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Carbonate dissolution at the seafloor: fluxes and drivers from a novel in situ porewater sampler

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Carbonate dissolution at the seafloor: fluxes and drivers from a novel in situ porewater sampler

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Carbonate dissolution at the seafloor: Fluxes and drivers from a novel in situ porewater sampler

by

Jaclyn Elise Pittman Cetiner

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

August 2023

Copyright 2023 Jaclyn Elise Pittman Cetiner
ii
Acknowledgments
There are many people who have made my educational journey not only possible, but
enjoyable as well, which I feel fortunate to be able to say. Sometimes I feel that I somehow cheated
the system and had too much fun during my PhD; I became a seafaring oceanographer, built some
fun robotic ocean instruments, and learned something new about carbonate dissolution, all while
surrounded by great people.
I often think back to those who first fostered my passion in science and encouraged me to
keep asking questions. My undergraduate advisor, Christopher Harrison at San Diego State
University, brought me into his lab when I didn’t know anything about research, just that I was
excited about studying the Earth through a chemistry lens. My summer internship in Karen
Casciotti’s lab at Stanford University, under the mentorship of Brian Peters, was instrumental in
guiding me toward chemical oceanography, and I’m so glad I learned about the marine nitrogen
cycle for a summer.
When I was first applying to grad school, the number one advice I received was to find an
advisor who you get along with, and I accomplished this when coming to USC to work with Will
Berelson. Will’s wry style sense of humor brightened the lab, which he frequented more than many
PIs, often working side by side on the bench. I sincerely appreciate all the time and effort Will put
into training me to be an inquisitive and challenge-seeking scientist, from the hours spent going
over a presentation to the near instant feedback on the many, many versions of my manuscripts.
The other members of the Berelson group made the lab an even livelier, collaborative
environment. Nick Rollins, who can solve just about any problem in the lab (a surprising number
iii
using only zip ties), was indispensable in training me on carbonate chemistry analyses, as well as
working together to design and deploy the porewater sampler. Thank you to Sijia Dong, who
patiently taught me many things about the intricacies of carbonate chemistry, and thank you to
Abby Lunstrum who taught me how to see beyond the lab and think about how carbonate
chemistry relates to the rest of the world. Thanks to my other lab mates: Rucha Wani, Matt
Quinan, and Emma Johnson, for bringing so much energy into the lab and field.
I would like to thank my dissertation and qualifying exam committee members for their
guided teachings during my studies: Naomi Levine, Josh West, Jess Adkins, and Dave Hutchins. A
special thank you to Doug Hammond, who taught the first class I took at USC and has been
inspiring me with his passion for geochemistry ever since. Doug’s field trips and patient
explanations of rocks to a non-geologist encouraged even more questions, and I enjoyed many
geochemistry chats on the walks to and from the frisbee field (frisbee being the primary source of
my exercise during grad school).
Thank you to the National Science Foundation for funding this work; I hope that I have
managed to convince a number of people that studying carbonate dissolution is worth their tax
dollars. I am very grateful to the captain and crew of the R/V Sally Ride SR2113 cruise, which
allowed me to collect the data presented in this dissertation, as well as the science party who made
the experience unforgettable. Thank you also to the captain and crew of the R/V Yellowfin and
staff at Southern California Marine Institute who were always excited to go out on (yet another)
test cruise for this crazy contraption to suck up porewater from deep-sea sediments.
Thanks to the many other colleagues I have had the pleasure working with and learning
from, especially in the slightly bizarre setting that is a month-long research cruise. It was great fun
iv
working so closely with Holly Barnhart and the rest of the “Bladie Ladies” at sea. Frankie Pavia
has been a good friend and patient explainer of many chemical oceanography lab techniques.
I consider myself so fortunate to have been a part of the USC Earth Sciences Department,
where I instantly made so many friends, starting from the new grad camping trip. Specifically,
Naomi Rodgers, Sharadha Sathiakumar, Rachel Kelly, and Mathilde Wimez have become dear
friends and I can’t wait to visit each other wherever in the world we are. Also thanks to the Earth
Sciences department chair, Frank Corsetti, and staff, especially Karen Young and Vardui Ter-
Simonian, who have graciously helped me in running the Young Researchers Program and who
have supported underprivileged high school students to conduct research into our geology labs.
I am so thankful for the close friendships I have outside of grad school, namely Kiley Sauder
and Alessandra Nerio who have been my longest friends, and Ric DeSantiago, who has kept me
grounded during these five years and before (quite literally, through our many camping
adventures).
Finally, I am so grateful to my family for supporting me in my education, which took longer
than I’m sure we all thought. Thank you to my brother John, who is constantly amazing me with
the amount of physics and chemistry he can teach himself by exploding things in the backyard.
Thank you to my sister Holly, for being my best friend, a compassionate listener, and opening my
mind to so many new perspectives. Thank you to my mom, who has always been my biggest
cheerleader and never misses a chance to tell me how proud she is of me. Lastly, I am so fortunate
to have my husband, Barbaros, by my side. After being the first to tell me what a PhD is, over eight
years ago, Barbaros’ unwavering love, encouragement, and support are what truly carried me
through this journey.
v
Table of Contents
Acknowledgments ........................................................................................................................... ii
List of Tables ................................................................................................................................ viii
List of Figures ................................................................................................................................. ix
Abstract .......................................................................................................................................... xi
Chapter 1: Introduction .................................................................................................................. 1
1.1 The marine carbonate chemistry system ...................................................................... 2
1.2 Porewater collection for carbonate chemistry ............................................................... 4
1.3 Carbonate dissolution in deep-sea sediments ............................................................... 5
1.4 Dissertation outline ....................................................................................................... 6
References ........................................................................................................................... 8
Chapter 2: Novel device to collect deep-sea porewater in situ: A focus on benthic carbonate
chemistry ....................................................................................................................................... 11
Abstract ............................................................................................................................. 11
2.1 Introduction ................................................................................................................ 12
2.2 Materials and procedures ............................................................................................ 15
2.2.1 Overview of in situ porewater sampling ...................................................... 15
2.3 Validation of methods ................................................................................................. 28
2.3.1 Porewater draw depth .................................................................................. 28
2.3.2 Overlying water channeling/initial drawdown ........................................... 29
2.3.3 Flow rate ...................................................................................................... 30
2.4 Field results and discussion ......................................................................................... 31
2.4.1 Heterogeneity ............................................................................................... 32
2.4.2 Sample Depth .............................................................................................. 33
2.4.3 Alkalinity in situ vs. cores ............................................................................. 37
2.4.4 Alkalinity artifact vs. sedimentary CaCO3 ................................................... 38
2.4.5 In situ carbonate dissolution experiments .................................................... 41
2.5 Summary ..................................................................................................................... 42
2.5.1 Comparison to existing methods .................................................................. 43
2.6 Comments and recommendations .............................................................................. 43
2.7 Acknowledgements ..................................................................................................... 44
vi
2.8 References ................................................................................................................... 46
Chapter 3: Carbonate dissolution fluxes in deep-sea sediments as determined from in situ
porewater profiles in a transect across the saturation horizon ...................................................... 50
Abstract ............................................................................................................................. 50
3.1 Introduction ................................................................................................................ 51
3.2 Background ................................................................................................................. 53
3.2.1 Respiration-driven dissolution ..................................................................... 53
3.2.2 Mass balance ................................................................................................ 54
3.2.3 Benthic flux chambers .................................................................................. 54
3.2.4 Microelectrodes ............................................................................................ 55
3.2.5 Other approaches ........................................................................................ 55
3.2.6 Calcium profiles ........................................................................................... 56
3.2.7 Water-side vs. sediment-side ........................................................................ 56
3.3 Study area and methods ............................................................................................. 58
3.3.1 Field site ....................................................................................................... 58
3.3.2 Porewater sampling ...................................................................................... 60
3.3.3 Chemical analyses ........................................................................................ 61
3.4 Modeling, calibration, and uncertainty ...................................................................... 64
3.4.1 Calibrating carbonate system measurements .............................................. 64
3.4.2 Curve fitting ................................................................................................. 65
3.4.3 Flux calculations ........................................................................................... 66
3.4.4 Error propagation ........................................................................................ 67
3.5 Results ......................................................................................................................... 68
3.5.1 Porewater profiles ........................................................................................ 68
3.5.2 Solid phase ................................................................................................... 72
3.6 Discussion .................................................................................................................... 73
3.6.1 Defining flux across the sediment-water interface ....................................... 73
3.6.2 Dissolution flux via alkalinity ....................................................................... 74
3.6.3 Dissolution flux via carbon isotopes ............................................................. 75
3.6.4 Dissolution flux via calcium ......................................................................... 76
3.6.5 Comparison of three approaches ................................................................. 78
3.6.6 Comparisons to existing flux data ................................................................ 79
3.6.7 Respiration-driven dissolution ..................................................................... 80
3.6.8 Porewater omega ......................................................................................... 81
3.6.9 Sedimentary mass balance and paleo-implications ..................................... 84
3.7 Summary ..................................................................................................................... 86
3.8 Acknowledgements ..................................................................................................... 87
3.9 References ................................................................................................................... 89
Chapter 4: Conclusions ................................................................................................................. 95
vii
Appendices .................................................................................................................................... 98
Supplementary Material to Chapter 3 .............................................................................. 98
Ex situ processed porewater data .................................................................................... 104
In situ porewater data ..................................................................................................... 109

viii
List of Tables
Table 2.1. Purchased parts. .......................................................................................................... 24
Table 2.2. Custom-built parts fabricated at USC. ........................................................................ 28
Table 2.3. Stations visited on Cocos Ridge in Nov-Dec. 2021. Sediment characteristics and
SIPR draw rates reported. ............................................................................................................ 33
Table 2.4. Analyses performed on in situ collected porewater. Except for dissolved silicate, all
analyses were run on-board. ......................................................................................................... 35

Table 3.1. Station information of relevant bottom water and sediment properties. .................... 59
Table 3.2. Alkalinity, DIC, and CaCO3 dissolution fluxes in mmol/m2/day. CaCO3
dissolution fluxes based on three approaches: shallow alkalinity flux, stable isotopes of DIC,
and calcium flux. Weighted average and uncertainty of three approaches is reported. 1σ in
parentheses. ................................................................................................................................... 78
Table 3.3. Mass balance of Cocos Ridge surface sediments (0-1 cm), including mass
accumulation rates via Carbon-14, CaCO3 accumulation rate, CaCO3 rain rate (a derived
parameter based on the sum of dissolution and accumulation), and CaCO3 burial efficiency,
assuming a steady state. Rates and fluxes are in g/cm2/kyr. Dissolution fluxes (units
converted from Table 3.2) and subsequent calculations are averages based on the three
approaches to calculating dissolution rate (Table 3.2). ................................................................. 85

Table A.1. Core top (0-1 cm) solid phase properties. PIC 𝛿
13
C and POC 𝛿
13
C values were
used as end-members in Keeling analysis. .................................................................................. 100
Table A.2. Ex situ porewater data from Rhizon-filtered multi-cores. ........................................ 105
Table A.3. In situ porewater data collected from SIPR. ............................................................ 111

ix
List of Figures
Figure 2.1. SIPR assembled on multicorer. Letters correspond to those in Figure 2.3. Not
shown are core tubes attached to every other slot on Spyder, or lead bricks that help drive the
Spyder into the sediment. ............................................................................................................. 17
Figure 2.2. Flow chart of SIPR operations. Timing intervals are programmed into the
computer sent down with the multicorer. ..................................................................................... 19
Figure 2.3. SIPR components. (a) Needles, (b) short blade, (c) long blade, (d) exploded needle
showing Rhizon filter and o-ring, (e) exploded blade filter showing Supor filter with
polypropylene net filter, two o-rings, and six screws holding down window covering, (f) pinch
plate, (g) syringe rack showing 60-mL syringes with springs to pull plungers, (h) foot to record
SWI position, (i) quartz coils in protective box, (j) syringe rack detail showing burn wire
holding down the piano hinge that restrains the syringe plungers, (k) syringe rack detail
showing burn wire released and syringe plungers pulled up by springs. ...................................... 27
Figure 2.4. Mixing experiment in sample coil initially filled with fluoride spiked seawater, with
16 mL sample porewater drawn into coil. The dashed line indicates hypothetical plug flow
with no mixing. Solid circles are data points, and the solid line connects data points. ................ 30
Figure 2.5. SIPR flow rates at Station 2, as seen via camera. A subset of sediment depths is
indicated. ....................................................................................................................................... 32
Figure 2.6. Station 2 dissolved silicate from in situ porewater (blue) and ex situ shipboard-
processed core porewater (yellow), before (a) and after (b) silica depth correction. Arrows
denote CTD bottom water values. Different blue colors and symbols represent individual
blades over two deployments, demonstrating the reproducibility of SIPR devices. Uncertainty
of [Si] measurements is 2%. .......................................................................................................... 37
Figure 2.7. Station 2 porewater alkalinity from in situ vs. ex situ (shipboard- processed cores)
filtered porewater. SIPR points represent all blades and needles over two deployments. Arrow
denotes CTD bottom water value. Error bars are smaller than size of point. .............................. 38
Figure 2.8. Porewater alkalinity lost from shipboard-processed cores relative to in situ
porewater vs. weight % CaCO3 in sediment core samples. Data are binned for every 1–2 cm
interval. ......................................................................................................................................... 39
Figure 2.9. Alkalinity and DIC loss (Δ) in shipboard-processed cores com- pared to in situ
filtered SIPR porewater at all stations. Alkalinity:DIC 2:1 ratio line. Δ DIC uncertainty: 32.5
μmol/kg. Δ Alkalinity uncertainty: 2.8 μmol/kg. ........................................................................ 41
Figure 2.10. Station 5 δ
13
C of DIC. (a) Ambient in situ profile, and (b) ambient profile and
results of in situ calcite dissolution experiments. Dissolution experiments shown (diamonds)
are the aliquot farthest from the porewater–fill water interface. Yellow band indicates
potential range of isotopic exchange signal. Note change in x-axes. Sta. 5 has OLW with
Ωcalcite = 0.89. ................................................................................................................................ 42

