologicalproductivityandthezonalstructureof
the underlying thermocline are both primarily
governedbythestrengthofsurfacewinds,which
push warm surface water to the west and draw
cold,nutrient-richdeepwaterupward along the
equator in the east. The overall rate of respira-
tion in the tropical thermocline, and its parti-
tioningintotheintenseOMZbythermoclinedepth
fluctuations,shouldthereforebothbecloselylinked
to the atmospheric Walker circulation (fig. S7).
Indeed,historicalchangesinequatorialwindstress,
recordedinthezonalsea-levelpressuregradient
(Fig. 3C), imply a long-term weakening trend
followedbyarecentintensification(28–30).Both
thesefeaturesarereflectedinthermoclinedepth
variationsandthevolumeofanoxiainferredfrom
sedimentary d
15
N. Because the oxygen content of
the thermocline integrates respiration rates over
multiyearresidencetimeofregionalwatermasses,
windandthermoclinefluctuationsof shorter du-
rations are naturally filtered out of the sediment
record, leaving decadal and longer time scales
(31).ThetrendsinthePacificOMZinferredfrom
our d
15
N records originate from slowly varying
strength of the tropical Walker circulation and
its effect on O
2
demand, rather than changes in
O
2
supply due to gas solubility and thermocline
ventilation. Given the mechanistic link between
equatorial wind stress and the extent of anoxia,
decadaltocentennialvariationsinthesediment
d
15
N may provide a valuable new proxy for var-
iationsinequatorialPacificwindsandupwelling
atthosetimescales.
Our findings demonstrate that the contempo-
rary anoxic zone of the ETNP is currently not
largerthanithasbeeninthepast150years,and
thus cast doubt on the view that the recent ex-
pansionofthetropicalPacificOMZisareflection
ofglobal oceandeoxygenation driven by climate
warming (8). This expansion coincides with a pe-
riod of surface cooling (fig. S8) (32)andthermo-
cline shoaling in the eastern tropical Pacific that
runscountertotheprevailingclimate-warming
trendandpartlyaccountsforthehiatusinglobal
surfacewarming(33).Itisthereforealikelyman-
ifestation of theocean’s pervasivelow-frequency
variability,ratherthanaresponsetorisinggreen-
housegases.Instead,thedominantinfluenceof
anthropogenicclimatewarmingontropicalPacific
oxygen over the last 150 years has been the
weakening of the easterly trade winds and its
effect on respiratory O
2
demand. The resumption
SCIENCE sciencemag.org 8 AUGUST 2014 VOL 345 ISSUE 6197 667
Fig. 3.Centennial changes in
PacificOMZdenitrification
and its climatic forcing.
(A) Sedimentary d
15
Nfrom
coring sites, extended back to
1850. (B) Trends in depth of the
thermocline (13°C isotherm) in
the ETNP from a database of
historical temperature profiles
(27) since 1955 (bold line)
and from IPCC coupled climate
model runs (gray line).The
mean deepening across
19 individual model runs
represents the response to
historical radiative forcing,
predominantly from greenhouse
gases.The standard deviation
reflects the intermodel spread