Figure 3.1. Map of station locations on the Cocos Ridge. Thick contours denote every 1000
m; thin contours denote every 200 m. .......................................................................................... 58
Figure 3.2. Water column profiles of dissolved oxygen (a), and carbonate saturation state with
respect to calcite, with inset focused near omega = 1 (vertical line) (b). ....................................... 60
Figure 3.3. Cross plot of DIC of water column (“WC” as squares and blue fit) and porewater
(“PW” as circles and red fit), calculated from alkalinity and pH, vs measured DIC. ................... 65
x
Figure 3.4. Porewater alkalinity. Arrows denote bottom water value. Solid black lines denote
best fit; dashed lines are 1σ of fit. Error bars (1σ of duplicates) are smaller than size of point. .... 69
Figure 3.5. Porewater dissolved inorganic carbon (DIC). Arrows denote bottom water value.
Solid black lines denote best fit of measured DIC; dashed lines are 1σ of fit. Error bars (1σ of
duplicates) are size of point. .......................................................................................................... 69
Figure 3.6. Porewater 𝛿
13
C of DIC. Arrows denote bottom water value. Solid black lines
denote best fit; dashed lines are 1σ of fit. Error bars (1σ of duplicates) are size of point. ............. 70
Figure 3.7. Porewater Ca, normalized to Na. Arrows denote bottom water value. Solid black
lines denote best fit; dashed lines are 1σ of fit. Error bars are 1 standard error of five replicates.
....................................................................................................................................................... 71
Figure 3.8. pH of in situ collected porewater. Arrows denote bottom water value. ..................... 71
Figure 3.9. Porewater manganese profiles with interpreted depth of oxygen penetration
shown by arrows. Note difference in x-axes and the very low values of [Mn] at Station 3. ......... 72
Figure 3.10. Solid phase properties: % CaCO3, 𝛿
13
C of CaCO3, % organic C, and porosity. .... 73
Figure 3.11. Porewater [Ca] vs alkalinity. Black points in bottom left are bottom water values.
Solid points are from SWI to “breaking points” in alkalinity profiles: 10, 10, 7, 7 cm for
Stations 1, 5, 2, 3, respectively; hollow points are deeper than breaking points. The change
in slope from shallow to deep points represents where reaction stoichiometry changes. Error
bars are standard error of five replicates in [Ca] space; error bars in alkalinity space are
smaller than size of point. ............................................................................................................. 77
Figure 3.12. Porewater profile of saturation state with respect to calcite, calculated with
alkalinity and pH. Arrows denote bottom water omega to their respective stations. Vertical
lines show Ω = 1 for calcite and aragonite. Error bars are smaller than size of point; 1σ of
duplicate measurements range from 0.001 – 0.003 Ω. ................................................................. 82
Figure 3.13. Porewater [Sr] (μmol/kg) (following normalization to [Na]). Arrows represent
bottom water. Error bars are standard error of five replicates. .................................................... 84

Figure A.1. ICP-OES IAPSO consistency standards across all runs, by which calcium and
strontium samples were normalized. ............................................................................................. 99
Figure A.2. Water column alkalinity and DIC comparing measurements from this study vs
GLODAP data, including 1:1 lines. .............................................................................................. 99
Figure A.3. Keeling plots with intercepts listed for every 1 cm sediment horizon. ..................... 101
Figure A.4. Foraminifera ages determined by Carbon-14. ......................................................... 102
Figure A.5 Multivariate analysis determining the best predictors of dissolution flux. The
predictor variables were water column depth, bottom water Ωcalcite, bottom water oxygen,
oxygen penetration depth, and % CaCO3. Bottom water Ωcalcite and oxygen penetration
depth paired together were statistically the best predictor of dissolution flux (top plot). The
second and third best predictor pairs are also shown in the middle and bottom plots. .............. 103

xi
Abstract
Calcium carbonate plays an unusual role in the global C cycle. Its formation produces CO2,
yet the dissolution of calcium carbonate (CaCO3) in the ocean is one of Earth’s natural mechanisms
for neutralizing high concentrations of atmospheric carbon dioxide (CO2). Despite the importance
of CaCO3 dissolution in long-term climate regulation, there remain large uncertainties in the rate
of seafloor dissolution with respect to depth, water column saturation state, surface productivity,
and sediment composition. Specifically, the role and magnitude of respiration-driven dissolution is
not well understood, in which acid production via aerobic respiration of organic carbon fuels
CaCO3 dissolution. Here, I present results from a novel in situ porewater sampler that I led in
constructing and testing. It is capable of extracting deep-sea porewater with cm-scale resolution,
and my first paper describes its design and provides examples of ex situ sampling artifacts. I then
present a study of CaCO3 dissolution on the Cocos Ridge seafloor in the eastern equatorial Pacific
through in situ porewater measurements. Dissolution fluxes were calculated with three
independent approaches: alkalinity fluxes, stable isotopes of dissolved inorganic carbon (DIC)
combined with DIC fluxes, and dissolved calcium fluxes. Although the saturation state of calcite
thermodynamically controls whether dissolution is favored, these results show dissolution flux is
not correlated with the saturation state of overlying bottom water. Additionally, dissolution is
observed in sediments where the water column is saturated with respect to calcite, indicating a
process other than water column chemistry is responsible for dissolution. I conclude that the driver
of sedimentary dissolution in this region is respiration-driven dissolution. That dissolution occurs
within the sediment column, not only at the sediment-water interface, contradicts a paradigm in
the field (rapid, water-side control) and will alter paleoceanographic proxy signals. These
interpretations underscore the necessity to consider factors other than saturation state, such as
xii
organic matter rain and oxygenation when estimating global CaCO3 budgets and assessing the
oceans’ response to the current pulse in atmospheric CO2. Via aerobic respiration of organic
carbon, sedimentary CaCO3 dissolution has the potential to neutralize anthropogenic CO2 at
larger magnitudes than previously estimated.
1
Chapter 1: Introduction
The dissolution of calcium carbonate (CaCO3) in the ocean is one of Earth’s natural
mechanisms for neutralizing atmospheric greenhouse gases. The burning of fossil fuels, such as
coal, oil, and natural gas, emits carbon dioxide (CO2), the greenhouse gas responsible for the
majority of anthropogenic climate change. As emissions continue to rise, the ocean serves as a
significant sink of atmospheric CO2, absorbing 25-33% of anthropogenic CO2 since the start of
the Industrial Revolution (Sabine et al, 2004). Following the absorption of CO2 into seawater,
carbonic acid (H2CO3) is formed, lowering seawater pH approximately 0.1 units in the last 250
years (Doney et al, 2009), and in turn, lowering the saturation state of mineral carbonate. CO2
absorbed or produced in seawater may encounter the mineral calcium carbonate (CaCO3), which
can result in the following CaCO3 dissolution reaction, where the CO2 is converted to alkaline
bicarbonate (HCO3
-
):
CaCO3 + H2O + CO2 à Ca
2+
+ 2HCO3
-
(1.1)
The source of CO2 in Equation 1.1 can be from the atmosphere or from respiration of organic
carbon (Corg) via oxygen uptake.
Certain marine organisms form shells of CaCO3 that, following the death of the organism,
can either dissolve or become buried in sediments at the seafloor. Through photosynthesis, marine
phytoplankton convert CO2 to organic C and by eventually sinking, carbon arrives at the seafloor
on much shorter timescales than through global ocean mixing. Particulate organic C sinks through
the water column at rates on the order of 100 m/day, with sinking rates increasing with depth
(Berelson 2002). This process increases the rate at which atmospheric carbon is transported to the
deep sea, where it can then be neutralized through carbonate dissolution (Equation 1.1).
2
Equation 1.1 serves as the foundation for how the dissolution and preservation of CaCO3
responds to, or drives, changes in global climate over geologic timescales (Broecker & Peng, 1987).
There is evidence of anthropogenic CO2 already having been buffered by seafloor dissolution in
the North Atlantic (Sulpis et al, 2018), which is the region where the thermohaline circulation
forms sinking water that incorporates atmospheric gases. For marine CaCO3 dissolution (which
largely occurs in the deep sea) to fully neutralize anthropogenic CO2, the CO2 must first make its
way down to the deep sea. Estimates based on global ocean circulation put this atmospheric CO2
penetration to the deep sea on the order of 1000-2000 years (Broecker et al, 1984). Further, models
estimate it will take 6,000-10,000 years for CaCO3 dissolution to neutralize the current pulse in
CO2 we have created (Archer et al, 2009). Given its importance in regulating atmospheric CO2, it
is therefore necessary to understand the fundamentals of carbonate dissolution, including where
and how fast CaCO3 is dissolving in deep-sea sediments, which is the focus of this dissertation.

1.1 The marine carbonate chemistry system
The carbonate chemistry system in the ocean is based on the relative proportions of H2CO3
(dissolved CO2), HCO3
-
, and CO3
-2
. Given the relationship between these three species and [H+],
the pH of seawater is an indicator of which species will be most prevalent. Carbonate ion [CO3
-2
]
is the key parameter in the thermodynamic formulation of the solubility of mineral carbonate.
Mineral solubility can be defined by its saturation state in solution, in this case, seawater. The
saturation state of calcite, for example, is defined as
Ωcalcite = ([Ca
2+
]*[CO3
2-
])/K’sp (1.2)
where K’sp = [Ca
2+
]*[CO3
2-
] at calcite saturation. The saturation state of a mineral will determine
its propensity to dissolve or precipitate. The surface ocean is oversaturated with respect to
carbonate minerals. Organisms such as foraminifera (forams), coccoliths, and pteropods living in
3
surface ocean secrete CaCO3 shells. When they die, soft organic matter remineralizes quickly, and
shells are left to sink through the water column, until they either dissolve in the water column or
land on the seafloor. At the seafloor, they will either dissolve or get buried by further sediment
deposition, where they become part of the paleoceanographic record to be studied by geologists in
present day. The shallowest part of the sediment is characterized by numerous diagenetic reactions
that, in part, determine the fate of CaCO3 in the sediments. The fluxes of dissolved constituents in
deep-sea sediments represent large sinks and sources between sediments and the overlying
seawater.
There are two major polymorphs of CaCO3 that are produced in the ocean as shells, or
tests: calcite and aragonite. Foraminifera are single-celled zooplankton, heterotrophic organisms
that primarily secrete calcite tests and live either near the surface ocean (planktonic forams) or on
and within the sediment (benthic forams). Because foram tests record information about their
environments and calcite does not dissolve as easily as aragonite, forams provide a wealth of
paleoceanographic information, such as global temperatures (e.g., Emiliani, 1955) and
atmospheric CO2 changes (e.g., Broecker & Peng, 1989). Other species that secrete CaCO3 shells
are phytoplanktonic, autotrophic coccolithophores that secrete calcite liths, and corals and
zooplanktonic pteropods that secrete aragonite tests.
Aragonite and calcite, while both composed of CaCO3, have different mineral structures
that greatly affect their propensity for dissolution versus preservation. Much of the surface ocean
is oversaturated with respect to calcite (Ωcalcite > 1), meaning calcite is thermodynamically likely to
precipitate. Going deeper in the ocean, the higher pressure and lower temperature make calcite
more soluble which increases K’sp, and at some point, sinking calcite tests will cross the saturation
horizon, where Ωcalcite in the water column = 1. Below this depth, the saturation horizon (Ωcalcite <
1), calcite is more likely to dissolve. This thermodynamic relationship is straight forward, but the
4
study of chemical oceanography rarely stops there. While it is more likely for carbonate to dissolve
below the saturation horizon, thermodynamics is not the only factor in play; the kinetics of a
reaction is also important. Dissolution of carbonate on the sea floor may be slower than the rate at
which it is being supplied to the seafloor, leading to net preservation in the sediment column. There
is a sedimentary horizon known as the carbonate compensation depth (Murray & Renard, 1891),
also known as the calcite snowline, below which there is little to no carbonate found in the
sediment. One definition of the lysocline, and the definition I personally prefer due to its consistent
use of the word root “-cline” with other “-clines” in the ocean (e.g., thermocline), is the depth range
between the saturation horizon and the carbonate compensation depth (Boudreau, Middelburg,
Luo, 2018). Throughout the lysocline, the dissolution rate of calcite increases sufficiently until the
rate of dissolution is equal to the rate of calcite rain to the seafloor. Biological reactions can inhibit
or promote dissolution above the saturation horizon, particularly in the sediment due to oxidation
of raining organic matter (Emerson & Bender, 1981) (Equation 1.1). When this occurs, the
sedimentary lysocline can extend to depths shallower than the saturation horizon. It is this process
that I aim to shed light on with this dissertation: what happens when calcite hits the seafloor, where
does it dissolve, at what rate does it dissolve, and what controls dissolution?

1.2 Porewater collection for carbonate chemistry
Studying porewater is necessary for understanding early diagenetic reactions but is fraught
with artifacts. Porewater is most commonly collected from sediments ex situ, typically with a
sediment core being brought to the surface before filtered with Rhizons, sectioned for
centrifugation, or squeezed to express porewater. All ex situ porewater collection methods are
subject to the pressure and temperature artifact that affects porewater while it is still in contact with
sediment and travels from the seafloor to the sea surface. Sediment cores are exposed to much
5
lower pressures and often warmer temperatures while traveling up through the water column.
While the temperature artifact can be reversed for some dissolved species by placing the core in a
cold room and bringing it down to bottom water temperatures before extraction (Bischoff & Sayles,
1972), the pressure artifact is often irreversible. As pressure decreases, the saturation state of calcite
increases and this may drive alkalinity and calcium to precipitate as CaCO3 onto existing grains
(Murray et al, 1980), in the reverse of Equation 1.1. This precipitation consumes porewater
alkalinity and DIC, confounding interpretations of dissolution based on these parameters. Several
in situ porewater samplers have been developed over the past decades, such as the harpoon
WHIMP created by Sayles et al, 1973, and used by Sayles et al, 1976, Murray et al, 1980, and
Sayles, 1981; and the in situ whole core squeezer from Sayles & Dickinson, 1991. These devices
provided invaluable information regarding sampling artifacts of porewater collection, as well as
provided the first in situ studies of seafloor carbonate dissolution. However, the coarse depth
resolution of the harpoon is not detailed enough to resolve reactions happening in the top few cm
of sediments, where some of the most diagenetically interesting reactions occur. The whole core
squeezer, on the other hand, may miss reactions that occur below the top few centimeters.