as well as unforced variability,
which is not in phase between
models and observations.
(C) East-west difference in
sea-level pressure (DSLP)
across the equatorial Pacific
(5°N to 5°S) based on historical
reconstructions (HadSLP,
yellow line) and atmospheric
reanalysis (ECMWF ERA40,
cyan line) (39, 40). Black line is
10-year running mean of the average ofthe data sets.The DSLPis a measure oftradewind strength,
with larger values indicating stronger easterly surface winds.
RESEARCH | REPORTS
190
ofPacifictradewindslackeningpredictedunder
futureclimatewarmingshouldextendthe20th-
centurycontractionoftheOMZintothiscentury.
This wind-driven forcing may eventually be
overwhelmed by the stratification-driven deoxy-
genation of the ocean as a whole, as proxies of
anoxia from Pleistocene sediments point to a
larger tropical OMZ and greater N loss during
warm climates (19, 34, 35).Therelativeinfluence
oftheseeffectsandthetimescalesoverwhichthey
operateonthetropicalOMZremainunknown.
The predominant 20th-centurycontraction of
the North Pacific OMZ has important implica-
tionsforthebasin’sNcycle.Overcentennialtime
scales,theslowingpaceofNlosswouldhavere-
ducedtheNO
3
–
deficitrelativetoplanktonPO
4
3–
requirements throughout surface waters of the
N-limitedNorthPacific.Recentisotopicanalysis
of skeleton material from deep-sea corals near
Hawaii also exhibit a decreasing trend over this
timeperiod,whichhasbeeninterpretedasasig-
nal of increasing N inputsfromN
2
fixation (36).
However,becauseisotopicandstoichiometricsig-
nalsofdenitrificationaretransportedfromthe
anoxic zone into the subtropical gyre (37), the
reported coral trends may originate partly from
the OMZ. Any remaining signal attributable to
N
2
fixationwouldimplythattheecologicalniche
of diazotrophs in the central gyre is uncoupled
fromthemajorNlossintheOMZ(38),andthat
asubstantial imbalanceofthePacificNbudget
has persistedoverthe20thcentury.
REFERENCES AND NOTES
1. R. Vaquer-Sunyer, C. M. Duarte, Proc. Natl. Acad. Sci. U.S.A.
105, 15452–15457 (2008).
2. L. Bopp, C. Le Quere, M. Heimann, A. C. Manning, P. Monfray,
Global Biogeochem. Cycles 16,6-1–6-13 (2002).
3. R. J. Matear, A. C. Hirst, Global Biogeochem. Cycles 17, 1125
(2003).
4. F. Chan et al., Science 319, 920 (2008).
5. R. J. Diaz, J. Environ. Qual. 30, 275–281 (2001).
6. J. A. Koslow, J. Couture, Nature 502, 163–164 (2013).
7. R. E. Keeling, A. Körtzinger, N. Gruber, Annu. Rev. Mar. Sci. 2,
199–229 (2010).
8. L. Stramma, G. C. Johnson, J. Sprintall, V. Mohrholz, Science
320, 655–658 (2008).
9. L. Bopp et al., Biogeosciences 10, 6225–6245
(2013).
10. C. Deutsch, H. Brix, T. Ito, H. Frenzel, L. Thompson, Science
333, 336–339 (2011).
11. T. Ito, C. Deutsch, Global Biogeochem. Cycles 27, 1119–1128
(2013).
12. S. McClatchie, R. Goericke, R. Cosgrove, G. Auad, R. Vetter,
Geophys. Res. Lett. 37, L19602 (2010).
13. L. A. Codispoti, J. P. Christensen, Mar. Chem. 16, 277–300
(1985).
14. J. A. Brandes, A. H. Devol, T. Yoshinari, D. A. Jayakumar,