1.3 Carbonate dissolution in deep-sea sediments
Perhaps the earliest and most influential study of dissolution in the deep sea was carried out
by Peterson (1966), in which calcite spheres were suspended in the water column and their mass
loss measured after several months. Those mass loss measurements became the first dissolution
rates of calcite in the ocean and showed a dramatic increase in dissolution rates below 3500 m in
the Pacific Ocean (Peterson 1966). Since then, carbonate dissolution on and within the seafloor
has been the topic of numerous studies, including through mass balances, (e.g., Murray, Leborgne,
Dandonneau, 1997; Balch et al, 2007; Berelson et al 2007), in situ microelectrodes (e.g., Archer,
6
Emerson, Reimers, 1989; Reimers, Jahnke, McCorkle, 1992; Hales, Emerson, Archer 1994),
benthic flux chambers (e.g., Berelson, Hammond, O’Neill, et al, 1990; Berelson, Hammond,
Cutter 1990; Jahnke et al 1997), and global modeling efforts (e.g., Hales, Emerson, and Archer,
1994; Archer & Maier-Reimer, 1994; Archer, Morford, Emerson 2002). Carbonate dissolution is
considered especially significant at the seafloor, as 40-80% of all carbonate that sinks through the
water column dissolves near the sediment-water interface (SWI) (Emerson and Bender, 1981).
Berelson et al assembled a global budget of sinking CaCO3 estimates and found that 0.6 Gt
CaCO3/year sinks through the deep sea, 0.4 Gt CaCO3/year dissolves at the seafloor, and 0.1 Gt
CaCO3/year is subsequently buried (Berelson et al, 2007). This represents a large fraction of
CaCO3 that dissolves at the seafloor and reenters the water column as alkalinity. Subsequent
inverse modeling suggests the export of CaCO3 may be much greater, between 1-2 Gt/yr (Liang
et al. 2023). Budgets like this allow us to see the larger picture of how much CaCO3 is dissolving,
but even with the large amount of effort that has been put into understanding carbonate dissolution
in the ocean, there are still open questions regarding exactly where in the sediment carbonate
dissolves and which controls, if any, are the primary drivers of dissolution, such as oxygen, organic
matter, diffusive boundary layer, and/or calcite saturation state.

1.4 Dissertation outline
In Chapter 2, I present the novel device built to extract porewater in situ from deep-sea
sediments. I show the tests used to validate this method against traditional methods. I also explore
the artifacts associated with traditional methods and how they relate to various environmental
factors, such as % CaCO3 and water depth, demonstrating the importance of in situ porewater
filtration for the generation of accurate porewater data. I also present the first in situ CaCO3
dissolution experiments in deep-sea sediments, performed with isotopically labeled Ca
13
CO3
7
grains. Chapter 2 has been published in Limnology & Oceanography: Methods (Cetiner et al
2023).
In Chapter 3, I investigate seafloor CaCO3 dissolution in the Eastern Equatorial Pacific
utilizing the CaCO3 chemistry parameters collected from the device of the previous chapter. In a
transect spanning the Cocos Ridge, I quantify dissolution fluxes using three approaches to quantify
dissolution: 1) porewater alkalinity gradients, which provide a proxy for dissolution; 2) 𝛿
13
C of
DIC, which provides the end member isotopic value of DIC, and hence partitions the source of
DIC attributable to CaCO3 dissolution; and 3) porewater [Ca] gradients, which serve as a proxy
for carbonate dissolution occurring in sediments. I examine dissolution with respect to water depth,
overlying water calcite saturation state, and overlying water oxygen, to distinguish the extent to
which CaCO3 dissolution is controlled by respiration-driven dissolution. Further, I discuss the
implications of respiration-driven dissolution on paleoceanographic interpretations and the
capacity for CO2 neutralization via seafloor carbonate dissolution on shorter timescales than
previously thought. Chapter 3 will be submitted for publication to Geochimica et Cosmochimica
Acta.
In Chapter 4, I compile the conclusions of my dissertation research.
The Appendices of my dissertation include supplementary material to Chapter 3, as well
as tables of all porewater measurements provided or discussed in Chapters 2 and 3.

8
References
Archer D. E., Morford J. L. and Emerson S. R. (2002) A model of suboxic sedimentary diagenesis
suitable for automatic tuning and gridded global domains. Global Biogeochem Cycles 16, 17-1-17–
21.
Archer D., Eby M., Brovkin V., Ridgwell A., Cao L., Mikolajewicz U., Caldeira K., Matsumoto
K., Munhoven G., Montenegro A. and Tokos K. (2009) Atmospheric Lifetime of Fossil Fuel
Carbon Dioxide. Annu Rev Earth Planet Sci 37, 117–134.
Archer D., Emerson S. and Reimers C. (1989) Dissolution of calcite in deep-sea sediments: pH
and O2 microelectrode results. Geochim Cosmochim Acta 53, 2831–2845.
Archer D. and Maier-Reimer E. (1994) Effect of deep-sea sedimentary calcite preservation on
atmospheric CO2 concentration. Nature 367, 260–263.
Balch W., Drapeau D., Bowler B. and Booth E. (2007) Prediction of pelagic calcification rates using
satellite measurements. Deep Sea Res 2 Top Stud Oceanogr 54, 478–495.
Berelson W. M. (2002) Particle settling rates increase with depth in the ocean. Deep Sea Res 2 Top
Stud Oceanogr 49, 237–251.
Berelson W. M., Balch W. M., Najjar R., Feely R. A., Sabine C. and Lee K. (2007) Relating
estimates of CaCO3 production, export, and dissolution in the water column to measurements
of CaCO3 rain into sediment traps and dissolution on the sea floor:A revised global carbonate
budget. Global Biogeochem Cycles 21, 1–15.
Berelson W. M., Hammond D. E. and Cutter G. A. (1990a) In situ measurements of calcium
carbonate dissolution rates in deep-sea sediments. Geochim Cosmochim Acta 54, 3013–3020.
Berelson W. M., Hammond D. E., O’Neill D., Xu X. m., Chin C. and Zukin J. (1990b) Benthic
fluxes and pore water studies from sediments of the central equatorial north Pacific: Nutrient
diagenesis. Geochim Cosmochim Acta 54, 3001–3012.
Bischoff J. L. and Sayles F. L. (1972) Pore Fluid and Mineralogical Studies of Recent Marine
Sediments: Bauer Depression Region of East Pacific Rise. Journal of Sedimentary Research Vol.
42, 711–724.
Boudreau B. P., Middelburg J. J. and Luo Y. (2018) The role of calcification in carbonate
compensation. Nat Geosci 11, 894–900.
Broecker W., Mix A., Andree M. and Oeschger H. (1984) Radiocarbon measurements on
coexisting benthic and planktic foraminifera shells: potential for reconstructing ocean
ventilation times over the past 20,000 years. Nuclear Instruments and Methods in Physics Research,
331–339.
Broecker W. S. and Peng T. ‐H (1989) The cause of the glacial to interglacial atmospheric CO2
change: A Polar Alkalinity Hypothesis. Global Biogeochem Cycles 3, 215–239.
9
Broecker W. S. and Peng T.-H. (1987) The role of CaCO3 compensation in the glacial to
interglacial atmospheric CO2 change. Global Biogeochem Cycles 1, 15–29.
Cetiner J. E. P., Berelson W. M., Rollins N. E., Barnhart H. A., Liu X., Dong S., Byrne R. H. and
Adkins J. F. (2023) Novel device to collect deep-sea porewater in situ: A focus on benthic
carbonate chemistry. Limnol Oceanogr Methods 21, 82–97.
Doney S. C., Fabry V. J., Feely R. A. and Kleypas J. A. (2009) Ocean Acidification: The Other
CO2 Problem. Ann Rev Mar Sci 1, 169–192.
Emerson S. and Bender M. (1981) Carbon fluxes at the sediment-water interface of the deep-sea:
calcium carbonate preservation. J Mar Res 39, 139–162.
Emiliani C. (1955) Pleistocene Temperatures. Geology 63, 538–578.
Hales B., Emerson S. and Archer D. (1994) Respiration and dissolution in the sediments of the
western North Atlantic: estimates from models of in situ microelectrode measurements of
porewater oxygen and pH. Deep Sea Research Part I: Oceanographic Research Papers 41, 695–719.
Jahnke R. A., Craven D. B., McCorkle D. C. and Reimers C. E. (1997) CaCO3 dissolution in
California continental margin sediments: The influence of organic matter remineralization.
Geochim Cosmochim Acta 61, 3587–3604.
Liang, H., Lunstrum, A. M., Dong, S., Berelson, W. M. and John, S. G. (2023) Constraining
CaCO3 export and dissolution with an ocean alkalinity inverse model. Global Biogeochem Cycles
31.
Murray J. and Renard A. F. (1891) Report on Deep-sea Deposits Based on the Specimens Collected During the
Voyage of HMS Challenger in the Years 1872 to 1876.,
Murray J. W., Emerson S. and Jahnke R. (1980) Carbonate saturation and the effect of pressure
on the alkalinity of interstitial waters from the Guatemala Basin. Geochim Cosmochim Acta 44,
963–972.
Murray J. W., Leborgne R. and Dandonneau Y. (1997) JGOFS studies in the equatorial pacific.
Deep Sea Res 2 Top Stud Oceanogr 44, 1759–1763.
Peterson M. N. A. (1966) Calcite: Rates of dissolution in a vertical profile in the central pacific.
Science (1979) 154, 1542–1544.
Reimers C., Jahnke R. and McCorkle D. (1992) Carbon fluxes and burial rates over the continental
slope and rise off central California with implications for the global carbon cycle. Global
Biogeochem Cycles 6, 199–224.
Sabine C. L., Feely R. A., Gruber N., Key R. M., Lee K., Bullister J. L., Wanninkhof R., Wong
C. S., Wallace D., Tilbrook B., Millero F. J., Peng T.-H., Kozyr A., Ono T. and Rios A. F.
(2004) The Oceanic Sink for Anthropogenic CO2. Science (1979) 305, 367–371.
10
Sayles F. L. (1981) The composition and diagenesis of interstitial solutions-II. Fluxes and diagenesis
at the water-sediment interface in the high latitude North and South Atlantic. Geochim
Cosmochim Acta 45, 1061–1086.
Sayles F. L. and Dickinson W. H. (1991) The ROLAI^2D lander: A benthic lander for the study
of exchange across the sediment-water interface. Deep Sea Research Part A, Oceanographic Research
Papers 38, 505–529.
Sayles F. L., Mangelsdorf P. C., Wilson T. R. S. and Hume D. N. (1976) A sampler for the in situ
collection of marine sedimentary pore waters. Deep-Sea Research 23, 259–264.
Sayles F. L., Wilson T. R. S., Hume D. N. and Mangelsdorf P. C. (1973) In situ Sampler for Marine
Sedimentary Pore Waters: Evidence for Potassium Depletion and Calcium Enrichment. Science
(1979) 181, 154–156.
Sulpis O., Boudreau B. P., Mucci A., Jenkins C., Trossman D. S., Arbic B. K. and Key R. M.
(2018) Current CaCO3 dissolution at the seafloor caused by anthropogenic CO2. Proceedings of
the National Academy of Sciences, 201804250.

11
Chapter 2: Novel device to collect deep-sea porewater in situ:
A focus on benthic carbonate chemistry

This chapter was published in 2023 as:
Jaclyn E. P. Cetiner, William M. Berelson, Nick E. Rollins, Holly A. Barnhart, Xuewu Liu, Sijia
Dong, Robert H. Byrne, Jess F. Adkins (2023), Novel device to collect deep-sea porewater in situ:
A focus on benthic carbonate chemistry. Limnology & Oceanography: Methods, 21: 82-
97. https://doi.org/10.1002/lom3.10530

Abstract
We have designed, built, tested, and deployed a novel device to extract porewater from
deep-sea sediments in situ, constructed to work with a standard multi-corer. Despite the
importance of porewater measurements for numerous applications, many sampling artifacts can
bias data and interpretation during traditional porewater processing from shipboard-processed
cores. A well-documented artifact occurs in deep-sea porewater when carbonate precipitates
during core recovery as a function of temperature and pressure changes, while porewater is in
contact with sediment grains before filtration, thereby lowering porewater alkalinity and dissolved
inorganic carbon (DIC). Here we present a novel device built to obviate these sampling artifacts
by filtering porewater in situ on the seafloor, with a focus near the sediment-water interface on cm-
scale resolution, to obtain accurate porewater profiles. We document 1-10% alkalinity loss in
shipboard-processed sediment cores compared to porewater filtered in situ, at depths of 1600-3200
m. We also show that alkalinity loss is a function of both weight % sedimentary CaCO3 and water
column depth. The average ratio of alkalinity loss to DIC loss in shipboard-processed sediment
12
cores relative to in situ porewater is 2.2, consistent with the signal expected from carbonate
precipitation. In addition to collecting porewater for defining natural profiles, we also conducted
the first in situ dissolution experiments within the sediment column using isotopically labeled
calcite. We present evidence of successful deployments of this device on and adjacent to the Cocos
Ridge in the Eastern Equatorial Pacific across a range of depths and calcite saturation states.

2.1 Introduction
Accurate porewater measurements are necessary for myriad oceanographic and
geochemical applications: modeling fluxes of nutrients and major ions (Berelson et al. 1987; Martin
et al. 1991; Sun et al. 2016; Hou et al. 2019); sediment diagenesis (Froelich et al. 1979; Sayles 1979;
Emerson et al. 1980; Jahnke et al. 1994); paleoclimate reconstructions (Schrag et al. 1996; Higgins
and Schrag 2012; Blättler et al. 2019); constraining model projections of the oceans’ response to
climate perturbations (Archer et al. 1998; Sulpis et al. 2018), among many others. Of particular
interest to our group is an effort to understand benthic carbonate diagenesis, dissolution fluxes
(Jahnke et al. 1997; Martin and Sayles, 2006; Berelson et al. 2007), and address the sediment vs.
water-side control of alkalinity fluxes from carbonate dissolution, i.e., identifying to what extent
dissolution occurs at or below the sediment-water interface (SWI) (Boudreau 2013; Sulpis et al.
2017).
The invasion of anthropogenic CO2 leads to acidified surface water eventually making its
way to the deep ocean. There is evidence in the geologic record that Earth has naturally corrected
for high concentrations of CO2 and returned to a state of equilibrium (Foster and Rohling, 2013)
via reactions on the deep-sea floor that neutralize acid, but this process is thought to occur on the
order of 10,000 years (Archer et al. 2009). Critical outstanding questions remain regarding
precisely how fast dissolved CO2 will react with CaCO3, how this reaction occurs as a function of
13
seawater saturation state, and where in the sediment column dissolution occurs. However, typical
porewater processing, e.g., collecting a sediment core and filtering, squeezing, or centrifuging
sediment to isolate porewater, can result in sampling artifacts associated with pressure and
temperature changes that alter the concentrations of porewater species. Much work remains to be
done studying carbonate dissolution at and below the SWI, in large part because of the challenges
associated with obtaining artifact-free carbonate chemistry porewater profiles.
One outstanding question regarding benthic carbonate dissolution studies relates to the
role of organic carbon respiration and carbonic acid formation as a driver of dissolution (Emerson
and Bender 1981). Whether this process occurs mainly at the SWI or deeper within the sediment
column is difficult to assess without artifact-free, high resolution profiles. While carbonate
dissolution kinetics have recently been well studied in the lab (e.g., Subhas et al. 2015; Dong et al.
2018; Naviaux et al. 2019a) and water column (e.g., Dong et al. 2019; Naviaux et al. 2019b,
summarized by Adkins et al. 2021), accurate porewater profiles and in situ carbonate dissolution
experiments will also help complete our understanding of carbonate dissolution kinetics. Because
of the importance of studying benthic carbonate chemistry and because of the inherent difficulties
of obtaining artifact-free porewater, we developed a novel instrument to collect filtered porewater
in situ from deep-sea sediments and to perform the first in situ benthic dissolution rate experiments.
Conducting measurements in situ has long been recognized as a way to obviate pressure
and temperature artifacts associated with sample return to the surface. The work with in situ pH,
pCO2, and O2 microelectrodes (Reimers et al. 1992; Glud et al. 1994; Hales et al. 1994) provided
breakthrough capabilities, yet they could not probe more than a few cm into the sediment column,
potentially not reaching below the depth of oxygen penetration which is necessary to study the
suite of diagenetic reactions occurring in shallow sediments. The WHIMP harpoon sampler was
developed to compare porewater collected in situ to that collected through traditional methods of
14
porewater extraction, i.e., core recovery and squeezing (Sayles et al. 1973). Sayles et al. (1973)
found that Ca and Mg were depleted, and K enriched in cores processed onboard a ship compared
to in situ collection. This harpoon system was the basis for several foundational in situ porewater
studies (Sayles et al. 1976; Murray et al. 1980; Sayles, 1981), in which ex situ depletions and
enrichments were further documented and attributed to the pressure effect, i.e., the pressure
decrease from seafloor to surface drives changes in saturation states and mineral solubilities. A
whole-core squeezer was modified to operate in situ to capture high resolution porewater in the
top few cm (Sayles and Dickinson 1991). This and subsequent similar devices lead to important
studies on carbonate chemistry in Atlantic Ocean sediments (e.g., Martin et al. 2000). A rapid
equilibration in situ peeper system employed passive diffusion of porewater through filters over
several days and was able to measure spatial variability and heterogeneity due to
bioturbation/irrigation (Aller et al., 1998). We have built on some of these ideas to develop our in
situ porewater sampler. Notably, we have focused on collecting high resolution porewater profiles
from the top few cm below the SWI through 30 cm, as well as conducting the first carbonate
dissolution rate experiments, in situ within deep-sea sediments.
The traditional method of collecting porewater, via sediment core collection and on-board
filtration or centrifugation, allows the porewater to interact with sediment grains as the collected
core travels through the water column to an environment with decreased pressure and often
increased temperature. Temperature changes can cause certain ions to adsorb or desorb to
sediment particles, such as the observed increases of K and Cl and decreases of Mg and Ca
(Bischoff et al. 1970). It was later discovered that if sediment cores are brought down to in situ
temperatures before porewater extraction, the enrichments or depletions can be reversed (Bischoff
and Sayles 1972). However, the pressure artifact cannot be reversed and causes the saturation state
15
of CaCO3 to rise, resulting in carbonate precipitation within the core, lowering both porewater
alkalinity (Murray et al. 1980) and dissolved inorganic carbon (DIC) (Sauvage et al. 2014). CaCO3
is more likely to precipitate onto existing nucleation sites, especially existing carbonate grains,
rather than spontaneously in solution. Therefore, even though isolated porewater will experience
the same pressure and temperature changes as sediment cores, the lack of nucleation sites in the
isolated porewater will hinder CaCO3 precipitation. Sauvage et al. 2014 demonstrated that this
sediment-induced artifact was a function of the time a recovered core spent at low pressure prior
to porewater extraction, but there are many other factors that could affect the magnitude of this
alkalinity artifact, as we show here.
To investigate deep-sea porewater chemistry and conduct CaCO3 dissolution experiments,
we built SIPR (Sampling In situ PorewateR) to accomplish the following goals:
1. Filter deep-sea porewater in situ, thereby avoiding artifacts induced by pressure and
temperature changes during sediment core recovery when water remains in contact with sediment
grains;
2. Capture cm-scale resolution porewater profiles of carbonate chemistry parameters,
among other constituents, specifically focused near the sediment-water interface;
3. Perform isotopically labeled calcite dissolution rate experiments in situ within sediment
porewater.

2.2 Materials and procedures
2.2.1 Overview of in situ porewater sampling: The overarching purpose of SIPR is
to separate porewater from sediments in situ, thereby avoiding recovery artifacts. SIPR was
designed to be attached to a multi-corer, a standard oceanographic tool. As described below, all
16
SIPR components mount onto a multi-corer, preserving the ability to collect cores as well as deploy
SIPR for in situ porewater collection. Any ship that can deploy a multi-corer (4 or 8 unit corer)
can thus deploy SIPR.
Two different SIPR designs enter the sediment as the multi-corer frame hits the seafloor:
the “needles” and the “blades” (Figure 2.1). Blades and needles are mounted on the multi-corer
Spyder, the inner part of the multi-corer which drives the core tubes into the sediment under
normal operation. Shortly after inserting the needles or blades into the sediment, porewater is
drawn through filters from various depth horizons from 0 (SWI) to as deep as 30 cm.
The basic flow of operations is as follows (Figure 2.2): multi-corer is lowered to the seafloor
from a cable attached to the ship; needles and blades are inserted into sediment; syringes are
triggered via computer-controlled burn wire to draw porewater through filters and into a storage
coil; after drawing porewater, a pinch mechanism (also fired by burn wire) prevents further suction
before the device is removed from seafloor; multi-corer is lifted from seafloor and received on-
board where porewater samples are allocated for analyses or stored for at-home measurements.
Time between multi-corer recovery and partitioning of samples is less than five hours. Table 2.1
provides descriptions and part numbers for all purchased parts; Table 2.2 provides descriptions for
custom-made parts.
The Ocean Instruments MC-800 has eight positions available for core attachment. We use
four slots for needles and/or blades and the other four slots for standard multi-corer tubes. Thus,
during one deployment, porewater is collected in situ and cores obtained for ex situ porewater
extraction. Each draw from either blades or needles collects eight samples from separate depths.
Each blade or needle collects porewater from a place on the seafloor that is ~50 cm away from the
17
next blade or needle. Therefore, porewater samples from one deployment represent spatial
variability within about 1.5 m
2
of the seafloor.

Figure 2.1. SIPR assembled on multicorer. Letters correspond to those in Figure 2.3. Not shown
are core tubes attached to every other slot on Spyder, or lead bricks that help drive the Spyder into
the sediment.

Needles: The needles are individual polycarbonate cylinders, 0.25 cm inner diameter, 20
cm long, that hold a Rhizon filter stick inside a pointed tip with 1 cm tall open windows (Figure
2.3a,d). The Rhizon stick is secured inside the needle by an o-ring seal. Eight needles reside on one
Aluminum
pressure case
Quartz glass
sample coils (i)
Needles (a, d) Foot (h)
Short blade (b, e)
Camera
Light
Pinch plate (f)
MC-800
Multi-Corer
Long blade (c, e)
30 cm
Syringes (g, j, k)
Spyder
18
platform, and each needle can be adjusted in height, with a range of 1 cm below SWI to 20 cm
deep. These needles resemble the harpoon used by (Sayles et al. 1976), except each needle draws
only one sample and depth resolution can be focused near the SWI.
Blade: The blade is a SIPR filter system that was designed as a second method to collect
porewater with the added benefit that it can also be used to conduct in situ carbonate dissolution
experiments. The blade is made from Delrin plastic 1 cm thick and 30 cm wide. “Window”
openings are machined into both sides, directly opposite each other, at varying depths to hold
Supor filters (0.45 μm), secured and sealed with two o-rings and a frame held against the filter with
six screws (Figure 2.3b,c,e). Blades come in two lengths, the short blade has windows down to 11
cm, with filters spaced 0.5 cm apart near the SWI; the long blade has filter windows as deep as 30
cm below the SWI. Like the needles, the blade filter window openings are 1 cm tall, thus defining
the vertical resolution per sample. For both the needles and blades, sample windows were spaced
such that porewater drawn would not overlap with adjacent samples. Both needles and blades are
fitted into a retaining plate that 1) connects the blade to the multi-corer, 2) provides some resistance
to over-penetration and 3) prevents or slows overlying water (OLW) from channeling down toward
filter windows.
Carbonate dissolution experiments: In addition to collecting porewater for natural
porewater profiles, we also built the blades to conduct in situ carbonate dissolution experiments
within the sediment column. To conduct a dissolution experiment, the blade filter is replaced by a
heat-sealed filter “sandwich” containing 2 to 10 mg of isotopically labeled Ca
13
CO3 (calcite) grains.
Labeled calcite grains were purchased from Sigma-Aldrich and wet sieved to 18-53 µm using 18.2
MΩ DI water. Porewater is drawn past these grains and, accounting for isotopic exchange, an
enriched δ
13
C DIC signal will appear if dissolution occurs. A model translating δ
13
C ratios into
19
gross and net CaCO3 dissolution rates is the topic of another study (H. A. Barnhart unpubl.). The
actual amount of carbonate dissolved is small enough that DIC and alkalinity are unchanged by
the process, allowing these samples to still be used for defining porewater gradients.

Figure 2.2. Flow chart of SIPR operations. Timing intervals are programmed into the computer
sent down with the multicorer.

Sample coils: Porewater is drawn through the blade or needles, then through a short
piece of PVDF Kynar tubing (<10 cm) into quartz glass coils (Figure 2.3i). Quartz glass (0.5 cm
ID, 0.7 cm OD) was chosen because it does not allow gas diffusion and introduces no artifacts in
DIC, alkalinity, or silica. Tests were conducted in the lab to ensure that these constituents did not
change in seawater incubated for 6-12 hours in the coil. The inner diameter (0.5 cm) of the quartz
coil was chosen to minimize mixing between fill water and porewater during sample draw. Coils
Multi-corer
deployed off ship
Multi-corer lands
on seafloor
Blades and
needles inserted
into sediment
Burn wire triggers
syringes to draw
porewater through
filters
Filtered
porewater stored
in quartz coils
Burn wire triggers
pinch to halt
syringe suction
Burn wire triggers
foot to lock in place
at sediment-water
interface
Multi-corer lifted
from seafloor
Multi-corer
recovered on
deck
2-4 h 0 h
1 hr
0.25 h 2-8 h
0.25 hr
0.5 h 2-4 h
Porewater is
allocated or stored
for various analyses
3-5 h
20
are designed to conserve space and provide some structural support for otherwise fragile tubing.
Each set of quartz coils is secured in a clear polycarbonate box for additional security. The coils
have an internal volume of 30 mL. For samples without labeled calcite grains in filters, i.e., “natural
samples,” the entire 30 mL is removed, homogenized, then partitioned for analyses. For samples
with labeled calcite grains in filters, i.e., “dissolution experiments,” aliquots of 7-8 mLs are removed
in sequence, to capture a time sensitive signal of δ
13
C during dissolution.
Pinch: Once triggered, there is no mechanism to stop suction from the syringe until it hits
a set screw at 55 mL, even if the device is removed from the sediment. Therefore, a pinching
apparatus was designed to prevent suction before the multi-corer is lifted from the seafloor (Figure
2.3f). The pinch is a spring-loaded guillotine apparatus, held up by a burn wire, that pinches eight
short lengths of flexible tubing closed at a programmed time. These short pieces of tubing are part
of the continuous draw stream, so when they are pinched shut, there can be no more suction.
Syringes: The plungers of 60 mL Codan syringes were modified with an eye bolt attached
to springs (Figure 2.3g,j,k). The syringes and springs are mounted into a custom-machined rack.
The rack holds the springs in the extended position, acting to pull the plungers up. The plungers
are held down by a piano hinge that is secured by a burn wire. At a programmed time, the burn
wire releases the hinge and allows the springs to pull up the plungers. This is when suction begins
and porewater is drawn through SIPR filters. The negative pressure created by suction is at most
30 psi (21 dbar), which is negligible in comparison to water pressure at our sites, ranging from
1279-3256 dbar, and therefore does not affect the saturation state of calcite in a meaningful
magnitude. These syringes were chosen because they work well in deep-sea applications, e.g.,
benthic landers (Kononets et al. 2021).
21
Foot: A “foot” was designed to initially estimate the location of the SWI in relation to the
multi-corer frame, and thereby the depth of the SIPR windows. The foot has a rectangular Delrin
plate attached to a geared shaft that is free to slide relative to the multi-corer (Figure 2.3h). This
plate gently rests on the seafloor when the multi-corer lands (confirmed by the camera), then locks
in place via burn wire by the spring-loaded pressure of a locking gear. Once back on deck, the
position of the foot is measured relative to the multi-corer which defines how deep each of the
SIPR windows were relative to the SWI. Two of these devices are attached on opposite sides of the
multi-corer to establish the position of the SWI and determine if the seafloor is sloped.
Tubing and fittings: The tubing that connects SIPR parts is designed to minimize
internal volume, prevent the diffusive loss of CO2, and avoid contamination of measured
constituents. Minimal lengths of PVDF tubing are used between the blade/needles and quartz
coils; PVDF is used between quartz coils and pinch plate; silicon tubing in the pinch plate; and
Tygon between pinch plate and syringes. Luer lock two-way valves and barbs are used to connect
tubing.
Computer and pressure case: An aluminum pressure case (10 cm ID, 47 cm length,
rated to 10,000 m) holds an Arduino computer and 12 1.5-volt D batteries (2 parallel sets of 6
batteries in series for 9 V output). The computer controls a bank of relays (on a circuit board that
we fabricate) that fire burn wires at specified times. A tilt meter on the electronics package records
changes in the x, y, and z dimensions, informing if the ship’s cable ever displaced the multi-corer
while sitting on the seafloor for 6-12 hours. The burn wire program is set prior to deployment,
allowing for the time to get the multi-corer to the seafloor (typically deployed at 40 m/min).
Because the pressure case is the ground for the electrical circuit, it is mounted on the multi-corer
with care to isolate it with rubber sheeting so it does not ground to the frame.
22
Camera: A camera and light are also attached to the multi-corer for several purposes: 1)
observe position of blades and needles on the SWI, 2) record syringe draw rates, and 3) observe
potential disturbance of multi-corer during deployment. The Group B Inc. camera and light are
housed in their own pressure cases that are rated to 2800 m water depth. Therefore, the camera
and light can only be deployed up to this depth, even though the computer pressure case and multi-
corer can be deployed deeper. The camera and light package can be programmed to turn on and
off at specified times and can record up to six hours of video footage. As shown in Figure. 2.1, we
attach the camera and light low on the multicore to capture both syringe draw rate and insertion
of the blades and needles into sediment.
Fill water: The tubing and sample coils must be filled with a liquid prior to deployment
to prevent tubing collapse as a result of deep ocean pressures. This “fill water” must have a
chemical tracer to distinguish between fill water and porewater, as the waters will inevitably mix
during collection. Fill water should have carbonate parameters somewhat similar to the sample,
minimizing error when correcting for mixing. Here we have spiked surface seawater with fluoride
(starting concentration of 1.3 ppm F
-
, spiked to 10 ppm F
-
) and measured its dilution using a
fluoride electrode. The [F
-
] in a mixed sample determines how much fill water is present, and the
other measured constituents can be corrected with a two-endmember mixing equation. As
demonstrated by mixing experiments in the lab at typical flow rates (Figure 2.4), porewater with
<1% fill water is collected beyond the 6 mLs closest to the fill water – sample water interface.
Burn wire: We use burn wires to actuate autonomous deep-sea operations. A single wire
consists of a nylon-coated stainless-steel fishing leader (40 lb test) covered in heat-shrink tubing for
added abrasion resistance. A small section (~2 mm) of the coating is removed at the desired burn
location, exposing the bare wire to seawater. When the batteries in the pressure case are
23
programmed to fire, the resulting current is sufficient to corrode the wire. The circuit is completed
through seawater when electrons leave the steel (producing Fe
+2
) and travel through sea water to
the aluminum pressure case, which serves as the ground. The exposed steel corrodes in
approximately 2 minutes and releases whichever piece it held. Figure 2.3 shows an example of
burn wire holding down syringes (Figure 2.3j) and then, upon corrosion (Figure 2.3k), allowing the
springs to pull up syringe plungers.