S. W. A. Naqvi, Limnol. Oceanogr. 43, 1680–1689 (1998).
15. K. Kritee et al., Geochim. Cosmochim. Acta 92, 243–259
(2012).
16. M. A. Altabet, J. Geophys. Res. 94, 12771 (1989).
17. D. M. Sigman, K. L. Casciotti, in Encyclopedia of Ocean
Sciences, J. H. Steele, K. K. Turekian, S. A. Thorpe, Eds.
(Academic Press, London, 2001), vol. 4, pp. 1884–1894.
18. R. C. Thunell, D. M. Sigman, F. Muller-Karger, Y. Astor,
R. Varela, Global Biogeochem. Cycles 18, GB3001 (2004).
19. R. S. Ganeshram, T. F. Pedersen, S. E. Calvert, G. W. McNeill,
M. R. Fontugne, Paleoceanography 15, 361–376 (2000).
20. M. A. Altabet, M. J. Higginson, D. W. Murray, Nature 415,
159–162 (2002).
21. See supplementary materials on Science Online.
22. C. G. Castro, F. P. Chavez, C. A. Collins, Global Biogeochem.
Cycles 15, 819–830 (2001).
23. S. S. Kienast, S. E. Calvert, T. F. Pedersen, Paleoceanography
17,7-1–7-17 (2002).
24. M. F. Lehmann, S. M. Bernasconi, A. Barbieri, J. A. McKenzie,
Geochim. Cosmochim. Acta 66, 3573–3584 (2002).
25. M. G. Prokopenko et al., Earth Planet. Sci. Lett. 242, 186–204
(2006).
26. T. DeVries, C. Deutsch, F. Primeau, B. X. Chang, A. H. Devol,
Nat. Geosci. 5, 547–550 (2012).
27. J. I. Antonov, et al., in NOAA Atlas NESDros. Inf. Serv. 66
(U.S. Government Printing Office, Washington, DC, 2010).
28. G. A. Vecchi et al., Nature 441,73–76 (2006).
29. H. Tokinaga, S.-P. Xie, C. Deser, Y. Kosaka, Y. M. Okumura,
Nature 491, 439–443 (2012).
30. M. A. Merrifield, J. Clim. 24, 4126–4138 (2011).
31. T. Ito, C. Deutsch, Geophys. Res. Lett. 37, L03601 (2010).
32. T. M. Smith, R. W. Reynolds, T. C. Peterson, J. Lawrimore,
J. Clim. 21, 2283–2296 (2008).
33. M. H. England et al., Nature Clim. Change 4, 222–227
(2014).
34. A. van Geen et al., Paleoceanography 18, 1098 (2003).
35. S. L. Jaccard, E. D. Galbraith, Nat. Geosci. 5, 151–156 (2012).
36. O. A. Sherwood, T. P. Guilderson, F. C. Batista, J. T. Schiff,
M. D. McCarthy, Nature 505,78–81 (2014).
37. D. M. Sigman, P. J. DiFiore, M. P. Hain, C. Deutsch, D. M. Karl,
Geophys. Res. Lett. 36, L08605 (2009).
38. T. Weber, C. Deutsch, Proc. Natl. Acad. Sci. U.S.A. 111,
8741–8746 (2014).
39. R. Allan, T. Ansell, J. Clim. 19, 5816–5842 (2006).
40. S. M. Uppala et al., Q. J. R. Meteorol. Soc. 131, 2961
(2005).
ACKNOWLEDGMENTS
This work was supported by grants from the National Science
Foundation (OCE-0851483 to C.D.; OCE-1242313 to T.I.;
OCE-0727123 to W.B.; OCE-0624777 to J.M.), by the Gordon
and Betty Moore Foundation through Grant GBMF3775 to C.D., and
by the U.S. Geological Survey Coastal and Marine Geology Program
(J.C.). A grant from the Climate Center of Lamont-Doherty Earth
Observatory (LDEO) contributed to the collection and dating of
the Soledad Basin core. This is LDEO contribution number 7812.
SUPPLEMENTARY MATERIALS
www.sciencemag.org/content/345/6197/665/suppl/DC1
Materials and Methods
Figs. S1 to S8
References
17 February 2014; accepted 1 July 2014
10.1126/science.1252332
ARTIFICIAL BRAINS
Amillionspiking-neuronintegrated