24

Table 2.1. Purchased parts.
Part Description Part no. Vendor
Multi-Corer 8 sample tubes with an effective
penetration of > 45 cm
MC-800 Ocean
Instruments
Rhizon filters 5 cm porous part, flat tip, OOD 2.5 mm,
glass fiber strengthener, PE/PVC
tubing 12cm, female luer
19.21.23F Rhizosphere
Supor filters 0.45 μm, 25 mm 60172 Pall
Polypropylene
Net Filter
25 μm, 25 mm PP2502500 Millipore
13
C calcite Sieved to 18-53 μm. Geometric surface
area of 0.064 m
2
/g
SKU
492027
Sigma Aldrich
NaF Sodium fluoride powder 201154-5G Sigma Aldrich
PVDF (Kynar)
tubing
0.0625” ID, 0.125” OD 51105K21 McMaster-
Carr
Tygon tubing 0.125” ID, 0.25” OD F-4040-A Tygon
Silicon tubing 0.125” ID, 0.188” OD 06422-04 Cole-Parmer
Syringes 2.5 cm” ID, 60 mL capacity 628426 Codan
(Sweden)
Syringe
springs
7” long extension spring, 4.165 lbs per
inch
7749N879 McMaster-
Carr
O-ring (inner) Buna-N -019 9452K73 McMaster-
Carr
O-ring (outer) Buna-N -022 9452K76 McMaster-
Carr
Burn wire 7x7 stainless steel nylon-coated wire.
40 lb test
CM49-40T-
A
AFW
Heat-shrink
tubing
25‘ length, 0.06” ID 7856K41 McMaster-
Carr
25
Arduino Uno
computer
Rev3
Micro-controller with Adafruit data
logging shield, mechanical relay shield,
custom circuit board
A000046 Arduino
Tilt meter ADXL335- 5V ready triple-axis
accelerometer to record angle in x, y, z
directions and acceleration
163 Adafruit

Camera and
light package
Benthic underwater camera and light,
rated to 2800 m
N/A Group B Inc.
26

20 cm
11 cm
30 cm
35.5 cm
a b c
15.5 cm
33.5 cm
5 cm 1 cm
1.7 cm
wide
15.5 cm
20 cm
9 cm
16 cm
32 cm
d e f
g h i
1 cm
tall
j k
4 cm
27
Figure 2.3. SIPR components. (a) Needles, (b) short blade, (c) long blade, (d) exploded needle
showing Rhizon filter and o-ring, (e) exploded blade filter showing Supor filter with polypropylene
net filter, two o-rings, and six screws holding down window covering, (f) pinch plate, (g) syringe
rack showing 60-mL syringes with springs to pull plungers, (h) foot to record SWI position, (i)
quartz coils in protective box, (j) syringe rack detail showing burn wire holding down the piano
hinge that restrains the syringe plungers, (k) syringe rack detail showing burn wire released and
syringe plungers pulled up by springs.

Maintaining seafloor position: If the multi-corer is jostled or in any way disturbed by
the ship’s cable during a deployment, the blades and needles will move and compromise sample
integrity. To mitigate potential disturbance, an extra 5-10 m of cable is paid out after the multi-
corer has landed. A Benthos glass float is attached 20 m above the multi-corer to help float the
cable so it does not get tangled on the multi-corer. Additionally, we devised a simple “stiff arm”
consisting of a 2.54 cm diameter 3 m long PVC pipe attached to the cable immediately above the
Benthos float via hose clamps. The float keeps the cable off the multi-corer and this stiff arm keeps
the cable from wrapping around the float. Additionally, dynamic positioning on the R/V Sally
Ride (the ship on which we made our deep-sea deployments) maintained lateral position within
150 m
2
(<15x10 m in latitude and longitude) for up to 12 hours. During deployments described in
this paper, we were subject to 0.5-1.5 m waves and winds 5-20 knots.

28

Table 2.2. Custom-built parts fabricated at USC.
Part Material Description
Needles Polycarbonate
plastic
0.25 cm ID, 20 cm long, 1 cm diameter, tip with 1 cm
tall sample port
Blade Delrin plastic Short blade: 11 cm long, 30 cm wide, 1 cm thick.
Long blade: 30 cm long, 30 cm wide, 1 cm thick
Sample
coils
Quartz glass 0.5 cm ID, 0.7 cm OD. 30 mL capacity
Foot Delrin plastic 9 x 9 x 1.25 cm rectangular plate
Pressure
case
Aluminum 47 cm long cylinder 15 cm OD with walls 2.54 cm
thick, can withstand pressure >4000 m

2.3 Validation of methods
2.3.1 Porewater draw depth
Windows on the blade were spaced such that 60 mL of porewater (the maximum volume
drawn), centered in the window, would not overlap with adjacent windows. By drawing 60 mLs
(even though collecting only the last 30 mL for analyses), and assuming 74% porosity (lowest
porosity at these field sites), the calculated radius of sample collection is 2.7 cm. The blade windows
and needles are spaced 6 cm apart minimum, so overlapping sample volume is not likely. However,
flow through muddy sediment is not likely to be perfectly uniform in all directions and is more
likely to draw from shallower sediments where the porosity is higher. We conducted lab testing to
confirm that SIPR windows draw porewater from the intended sediment horizon. Tests were
conducted in the lab by inserting needles, blades, and Rhizon filters (inserted horizontally to mimic
core Rhizon processing) into a container of muddy sediment. Porewater profiles from both SIPR
methods indicated that needles and blades draw the same water as Rhizons, with the depth horizon
29
centered around the midpoint of the filter. Bottom water entrainment into shallow SIPR windows
will also contaminate samples. Lab testing, again with isotopically spiked overlying water, showed
that when the retaining plate is securely placed on the sediment, bottom water entrainment is
negligible.
2.3.2 Overlying water channeling/initial drawdown
Another concern was that during insertion of the blade or needles, overlying water (OLW)
would be carried down into the sediments or move through channels opened during insertion.
Needle and blade windows have internal volumes of 1 cm
3
and 0.6 cm
3
, respectively, that are
exposed as multi-corer travels through the water column. This volume is filled with OLW before
insertion into sediment. A lab test was conducted by filling a bucket with mud and seawater, spiking
OLW with
13
C-enriched DIC and inserting the blade into the mud. By measuring the
13
C signal
from sample DIC, we determined how much OLW is brought down in this process. As expected,
this volume was at most 1 mL. Thus, we know that in addition to the fill water, there is also an
extra ~1 mL OLW that enters the window prior to sample draw. To obviate this contamination
of our sample, we aim to draw more porewater than the 30 mL volume of the sample coil. That
way, this OLW contamination is flushed out. If less than 30 mLs were drawn, the first 6 mLs of
porewater were discarded to account for mixing of fill water and OLW with porewater (Figure
2.4).
30

Figure 2.4. Mixing experiment in sample coil initially filled with fluoride spiked seawater, with 16
mL sample porewater drawn into coil. The dashed line indicates hypothetical plug flow with no
mixing. Solid circles are data points, and the solid line connects data points.

2.3.3 Flow rate
Knowing the porewater draw rate into the sampler helps us determine deployment timing
and is also important in calculating dissolution rate from the labeled carbonate experiments. Flow
rate is determined by either 1) dividing volume recorded on syringe by suction time, or 2) observing
the speed of syringe plunger movement via a camera. Syringes often draw the maximum volume,
31
so the first method is not always viable, thus reported rates from Stations 2-5 are those viewed from
the camera. Rates from Station 1 are from the first method of determining flow rate, as that site
was deeper than the camera’s depth rating of 3000 m. Both syringe travel and camera yield similar
flow rates, as did lab testing which showed that volume draw vs. time is generally a linear function
after 0.5 hr. An example of flow rates from Station 2, characterized primarily by carbonate clays,
is shown in Figure 2.5. Natural samples have filters on both sides of the blade, whereas dissolution
samples have a filter on only one side and therefore draw more slowly than natural samples.
Obtaining a full draw (>50 mL) takes at least 2 hrs. Average natural sample flow rates are reported
in Table 2.3. Flow rates vary among different sediment types and the rate is often fastest in the
upper 5 cm, but otherwise does not scale with depth. The flow rates we found during deep sea
deployments in muddy sediments (Table 2.3) were comparable to those found during testing off
the San Pedro Shelf (400 m). Clearly, the draw rate in sandy sediments was very fast compared to
muds.
2.4 Field results and discussion
The tests described above validated that SIPR works as intended in the lab, and the
following measurements verified the performance of SIPR in the field. Because of the low volume
collected from SIPR (max 30 mLs in sample coils), all analyses must be run on the smallest amount
possible. Table 2.4 summarizes the analyses run and relevant instrument information. Total
alkalinity was measured using the mvMICA system (X. Liu pers. comm.), able to determine
alkalinity on 1.5 mL of sample with a precision of 2 μM. mvMICA is novel in its low volume
requirement, but it was built upon established methods (Liu et al., 2006; Wang et al., 2007; Liu,
Patsavas, & Byrne 2011; Liu et al., 2013).
32

Figure 2.5. SIPR flow rates at Station 2, as seen via camera. A subset of sediment depths is
indicated.

2.4.1 Heterogeneity
It is possible for heterogeneity in natural sediment systems to drive differences in porewater
profiles not due to any artifact or instrument malfunction. To test the extent that heterogeneity
impacts SIPR profiles, we utilized the adjustable nature of the needles. At a test station in San
Clemente Basin (1975 m water depth), all eight needles were set to the same depth (20 cm). The
needles were laterally spaced 2 cm from one another and occupied an area of 24.4 cm
2
. Aside from
one needle that leaked (leaks are diagnosed by high % fill solution in sample water), the remaining
seven needles contained an average dissolved silicate value of 292.6 μM with a standard deviation
of 7.3 μM. This is the same value within error from blade windows at the same depth. The
variability in [Si] values is only slightly higher than our analytical precision (±2%, one standard
deviation of replicates). These results confirmed that in situ samples drawn from the same depth
yield similar concentration values. It also suggests that, at least in San Clemente Basin, any natural
0
20
40
60
0 1 2 3 4 5 6 7
Volume (mL)
Suction time (hr)
9 cm
8 cm
1 cm
24 cm
9 cm
7 cm
10 cm
34 cm
Natural porewater
13
C dissolution experiments
33
heterogeneity in porewater [Si] must be smoothed out by our sampling or does not exist. However,
we do recognize that heterogeneity will exist in sediments, largely due to infaunal activity, but this
was not observed in San Clemente Basin and would be expected to be less common at deeper sites
further from coastal upwelling. We also recognize that heterogeneity in other constituents, such as
carbonate parameters, may be greater than for dissolved [Si].
Table 2.3. Stations visited on Cocos Ridge in Nov-Dec. 2021. Sediment characteristics and SIPR
draw rates reported.
Station
#
Station
coordinates
Water
Depth
(m)
Sediment type Porosity
range
(%)
Average
SIPR
flow rate
(mL/hr)
Weight
%
CaCO3
range
1 6° 47.08’ N,
88° 15.65’ W
3223 Low carbonate
clay
84 - 94 7 0 – 18
5 6° 35.99’ N,
86° 41.75’ W
2911 Medium
carbonate clay
81 - 93 6 10 – 33
2 5° 57.48’ N,
87° 57.40’ W
2650 Medium
carbonate clay
78 - 88 13 57 – 64
3 5° 10.17’ N,
86° 35.59’ W
1630 High carbonate
clay
74 - 83 70 75 – 81
4 4° 48.72’ N,
88° 37.31’ W
1274 Foraminiferal
sand
74 - 80 3600 91 – 93

2.4.2 Sample Depth
As is true with any porewater profile work, assigning depth to a particular sample is key to
defining the profile shape. We have 3 independent methods of determining depth of the blade and
needle windows:
1. Relative depth: The foot is designed to sit on, and thereby mark, the sediment-water
interface. The foot is a physical method to compare where the blades and needles are in space,
34
relative to the SWI. Upon recovery of the multicore, the location of the feet was measured to assign
temporary depths and to ensure that the blades and needles were inserted into the sediment. There
are two feet on the multi-corer, so their position relative to the frame is extrapolated to the blades
and needles, providing relative depth of SIPR windows between multiple blades and needles.
2. Visual depth: The camera field of vision captures one blade to provide evidence of the
depth of blade penetration. The video captures insertion of a blade into the sediment and shows if
the multi-corer was disturbed during the suction time. The multi-corer was never displaced from
the sediment in all ten deployments. From markings on the blade, the camera provides visual
evidence of the penetration and thus the window depths.
3. Chemical depth: For each deployment, dissolved silicate [Si] profiles were compared
between SIPR blades/needles and shipboard-processed cores filtered with Rhizons (Figure 2.6).
Cores have definite depth assignments due to the relatively gentle insertion of the multi-corer into
sediment and resulting interface preservation. Previous work with [Si] gives us confidence that
shipboard-processed cores reflect accurate porewater gradients near the SWI (McManus et al.
1995, Berelson et al. 1997, Hou et al. 2019).