circuitwithascalablecommunication
networkandinterface
PaulA.Merolla,
1
*JohnV.Arthur,
1
*RodrigoAlvarez-Icaza,
1
*AndrewS.Cassidy,
1
*
JunSawada,
2
*FilippAkopyan,
1
*BryanL.Jackson,
1
*NabilImam,
3
ChenGuo,
4
YutakaNakamura,
5
BernardBrezzo,
6
IvanVo,
2
StevenK.Esser,
1
RathinakumarAppuswamy,
1
BrianTaba,
1
ArnonAmir,
1
MyronD.Flickner,
1
WilliamP.Risk,
1
RajitManohar,
7
DharmendraS.Modha
1
†
Inspired by the brain’s structure, we have developed an efficient, scalable, and flexible
non–von Neumann architecture that leverages contemporary silicon technology.To
demonstrate,we built a 5.4-billion-transistor chip with 4096 neurosynaptic cores
interconnected via an intrachip network that integrates 1 million programmable spiking
neurons and 256 million configurable synapses. Chips can be tiled in two dimensions via
an interchip communication interface, seamlessly scaling the architecture to a cortexlike
sheet of arbitrary size.The architecture is well suited to many applications that use
complex neural networks in real time, for example, multiobject detection and classification.
With 400-pixel-by-240-pixel video input at 30 frames per second, the chip consumes
63 milliwatts.
A
long-standing dream (1, 2) has been to
harness neuroscientific insights to build a
versatilecomputerthatisefficientinterms
of energy and space, homogeneously scal-
able to large networks of neurons and
synapses, and flexible enough to run complex
behavioral modelsoftheneocortex(3,4)aswell
asnetworksinspiredbyneuralarchitectures(5).
No such computer exists today. The von
Neumann architecture is fundamentally ineffi-
cientandnonscalableforrepresentingmassively
interconnected neural networks (Fig. 1) with re-
specttocomputation,memory,andcommunica-
tion(Fig.1B).Mixedanalog-digitalneuromorphic
approaches have built large-scale systems (6–8)
to emulate neurobiology by using custom com-
putationalelements,forexample,siliconneurons
(9, 10), winner-take-all circuits (11), and sensory
circuits (12). We have found that a multiplexed
digitalimplementationofspikingneuronsismore
efficient than previous designs (supplementary
sectionS3)andenablesone-to-onecorrespondence
betweensoftwareandhardware(supplementary
668 8 AUGUST 2014 VOL 345 ISSUE 6197 sciencemag.org SCIENCE
1
IBM Research–Almaden, 650 Harry Road, San Jose, CA
95120, USA.
2
IBM Research–Austin, 11501 Burnet Road,
Austin, TX 78758, USA.
3
Cornell University, 358 Upson Hall,
Ithaca, NY 14853 USA.
4
IBM Engineering and Technology
Services, San Jose Design Center, 650 Harry Road, San Jose,
CA 95120, USA.
5
IBM Research–Tokyo, Nippon Building Fund
Toyosu Canal Front Building, 5-6-52 Toyosu, Koto-ku Tokyo
135-8511, Japan.
6
IBM T. J. Watson Research Center, 101
Kitchawan Road, Yorktown Heights, NY 10598, USA.
7
Cornell
Tech, 111 Eighth Avenue No. 302, New York, NY 10011, USA.
*These authors contributed equally to this work. †Corresponding
author. E-mail: dmodha@us.ibm.com
RESEARCH | REPORTS
191
Appendix B.
Data from Chapter 2

Data in this appendix includes data related to porosity calculations of both Santa
Monica Basin and Pescadero Slope multicores,
210
Pb activity data from the
Pescadero Slope multicore used for age model construction, δ
15
N
sed
from both
multicores, average yearly salinity measurements from the Santa Monica Basin
from the CalCOFI Survey compared, Δδ
15
N
sed
, the average Pacific Decadal
Oscillation Index for from 1900 to 2010, and a supplemental figure of weight
percent nitrogen compared to weight percent carbon from the Pescadero Slope
multicore.

192

193

194

195

196

197

198

199

200
Appendix C.
Data from Chapter 3 and Chapter 4

The tables below contain varve counts, δ
15
N
sed
, weight percent organic carbon, and
weight percent biogenic silica data from the Pescadero Slope cores PESC-MC1
and PESC-GC3.

201

202

203

204

205

206

207

208

209

210

211

212

213

214

215

216

217

218

219

220

221

222
Chapter 4 Data

The X-ray Fluorescence (XRF) analysis for Ti, K, Ca, Sr, and Ba from the
Pescadero Slope will be made available in an online database due to the large
volume of data (>2500 individual data points for each analysis).

The δ
15
N
sed
and XRF data from the Soledad Basin is available through Alexander
van Geen at Lamont-Doherty Earth Observatory of Columbia University. His
contact information is: avangeen@ldeo.columbia.edu.
223

Abstract (if available)

Abstract Oxygen minimum zones (OMZs), located below highly productive marine regions, are sites of microbially-mediated denitrification and biogeochemical cycling that have global significance. The intensity of OMZs fluctuate naturally, however, the degree of these fluctuations and a comprehensive understanding of the factors that drive these fluctuations on interannual to centennial time scales is lacking. This dissertation investigates variations in the intensity of water column denitrification and the OMZ and the transport of that signal northward through the analysis of nitrogen isotopes in laminated marine sediments at three sites in the eastern tropical north Pacific (ETNP). The first portion of this study investigates the northward transport of the ETNP denitrification signal by the California Undercurrent, and the implications of nitrogen isotopes serving as a tracer of ocean circulation. Comparisons between salinity in the thermocline off Southern California (Santa Monica Basin) and the difference between δ¹⁵Nsed within age equivalent sediments from a southern (Pescadero Slope) and northern (Santa Monica Basin) site verify this concept. Trends in this parameter, Δδ¹⁵Nsed, relate to Pacific Decadal Oscillation (PDO) phase changes between 1900 and 1990. We hypothesize that the decline in Δδ¹⁵Nsed during warm PDO phases is due to a strengthening of the California Undercurrent transporting ¹⁵N-enriched nitrate from the ETNP northward. The deviation from this trend after 1990 suggests recent changes in circulation and/or California Current water nutrient biogeochemistry. A high-resolution (near annual) record of δ¹⁵Nsed from the Pescadero Slope in the Gulf of California is presented in the second portion of this study and suggests that the OMZ fluctuates between maximum (10.5‰) and minimum (8.0‰) δ¹⁵Nsed values which create hard ceilings and floors between which OMZ intensity has varied over the past 1200 years. A comparative analysis of the relationship between δ¹⁵NO₃⁻ and [O₂] in Pescadero and nearby sites suggests that the observed range of δ¹⁵Nsed values is equivalent to a 45 μM fluctuation in O₂ content and that these changes can occur in less than 40 years. Our findings show that the OMZ typically intensifies quickly and contracts gradually

Linked assets

University of Southern California Dissertations and Theses

Conceptually similar

PDF

Anaerobic iron cycling in an oxygen deficient zone

PDF

Diagenesis of C, N, and Si in marine sediments from the Western Tropical North Atlantic and Eastern Subtropical North Pacific: pore water models and sedimentary studies

PDF

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

PDF

Oxygen uptake rates in the thermocline of coastal waters: assessing the role of carbon and inorganic nutrient inputs

PDF

Germanium and silicon isotope geochemistry in terrestrial and marine low-temperature environments

PDF

Recent variability in the hyrdological cycle of tropical Asia from oxygen isotopes of tree celulose

PDF

Modeling deep ocean water and sediment dynamics in the eastern Pacific Ocean using actinium-227 and other naturally occurring radioisotopes

PDF

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

Asset Metadata

Creator Tems, Caitlin Elizabeth (author)

Core Title Tracking fluctuations in the eastern tropical north Pacific oxygen minimum zone: a high-resolution geochemical evaluation of laminated sediments along western North America

School College of Letters, Arts and Sciences

Degree Doctor of Philosophy

Degree Program Geological Sciences

Publication Date 04/08/2016

Defense Date 03/21/2016

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

Tag denitrification,eastern tropical north Pacific,geochemistry,laminated marine sediments,nitrogen isotopes,OAI-PMH Harvest,oxygen minimum zone

Format application/pdf (imt)

Language English

Contributor Electronically uploaded by the author (provenance)

Advisor Berelson, William M. (committee chair), Capone, Douglas G. (committee member), Corsetti, Frank (committee member)

Creator Email caty.tems@gmail.com,tems@usc.edu

Permanent Link (DOI) https://doi.org/10.25549/usctheses-c40-227010

Unique identifier UC11278407

Identifier etd-TemsCaitli-4238.pdf (filename),usctheses-c40-227010 (legacy record id)

Legacy Identifier etd-TemsCaitli-4238.pdf

Dmrecord 227010

Document Type Dissertation

Format application/pdf (imt)

Rights Tems, Caitlin Elizabeth

Type texts

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

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

Repository Name University of Southern California Digital Library

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