35
Table 2.4. Analyses performed on in situ collected porewater. Except for dissolved silicate, all
analyses were run on-board.
Analysis Instrument
(reference)
Volume
(mL)
Standard Precision
(one σ of
duplicates)
Lab
Dissolved
Inorganic
Carbon
(DIC)
Picarro Cavity Ring-
Down Spectrometer
(Subhas et al., 2015)
4 Dickson
CO2
Reference
Material
seawater
23 μmol/kg Berelson
(USC)
δ
13
C of
DIC
Picarro Cavity Ring-
Down Spectrometer
(Subhas et al., 2015)
4 (same
water
as DIC)
Optical
calcite
powder
0.15 ‰ Berelson
(USC)
Total
Alkalinity
Minimal Volume
Multiparameter
Inorganic Carbon
Analyzer (mvMICA) (X.
Liu et al. unpubl.)
1.5 Dickson
CO2
Reference
Material
seawater
2 μM Byrne
(USF)
pH mvMICA (X. Liu et al.
unpubl.)
1.5 N/A 0.001 Byrne
(USF)
Dissolved
silicate
Spectrophotometer
(Hou et al., 2019)
1 Artificial
seawater
2% Berelson
(USC)
Fluoride Orion fluoride electrode
(Rix et al., 1976)
1 TISAB 0.1 ppm Adkins
(Caltech)

From three shipboard-processed cores, we generate an average [Si] porewater profile. The
top 0-10 cm of core profiles were fitted with a linear, second, or third order polynomial (best fit).
SIPR blades/needles were then shifted in depth, up or down, to minimize the difference in SIPR
and core [Si], producing a “silica depth.” Each blade/needle set must only have one depth offset
by which all windows are shifted. On average, this [Si] adjustment in depth was < ±1 cm compared
to relative depth. This [Si] depth adjustment of in situ samples does not change the magnitude of
36
the concentration change, only the position relative to the SWI. After the [Si] depth correction,
individual blades and needle sets collapse onto the core [Si] gradient (Figure 2.6b).
We consider the chemical depth assignment to be the most accurate. Relative depth can
be influenced by seafloor topography; visual depth only applies to half the blades or needles that
fit in the field of vision. Most importantly, where we have camera evidence of blade penetration,
the assigned chemical depths agree with visual depths.
We found a difference between in situ and ex situ dissolved [Si] deeper in the core (>10
cm) (Figure 2.6). The in situ dissolved [Si] profiles reach higher values than the ex situ core profiles.
This offset is seen after we adjust core [Si] for a temperature correction (McManus et al. 1995).
We do not have a ready explanation for this offset but have perhaps documented yet another
pressure/temperature/precipitation artifact between in situ and ex situ porewater collection.
Unless stated otherwise, reported sediment depths are in terms of this silica correction.
Additionally, [Si] can diagnosis bottom water entrainment into shallow SIPR windows. If
a sample falls off the trend defined by the majority of data and has a lower [Si] value, i.e., closer to
bottom water, we assume that sample must have had channeling of bottom water into the sample
window, and is therefore removed from further analyses.

150 200 250 300 350 400
Dissolved silicate ( M)
0
10
20
30
40
Sediment depth (cm)
Station 2 - Silica
150 200 250 300 350 400
Dissolved silicate ( M)
0
10
20
30
40
Sediment depth (cm)
Station 2 - Silica
(a) Relative depth assignment (b) Chemical depth assignment
2% uncertainty
150 200 250 300 350 400
Silica ( M)
0
5
10
15
20
25
30
35
40
Sediment depth (cm)
Station 2 - Silica
Blade 2, Drop 1
Blade 3, Drop 1
Blade 4, Drop 1
Needles, Drop 1
Blade 3, Drop 2
Blade 5, Drop 2
Blade 6, Drop 2
Needles, Drop 2
Core Rhizons
150 200 250 300 350 400
Dissolved Silicate ( M)
0
5
10
15
20
25
30
35
40
Sediment depth (cm)
Station 2 - Silica
Blade 2, Drop 1
Blade 3, Drop 1
Blade 4, Drop 1
Needles, Drop 1
Blade 3, Drop 2
Blade 5, Drop 2
Blade 6, Drop 2
Needles, Drop 2
Core Rhizons Shipboard-
processed cores
37
Figure 2.6. Station 2 dissolved silicate from in situ porewater (blue) and ex situ shipboard-processed
core porewater (yellow), before (a) and after (b) silica depth correction. Arrows denote CTD bottom
water values. Different blue colors and symbols represent individual blades over two deployments,
demonstrating the reproducibility of SIPR devices. Uncertainty of [Si] measurements is 2%.

2.4.3 Alkalinity in situ vs. cores
There is a known artifact that occurs when dissolved Ca
2+
and CO3
2-
ions precipitate as
CaCO3 onto existing nucleation sites (e.g., sediment grains) as pressure and temperature change
during core recovery from deep water prior to porewater filtration (Murray et al. 1980; Sauvage
et al. 2014). Due to this artifact, alkalinity is lowered in porewater extracted from cores processed
on-board. Our data confirm this artifact and demonstrate how it impacts the profile gradient and
structure (Figure 2.7). This artifact not only reduces the maximum alkalinity value, but drastically
changes the slope and shape of the profile, especially in the top few cm. Importantly, fluxes
calculated from core profiles would be much lower than those from porewater collected in situ.

2300 2400 2500 2600 2700 2800 2900 3000
Alkalinity ( mol/kg)
0
10
20
30
40
Sediment depth (cm)
Station 2 - Alkalinity
In situ SIPR
Shipboard-processed cores
38
Figure 2.7. Station 2 porewater alkalinity from in situ vs. ex situ (shipboard- processed cores)
filtered porewater. SIPR points represent all blades and needles over two deployments. Arrow
denotes CTD bottom water value. Error bars are smaller than size of point.

2.4.4 Alkalinity artifact vs. sedimentary CaCO3
To investigate potential explanations for this alkalinity sampling artifact, the alkalinity loss
from ex situ shipboard-processed cores relative to in situ porewater was compared to weight %
CaCO3 in sediments (Figure 2.8). This comparison is made more robust by our range in weight %
CaCO3: between 1630 m and 3223 m water depth, we collected sediments that range from 0-80
weight % CaCO3. The magnitude of the artifact was quantified by taking the average of all SIPR
points within 1 cm intervals, taking the average of all core points within 1 cm intervals, then
calculating the % loss between SIPR and core values at each cm horizon. Weight % CaCO3 was
measured by sectioning cores at 1-2 cm intervals, grinding the sediment to a powder, then
acidifying the sediment with 10% H3PO4 in an evacuated container to convert CaCO3 to CO2.
This CO2 was then measured on a Picarro Cavity Ring-Down Spectrometer.
Three potential mechanisms for alkalinity loss were examined in this study: 1) the effect of
weight % CaCO3 (higher weight % CaCO3 has more potential for alkalinity loss by providing more
nucleation sites onto which carbonate can precipitate), 2) the effect of water column depth (based
on the known relationship between pressure, temperature, and the saturation state of carbonate),
and 3) time between core recovery and processing (as found by Sauvage et al. 2014). Sites deeper
than 2600 m (Stations 1, 2, 5) exhibit a trend of more alkalinity loss with increased carbonate
content. Contrary to the pressure artifact definition, Station 1, the deepest site, shows the smallest,
39
albeit non-negligible, alkalinity artifact and it is the core with the lowest % CaCO3. From 3200 –
2600 m, weight % CaCO3 is positively correlated with % alkalinity loss in ex situ processed cores.
However, Station 3 at 1600 m has the highest weight % CaCO3, yet has an average
alkalinity artifact smaller than that at Station 2, where there is less sedimentary carbonate.
Therefore, weight % CaCO3 is not the only factor influencing alkalinity loss in cores; water column
depth may also play a role, as could other sedimentary features and composition.

Figure 2.8. Porewater alkalinity lost from shipboard-processed cores relative to in situ porewater
vs. weight % CaCO3 in sediment core samples. Data are binned for every 1–2 cm interval.

Lastly, we compared time between core recovery and processing to see if this impacted
alkalinity. Sauvage et al. 2014 measured an average loss of 8.9% alkalinity in cores due to CaCO3
precipitation after waiting 3-7 hours to extract porewater, compared to cores processed in <2 hrs.
Our cores were placed in a cold van until they reached bottom water temperature. Three cores
were processed for porewater at each station; the time between core recovery and Rhizon-
0
4
8
12
0 25 50 75 100
% Alkalinity loss
Weight % CaCO
3
Station 1 (3223 m)
Station 2 (2650 m)
Station 3 (1630 m)
Station 5 (2911 m)
0
2
4
6
8
10
12
0 20 40 60 80 100
% Alkalinity Loss
Weight % CaCO
3
Station 1 (3223 m)
Station 2 (2650 m)
Station 3 (1630 m)
Station 5 (2911 m)
0
2
4
6
8
10
12
0 20 40 60 80 100
% Alkalinity Loss
Weight % CaCO
3
Station 1 (3223 m)
Station 2 (2650 m)
Station 3 (1630 m)
Station 5 (2911 m)
0
4
8
12
0 25 50 75 100
% Alkalinity loss
Weight % CaCO
3
Station 1 (3223 m)
Station 2 (2650 m)
Station 3 (1630 m)
Station 5 (2911 m)
40
processing ranged from 8 to 27 hours. There was no systematic difference among the alkalinity
profiles from three cores/station processed at different times. We believe we do not see the same
sampling time artifact as Sauvage et al. 2014 because our shortest wait-time is longer than their
longest wait-time and this effect can only be seen within the first 2 hours after recovery.
Alkalinity loss was also compared to DIC loss in cores relative to SIPR (Figure 2.9). The
premise being if carbonate precipitation were responsible for this artifact, an Alk:DIC (Δ) loss ratio
of 2:1 would be diagnostic of this mechanism. Both alkalinity and DIC losses were calculated by
taking the average of all SIPR points within 1 cm intervals and subtracting the average of all
shipboard-processed core points for the same horizon. While there is significant scatter in the data,
possibly due to the relatively high uncertainty of our DIC measurements (±23 µmol/kg), the
average alkalinity loss/DIC loss is equal to 2.2, consistent with the signal expected from carbonate
precipitation.

0
100
200
300
0 100 200
Alkalinity loss (µmol/kg)
DIC loss (µmol/kg)
2:1 line
41
Figure 2.9. Alkalinity and DIC loss (Δ) in shipboard-processed cores com- pared to in situ filtered
SIPR porewater at all stations. Alkalinity:DIC 2:1 ratio line. Δ DIC uncertainty: 32.5 μmol/kg. Δ
Alkalinity uncertainty: 2.8 μmol/kg.

Assigning mechanistic controls of this artifact is beyond the scope of this work, but we
believe this is the first study to show a large range of this alkalinity artifact as a function of deep-
sea sediment type and depth. Weight % CaCO3 and water column depth are demonstrated as
important factors contributing to artifact magnitude; as others have suggested (Murray et al. 1980,
Sauvage et al. 2014), carbonate precipitation during core recovery is the likely driver.
2.4.5 In situ carbonate dissolution experiments
At all stations, we conducted the first in situ carbonate dissolution rate experiments in
sediments. Isotopically labeled calcite grains were placed in blade windows for porewater to be
drawn past. A measured enrichment of
13
C in DIC comes from two processes: isotopic exchange
at the solid-solution interface and calcite dissolution. The first necessary measurement is to define
the ambient δ
13
C DIC profile (Figure 2.10a) in the absence of any labeled carbonate. As porewater
is drawn past labeled calcite grains, isotopic exchange between the grains and adjacent porewater
can occur independently of dissolution. Such isotopic exchange signals have been documented in
benchtop experiments in supersaturated seawater (Subhas et al. 2015). The relative contributions
of δ
13
C from isotopic exchange and dissolution is currently being studied (H. A. Barnhart unpubl.).
However, based on in situ dissolution experiments in supersaturated porewater, we estimate the
effect of isotopic exchange is 10 – 20 ‰. An example from Station 5 (overlying water Ωcalcite =
0.89) shows significant enrichment relative to ambient δ
13
C beyond the exchange estimate,
indicating a contribution to the δ
13
C signal from net carbonate dissolution (Figure 2.10b). This
42
occurs at the highest magnitude in the top 8 cm, likely where sediment is oxygenated, evidence of
aerobic respiration-driven dissolution (Emerson and Bender, 1981; J. E. P. Cetiner unpubl.).
Dissolved manganese is depleted in these porewaters from the SWI to 8 cm (F. J. Pavia pers.
comm.), indicative of the oxygen penetration depth. These results show that in situ dissolution
experiments are a viable method for quantifying carbonate dissolution rates in sediments and
qualitatively defining depth of dissolution.

Figure 2.10. Station 5 δ
13
C of DIC. (a) Ambient in situ profile, and (b) ambient profile and results
of in situ calcite dissolution experiments. Dissolution experiments shown (diamonds) are the aliquot
farthest from the porewater–fill water interface. Yellow band indicates potential range of isotopic
exchange signal. Note change in x-axes. Sta. 5 has OLW with Ωcalcite = 0.89.

2.5 Summary
SIPR (Sampling In situ PorewateR) was built to filter porewater from deep-sea sediments
in situ, avoid sampling artifacts associated with traditional core recovery methods, and obtain high
quality samples for accurately characterizing porewater carbonate chemistry (DIC, δ
13
C of DIC,
Total Alkalinity, and pH). SIPR was successfully deployed nine times on and adjacent to the Cocos
Ridge in the eastern equatorial Pacific, from 1300-3200 m, filtering porewater from sediment types
0 20 40 60 80 100
13
C
DIC
0
10
20
30
Sediment depth (cm)
Station 5 - C13
-3 -2 -1 0
13
C
DIC
0
10
20
30
Sediment depth (cm)
Station 5 - C13
0.15 ‰ uncertainty
2300 2400 2500 2600 2700 2800 2900 3000 3100
Alkalinity ( mol/kg)
0
5
10
15
20
25
30
35
40
Sediment depth (cm)
Station 2 - Alkalinity
In Situ Blade
Cores
Natural porewater
13
C dissolution experiments
(a) (b)
43
ranging from low to high carbonate clay and foraminiferal sand. In comparing alkalinity profiles
from in situ and ex situ shipboard-processed cores, the ex situ porewater exhibits a significant
reduction in alkalinity, attributed to carbonate precipitation in cores, consistent with the ratio of
alkalinity to DIC loss of 2. This artifact leads to a large difference in the steepness of the gradient
(dC/dz) of in situ vs. ex situ profiles. This difference in gradient slope and shape will have profound
impacts on the calculated fluxes of carbonate parameters from deep-sea sediments. We also
established that this artifact (loss of 1-10% alkalinity) is a consequence of sediment % CaCO3 and
the water depth from which cores were processed ex situ.
2.5.1 Comparison to existing methods
Traditional methods for collecting deep-sea porewater include filtering from cores using
Rhizons, centrifugation, and whole-core squeezers. Compared to these methods, SIPR is more
time-intensive in both preparation and deployment: 20-24 person-hours for cleaning, blade/needle
prep in lab, and deck prep (attaching all parts to multi-corer, setting up electronics), and 6-18 hours
total deployment time. Despite these time constraints, it is clear that in situ porewater collection is
necessary to generate high resolution and accurate porewater profiles for carbonate species.

2.6 Comments and recommendations
SIPR was built with a future application for trace metals in mind, hence every surface that
touches the porewater is metal-free. The multi-corer and parts attaching SIPR to the multi-corer
are metal, but porewater is never in contact with these surfaces. Porewater is filtered through a
Supor filter, pulled through Delrin plastic (blade) or polycarbonate sheathing (needle), drawn
through plastic tubing, and stored in quartz glass sample coils. Sample coils could be fabricated
from plastic for trace metal applications, so long as there is no need to measure gas species. Further
44
testing would need to be conducted to ensure these surfaces remain trace metal clean after acid
washing, for example. Other work has clearly documented the sampling artifact of additional
constituents besides carbonate species, such as O2 (Glud et al., 1994), the nitrogen system (NH4
+
,
NO3
-
, NO2
-
) (Aller et al., 1998), and dissolved organic matter (Hall et al., 2007), all of which would
therefore require in situ filtration to obtain the most accurate porewater profiles, and for which
SIPR could be adapted and deployed. SIPR would also be suitable for anoxic environments, as the
quartz sample coils are gas impermeable and allow for separation between porewater (anoxic) and
fill water (oxic). In addition to deep-sea studies, SIPR has potential for coastal applications as well.
In sandy sediments, it is difficult to collect full cores due to the resistance to core penetration in
sands. We demonstrated that SIPR is capable of collecting in situ porewater from sands up to 35
cm deep at Station 4 (data not shown). SIPR has a water column depth constraint set by the
pressure case, which was built for use with the Alvin submersible and rated to 10,000 m.
In summary, through 3 years of testing, we have developed a unique oceanographic
instrument that will serve to further benthic chemistry research by providing 10’s of mLs of artifact-
free, in situ filtered porewater for carbonate chemistry constituents, among other applications.
Pairing in situ porewater profiles with other in situ techniques (e.g., pH and O2 microprofiling,
benthic chambers, eddy correlation studies) could lead to deeper understanding of carbonate
dissolution on the seafloor.

2.7 Acknowledgements
This work was supported by NSF Ocean Acidification (OCE-1834475). J.E.P.C. thanks the
USC Wrigley Institute for Environmental Studies for PhD funding. H.A.B. thanks the Resnick
Sustainability Institute for PhD funding. The authors would like to thank the editors, as well as
45
David Burdige and an anonymous reviewer for their helpful questions and comments that
improved the manuscript. The authors would like to sincerely thank Kalla Fleger for alkalinity and
pH measurements on board; Matthew Quinan for dissolved silicate measurements; Rucha Wani
and Emma Johnson for their contributions to the SIPR team at sea; and the science party, captain,
and crew of the R/V Sally Ride SR2113 cruise. We also appreciate the discussions and laboratory
support from Abby Lunstrum and assistance with silicate measurements from Doug Hammond.
The authors also thank the Southern California Marine Institute and the captain and crew of the
R/V Yellowfin for allowing us many SIPR test deployments.

46
2.8 References
Adkins, J. F., J. D. Naviaux, A. V. Subhas, S. Dong, and W. M. Berelson. 2021. The Dissolution
Rate of CaCO3 in the Ocean. Annual Review of Marine Science 13: 1–24.
doi:10.1146/annurev-marine-041720-092514
Aller, R. C., P. O. J. Hall, P. D. Rude, and J. Y. Aller. 1998. Biogeochemical heterogeneity and
suboxic diagenesis in hemipelagic sediments of the Panama Basin. Deep Sea Res., Part I 45:
133-165.
Archer, D., H. Kheshgi, and E. Maier-Reimer. 1998. Dynamics of fossil fuel CO2 neutralization
by marine CaCO3. Global Biogeochem. Cycles 12: 259–276. doi:10.1029/98GB00744
Archer, D., M. Eby, V. Brovkin, and others. 2009. Atmospheric Lifetime of Fossil Fuel Carbon
Dioxide. Annu. Rev. of Earth and Planetary Sciences 37: 117–134.
doi:10.1146/annurev.earth.031208.100206
Berelson, W. M., D. E. Hammond, and K. S. Johnson. 1987. Benthic fluxes and the cycling of
biogenic silica and carbon in two southern California borderland basins. Geochim.
Cosmochim. Acta 51: 1345–1363. doi:10.1016/0016-7037(87)90320-6
Berelson, W. M., R. F. Anderson, J. Dymond, and others. 1997. Biogenic budgets of particle rain,
benthic remineralization and sediment accumulation in the equatorial Pacific. Deep Sea Res.,
Part II 44: 2251–2282. doi:10.1016/S0967-0645(97)00030-1
Berelson, W. M., W. M. Balch, R. Najjar, R. A. Feely, C. Sabine, and K. Lee. 2007. Relating
estimates of CaCO3 production, export, and dissolution in the water column to measurements
of CaCO3 rain into sediment traps and dissolution on the sea floor:A revised global carbonate
budget. Global Biogeochem. Cycles 21: 1–15. doi:10.1029/2006GB002803
Bischoff, J. L., R. E. Greer, and A. O. Luistro. 1970. Composition of interstitial waters of marine
sediments: temperature of squeezing effect. Science 167: 1245–1246.
Bischoff, J. L., and F. L. Sayles. 1972. Pore Fluid and Mineralogical Studies of Recent Marine
Sediments: Bauer Depression Region of East Pacific Rise. J. Sediment. Res. 42: 711–724.
doi:10.1306/74d72617-2b21-11d7-8648000102c1865d
Blättler, C. L., J. A. Higgins, and P. K. Swart. 2019. Advected glacial seawater preserved in the
subsurface of the Maldives carbonate edifice. Geochim. Cosmochim. Acta 257: 80–95.
doi:10.1016/j.gca.2019.04.030
Boudreau, B. P. 2013. Carbonate dissolution rates at the deep ocean floor. Geophys. Res. Lett. 40:
744–748. doi:10.1029/2012GL054231
Dong, S., A. V. Subhas, N. E. Rollins, J. D. Naviaux, J. F. Adkins, J. F., and W. M. Berelson. 2018.
A kinetic pressure effect on calcite dissolution in seawater. Geochim. Cosmochim. Acta 238:
411–423. doi:10.1016/j.gca.2018.07.015
47
Dong, S., W. M. Berelson, N. E. Rollins, and others. 2019. Aragonite dissolution kinetics and
calcite/aragonite ratios in sinking and suspended particles in the North Pacific. Earth Planet.
Sci. Lett. 515: 1–12. doi:10.1016/j.epsl.2019.03.016
Emerson, S., R. Jahnke, M. Bender, P. Froelich, G. Klinkhammer, C. Bowser, and G. Setlock.
1980. Early diagenesis in sediments from the eastern equatorial Pacific, I. Pore water nutrient
and carbonate results. Earth Planet. Sci. Lett. 49: 57–80. doi:10.1016/0012-821X(80)90150-8
Emerson, S., and M. Bender. 1981. Carbon fluxes at the sediment-water interface of the deep-sea:
calcium carbonate preservation. J. Mar. Res. 39: 139–162.
Foster, G. L., and E. J. Rohling. 2013. Relationship between sea level and climate forcing by CO2
on geological timescales. Proc. Natl. Acad. Sci. 110: 1209–1214.
doi:10.1073/pnas.1216073110
Froelich, P. N., G. P. Klinkhammer, M. L. Bender and others. 1979. Early oxidation of organic
matter in pelagic sediments of the eastern equatorial Atlantic: suboxic diagenesis. Geochim.
Cosmochim. Acta 43: 1075-1090. doi:10.1016/0016-7037(79)90095-4
Glud, R. N., J. K. Gundersen, B. B. Jørgensen, N. P. Revsbech, and H. D. Schulz. 1994. Diffusive
and total oxygen uptake of deep-sea sediments in the eastern South Atlantic Ocean: in situ and
laboratory measurements. Deep Sea Res., Part I 41: 1767–1788. doi:10.1016/0967-
0637(94)90072-8
Hales, B., S. Emerson, and D. Archer. 1994. Respiration and dissolution in the sediments of the
western North Atlantic: estimates from models of in situ microelectrode measurements of
porewater oxygen and pH. Deep Sea Res., Part I 41: 695–719. doi:10.1016/0967-
0637(94)90050-7
Higgins, J. A., and D. P. Schrag. 2012. Records of Neogene seawater chemistry and diagenesis in
deep-sea carbonate sediments and pore fluids. Earth Planet. Sci. Lett. 357–358: 386–396.
doi:10.1016/j.epsl.2012.08.030
Hou, Y., D. E. Hammond, W. M. Berelson, N. Kemnitz, J. F. Adkins, and A. Lunstrum. 2019.
Spatial patterns of benthic silica flux in the North Pacific reflect upper ocean production. Deep
Sea Res., Part I 148: 25–33. doi:10.1016/j.dsr.2019.04.013
Jahnke, R. A., D. B. Craven, and J. F. Gaillard. 1994. The influence of organic matter diagenesis
on CaCO3 dissolution at the deep-sea floor. Geochim. Cosmochim. Acta 58: 2799–2809.
doi:10.1016/0016-7037(94)90115-5
Jahnke, R. A., D. B. Craven, D. C. McCorkle, and C. E. Reimers. 1997. CaCO3 dissolution in
California continental margin sediments: The influence of organic matter remineralization.
Geochim. Cosmochim. Acta 61: 3587–3604. doi:10.1016/S0016-7037(97)00184-1
Kononets, M., A. Tengberg, M. Nilsson, and others. 2021. In situ incubations with the
Gothenburg benthic chamber landers: Applications and quality control. J. Mar. Sys. 214.
doi:10.1016/j.jmarsys.2020.103475
48
Liu, X., Z. A. Wang, R. H. Byrne, E. A. Kaltenbacher, and R. E. Bernstein. 2006.
Spectrophotometric measurements of pH in-situ: laboratory and field evaluations of
instrumental performance. Environ. Sci. Technol. 40: 5036-5044. doi:10.1021/es0601843
Liu, X., M. C. Patsavas, and R. H. Byrne. 2011. Purification and characterization of meta-Cresol
Purple for spectrophotometric seawater pH measurements. Environ. Sci. Technol. 45: 4862-
4868. doi:10.1021/es200665d
Liu, X., R. H. Byrne, L. Adornato, K. K. Yates, E. Kaltenbacher, X. Ding, and B. Yang. 2013. In
situ spectrophotometric measurement of dissolved inorganic carbon in seawater. Environ. Sci.
Technol. 47: 11106-11114. doi:10.1021/es4014807
Martin, J. B., M. Kastner, and H. Elderfield. 1991. Lithium: sources in pore fluids of Peru slope
sediments and implications for oceanic fluxes. Mar. Geol. 102: 281–292. doi:10.1016/0025-
3227(91)90012-S
Martin, W. R., A. P. McNichol, and D. C. McCorkle. 2000. The radiocarbon age of calcite
dissolving at the sea floor: Estimates from pore water data. Geochim. Cosmochim. Acta 64:
1391-1404. doi:10.1016/S0016-7037(99)00424-X
Martin, W. R., and F. L. Sayles. 2006. Organic matter oxidation in deep-sea sediments:
Distribution in the sediment column and implications for calcite dissolution. Deep Sea Res.,
Part II 53: 771–792. doi:10.1016/j.dsr2.2006.01.017
McManus, J., D. E. Hammond, W. M. Berelson, T. E. Kilgore, D. J. Demaster, O. G. Ragueneau,
and R. W. Collier. 1995. Early diagenesis of biogenic opal: Dissolution rates, kinetics, and
paleoceanographic implications. Deep Sea Res., Part II 42: 871–903. doi:10.1016/0967-
0645(95)00035-O
Murray, J. W., S. Emerson, and R. Jahnke. 1980. Carbonate saturation and the effect of pressure
on the alkalinity of interstitial waters from the Guatemala Basin. Geochim. Cosmochim. Acta
44: 963–972. doi:10.1016/0016-7037(80)90285-9
Naviaux, J. D., A. V. Subhas, N. E. Rollins, S. Dong, W. M. Berelson, and J. F. Adkins. 2019.
Temperature dependence of calcite dissolution kinetics in seawater. Geochim. Cosmochim.
Acta 246: 363–384. doi:10.1016/j.gca.2018.11.037
Naviaux, J. D., A. V. Subhas, S. Dong, N. E. Rollins, X. Liu, R. H. Byrne, W. M. Berelson, J. F.
Adkins. 2019. Calcite dissolution rates in seawater: Lab vs. in-situ measurements and inhibition
by organic matter. Mar. Chem. 215. doi:10.1016/j.marchem.2019.103684
Reimers, C., R. Jahnke, and D. McCorkle. 1992. Carbon fluxes and burial rates over the
continental slope and rise off central California with implications for the global carbon cycle.
Global Biogeochem. Cycles 6: 199–224.
Rix, C. J., A. M. Bond, and J. D. Smith. 1976. Direct determination of fluoride in sea water with
a fluoride selective ion electrode by a method of standard additions. Anal. Chem. 48: 1236-
1239. doi:10.1021/ac50002a044
49
Sauvage, J., A. J. Spivack, R. W. Murray, and S. D’Hondt. 2014. Determination of in situ dissolved
inorganic carbon concentration and alkalinity for marine sedimentary porewater. Chem. Geol.
387: 66–73. doi:10.1016/j.chemgeo.2014.06.010
Sayles, F. L., T. R. S. Wilson, D. N. Hume, and P. C. Mangelsdorf. 1973. In situ Sampler for
Marine Sedimentary Pore Waters: Evidence for Potassium Depletion and Calcium
Enrichment. Science 181: 154–156. doi:10.4213/mzm2562
Sayles, F. L., P. C. Mangelsdorf, T. R. S. Wilson, and D. N. Hume. 1976. A sampler for the in situ
collection of marine sedimentary pore waters. Deep Sea Res. 23: 259–264. doi:10.1016/0011-
7471(76)91331-0
Sayles, F. L. 1979. The composition and diagenesis of interstitial solutions-I. Fluxes across the
seawater-sediment interface in the Atlantic Ocean. Geochim. Cosmochim. Acta 43: 527–545.
doi:10.1016/0016-7037(79)90163-7
Sayles, F. L. 1981. The composition and diagenesis of interstitial solutions-II. Fluxes and diagenesis
at the water-sediment interface in the high latitude North and South Atlantic. Geochim.
Cosmochim. Acta 45: 1061–1086. doi:10.1016/0016-7037(81)90132-0
Sayles, F. L., and W. H. Dickinson. 1991. The ROLAI^2D lander: A benthic lander for the study
of exchange across the sediment-water interface. Deep Sea Res. 38 505–529.
Schrag, D. P., G. Hampt, and D. W. Murray. 1996. Pore fluid constraints on the temperature and
oxygen isotopic composition of the glacial ocean. Science 272: 1930–1932.
doi:10.1126/science.272.5270.1930
Subhas, A. V., N. E. Rollins, W. M. Berelson, S. Dong, J. Erez, and J. F. Adkins. 2015. A novel
determination of calcite dissolution kinetics in seawater. Geochim. Cosmochim. Acta 170 51–
68. doi:10.1016/j.gca.2015.08.011
Sulpis, O., C. Lix, A. Mucci, and B. P. Boudreau. 2017. Calcite dissolution kinetics at the sediment-
water interface in natural seawater. Mar. Chem. 195: 70–83.
doi:10.1016/j.marchem.2017.06.005
Sulpis, O., B. P. Boudreau, A. Mucci, C. Jenkins, D. S. Trossman, B. K. Arbic, and R. M. Key.
2018. Current CaCO3 dissolution at the seafloor caused by anthropogenic CO2. Proc. Natl.
Acad. Sci. doi:10.1073/PNAS.1804250115
Sun, X., J. Higgins, and A. V. Turchyn. 2016. Diffusive cation fluxes in deep-sea sediments and
insight into the global geochemical cycles of calcium, magnesium, sodium and potassium. Mar.
Geol. 373: 64–77. doi:10.1016/j.margeo.2015.12.011
Wang, Z. A., X. Liu, R. H. Byrne, R. Wanninkhof, R. E. Bernstein, E. A. Kaltenbacher, and J.
Patten. 2007. Anal. Chim. Acta 596: 23-36. doi:10.1016/j.aca.2007.05.048

50
Chapter 3: Carbonate dissolution fluxes in deep-sea sediments
as determined from in situ porewater profiles in a transect
across the saturation horizon

Abstract
Despite the importance of marine CO2 for long-term climate regulation and the numerous
studies of its cycling, there still remain large uncertainties for rates and mechanisms of seafloor
carbonate dissolution, especially with respect to calcite saturation and the impact of metabolic CO2
production. Here, we present results from an in situ porewater sampler deployed at the Cocos
Ridge in the eastern equatorial Pacific, where we studied seafloor carbonate dissolution from 1600
to 3200 m, where bottom waters ranged from Ωcalcite 1.0 to 0.84. We present porewater profiles of
alkalinity, pH, DIC, 𝛿
13
C of DIC, Ωcalcite, manganese, calcium, and strontium from 0-35 cm and
solid phase porosity, % CaCO3, 𝛿
13
C of PIC, % organic C, and 𝛿
13
C of organic C with cm-scale
resolution. We calculate dissolution fluxes using three independent approaches: alkalinity fluxes
based on porewater gradients, 𝛿
13
C of DIC combined with DIC fluxes, and calcium fluxes based
on porewater gradients. The three approaches result in carbonate dissolution fluxes at four sites of
0.04 to 0.10 mmol CaCO3/m
2
/day (± 20-25%). The magnitude of dissolution fluxes is not
correlated with bottom water saturation state (Ωcalcite), bottom water dissolved oxygen, or
sedimentary CaCO3 content. We observe dissolution occurring at all stations, including where
bottom water is saturated with respect to calcite and present evidence that this occurs through
respiration-driven dissolution. At all stations, porewater Ωcalcite decreases below bottom water
values in the top 5-15 cm before increasing toward saturation, providing evidence of respiration-
51
driven dissolution within the sediment column. Using the 𝛿
13
C of DIC, we partition the observed
DIC fluxes and find 20-50% of DIC sourced from CaCO3 dissolution, with the remaining sourced
from organic matter respiration. Our results provide evidence of sediment-side control of
dissolution fluxes in these deep-sea sediments and that net chemical erosion of old (5,000-10,000
years old) and deep carbonate is occurring. We present a sedimentary mass balance, assembled
with mass accumulation rates and dissolution fluxes, and calculate CaCO3 burial efficiencies
between 3 and 75%, correlating with depth from deepest to shallowest station. Via aerobic
respiration of organic carbon, sedimentary CaCO3 dissolution has the potential to neutralize
anthropogenic CO2 at larger magnitudes than previously estimated.

3.1 Introduction
The ocean has absorbed 25-33% of fossil fuel emissions since the start of the Industrial
Revolution (Sabine et al., 2004). This influx of CO
2 has led to a decrease in surface seawater pH
of roughly 0.1 units (Doney et al., 2009). Based on modeling studies, ocean buffering will neutralize
anthropogenic CO
2
on timescales of 10,000 years (Archer et al., 2009). There is already evidence
of anthropogenic CO
2
being buffered by deep-sea carbonate dissolution in some areas of the North
Atlantic (Sulpis et al., 2018), but it will take on the order of 1000-2000 years, i.e., global ocean
mixing (Broecker et al., 1984), for anthropogenic CO
2
to penetrate throughout the deep ocean.
Carbonate (CaCO3) dissolution, particularly in the deep ocean, is a natural mechanism through
which the planet neutralizes atmospheric CO
2
. When anthropogenic CO
2
does arrive at the
seafloor, knowing the rate of CaCO3 dissolution and the parameters controlling dissolution will
help us better understand how the ocean will respond to anthropogenic climate change.
52
Quantifying rates of CaCO3 dissolution has been of interest to many researchers, and
several methodologies have been employed to constrain the marine carbonate budget, such as
benthic landers (e.g., Berelson et al., 1990; Sayles and Dickinson, 1991; Jahnke et al., 1997), in situ
microelectrodes (e.g., Archer et al., 1989; Reimers et al., 1992), mass balances (e.g., Balch et al.,
2007), and modeling efforts (e.g., Broecker and Broecker, 1974; Archer and Maier-Reimer, 1994).
In addition to dissolution rate, the location in the sediment where dissolution occurs is also of
interest. Specifically, whether dissolution occurs primarily in porewater deeper in the sediment
(Archer, 1991) or in the diffusive boundary layer between bottom water and sediments (Boudreau,
2013), is still under debate. While the parameters driving dissolution, such as saturation state,
bottom water oxygen concentration, organic carbon rain, and respiration-driven dissolution
(Emerson and Bender, 1981) have been the topic of many studies, we are still unsure of the relative
importance of these factors and their control on sedimentary dissolution.
Here, we present the first high resolution porewater profiles obtained from porewater
extracted in situ at sites on and adjacent to the Cocos Ridge in the eastern tropical equatorial
Pacific. We report pH, Total Alkalinity, Dissolved Inorganic Carbon (DIC) and 𝛿
13
C of DIC
measurements on in situ collected porewater and interpret these profiles via Fick’s Law to estimate
seafloor carbonate dissolution fluxes using three approaches: alkalinity gradients, 𝛿
13
C stable
isotopes and DIC gradients, and calcium gradients. We also present a budget for CaCO3 rain,
burial, and dissolution on the Cocos Ridge seafloor. We establish that dissolution is occurring in
the top 5-15 cm of the sediments and that respiration-driven dissolution produces net chemical
erosion of carbonate at our deepest site.

53
3.2 Background
Recently, dissolution kinetics have been well studied in bench top experiments and the
water column to help with the parameterization of the dissolution rate law (e.g., Dong et al., 2018;
Naviaux et al., 2019a; Subhas et al., 2015; Adkins et al., 2021). An outstanding question is whether
these dissolution kinetics will apply to dissolution with the sediment (Barnhart et al, in prep). In
order to address this question, porewater carbonate parameters and subsequent net dissolution
fluxes must be well defined.
3.2.1 Respiration-driven dissolution
Respiration-driven dissolution, i.e., the production of acid via aerobic respiration of organic
carbon and the subsequent dissolution of CaCO3 neutralizes CO
2
produced via the following
coupled reactions (Emerson and Bender, 1981):
Corg + O
2
à CO
2
; CO
2
+ H
2
O + CaCO3 à 2HCO
3
-
+ Ca
2+
(3.1)
Martin and Sayles (1996) found that significant dissolution could be occurring in sediments above
the saturation horizon due to metabolic CO
2
production, based on organic carbon oxidation rates
and dissolution rates estimated from in situ whole-core squeezed porewater measurements of [Ca],
alkalinity, and DIC in situ collected porewater. The depth of oxygen penetration may significantly
affect the rate at which respiration-driven dissolution can occur, regardless of the bottom water
saturation state (Martin and Sayles, 2006). Additionally, even if the rain rates of organic carbon
are higher than those of carbonate, carbonate dissolution may not occur at the magnitude
predicted if the organic carbon is oxidized very near to the SWI, thereby allowing bottom water
CO3
-2
to neutralize metabolic CO
2
(Sayles et al., 2001). Whether this process occurs primarily at
the sediment-water interface or within the sediment column is difficult to assess without accurate,
54
high-resolution profiles. Accurate porewater profiles will help complete our understanding of
seafloor carbonate dissolution and the processes responsible for driving dissolution.
3.2.2 Mass balance
There have been many studies on carbonate production, export, remineralization, and
burial that attempt to constrain the marine carbonate system. The Joint Global Ocean Flux Study
(JGOFS) assessed carbon fluxes between the atmosphere, surface ocean, deep ocean, and sediment
in the equatorial Pacific, as summarized by Murray et al. (1997), an important region in which to
study carbon fluxes as it is the largest natural source of CO
2
to the atmosphere (Tans et al., 1990).
Sediment traps (Berelson et al., 2007), calcification rates derived via satellite data (Balch et al.,
2007), CO
2
fluxes from ocean to atmosphere (Murray et al., 1995), and models of excess alkalinity
and water mass ages (Feely et al., 2002) have all been used to constrain various aspects of the global
carbonate budget. Insofar as carbonate accumulation rate represents the difference between rain
to the seafloor and net dissolution, this rather common approach has been applied at various sites.
Here, we use C-14 dated foraminifera to estimate sediment accumulation and combine this with
dissolution flux to constrain rain rate, assuming a mass balance, and derive rain rates of PIC to the
sea floor and carbonate preservation efficiency.
3.2.3 Benthic flux chambers
Much work has been done to determine seafloor carbonate dissolution fluxes by direct
measurements using benthic flux chamber incubations. Near our study site, in situ seafloor
dissolution fluxes were determined by chambers on landers in a transect in the central equatorial
North Pacific (Berelson et al., 1990b). These studies showed that in carbonate-rich sites, at least
60% of seafloor dissolution could be attributed to organic matter oxidation (Berelson et al., 1990a).
Along the equatorial eastern Pacific, benthic chamber derived CaCO3 dissolution fluxes ranged
55
from 0.2-0.7 mmol/m
2
/day; this study also found that organic carbon rain rates combined with
bottom water undersaturation become the driving force of dissolution—these results are based on
modeling fluxes using a single dissolution rate constant and dissolution rate as a continuous
function of saturation state (Berelson et al., 1994). A study by Jahnke et al. found that modeled
rates of carbonate dissolution under-estimated rates measured in chambers, indicating that
metabolic CO
2
must significantly contribute to the observed dissolution rates (Jahnke et al., 1997).
The ROLAI
2
D lander captured fluxes of alkalinity and DIC; these flux ratios were used as a proxy
for carbonate dissolution and respiration (Sayles and Dickinson, 1991).
3.2.4 Microelectrodes
Microelectrodes can provide high resolution profiles that can be modeled for estimates of
CaCO3 dissolution within sediments; the major benefits of microelectrodes include mm-scale depth
precision and in situ measurement capability. Oxygen and pH microelectrodes, in conjunction
with numerical models, have be used to calculate dissolution rates in the top few cm of sediment
(Archer et al., 1989; Reimers et al., 1992), and measure oxygen flux into sediment, a proxy for
aerobic respiration that could drive dissolution (Hales et al., 1994). However, there is uncertainty
in microelectrode pH data interpretation given the multiple reactions that can affect pH. The use
of in situ pCO
2
microelectrodes also provides evidence that a fraction of sedimentary dissolution
occurs through respiration-driven processes (Hales and Emerson, 1997).
3.2.5 Other approaches
The modeling approach to understanding benthic dissolution began with a simple model
to predict the location of the sedimentary lysocline (Broecker and Broecker, 1974), which laid the
foundation for studies regarding the significance of the water column saturation horizon with
respect to benthic dissolution. These models have evolved to include PIC:POC ratios of raining
56
material (Archer and Maier-Reimer, 1994) and also involve formulations of benthic dissolution
kinetics. Hales et al. (1994) have modeled benthic dissolution using a kinetic formulation modified
from that originally proposed by Keir (1980) and require metabolic dissolution to explain fluxes.
An oxic and suboxic diagenetic model was able to successfully reproduce solid phase and porewater
NO3
-
, Mn, Fe, S, and organic C to find that in the deep sea, oxidation by O2 accounted for 95%
of organic C respiration (Archer et al., 2002). Using porewater obtained with an in situ whole-core
squeezer, (Martin et al., 2000) modeled the input of
13
C:
12
C DIC into porewater as a proxy for
determining the amount of DIC that came from carbonate versus from organic matter respiration.
3.2.6 Calcium profiles
The flux of dissolved calcium into a benthic chamber may be considered a more definitive
proxy for CaCO3 dissolution (Jahnke et al., 1997) given the many sources of alkalinity but fewer
major reactions involving dissolved Ca. However, the measurement precision needed to observe
small changes in [Ca] can be difficult to achieve. Measurements of excess [Ca] with respect to Na
in porewater has also been considered a signal of sedimentary dissolution (Green et al., 1998) in
coastal sediments. Calcium to alkalinity ratios have been used as a measure of CaCO3 production
and dissolution (Jahnke and Jahnke, 2004; Steiner et al., 2021) and some studies have successfully
measured porewater [Ca] at high enough precision to interpret porewater profiles in terms of
cation exchange processes (Steiner et al., 2022).
3.2.7 Water-side vs. sediment-side
An ongoing debate regarding seafloor dissolution is whether dissolution is under water-side
or sediment-side control, specifically in the diffusive-boundary layer versus the porewater. Some
studies point to models and sediment bed reactors that indicate seafloor carbonate dissolution is
entirely dependent on boundary layer conditions (e.g., Boudreau, 2013; Sulpis et al., 2017). One
57
major implication of water-side control is the prediction that there is no change in porewater pH
or calcite saturation state (Ωcalcite) downcore (Boudreau et al., 2020). Because porewater has been
so difficult to sample accurately in the past, no artifact-free, high-resolution Ωcalcite porewater
profiles have previously been generated to test this implication.
Here, we present porewater profiles of alkalinity, DIC, pH, 𝛿
13
C of DIC, and [Ca] obtained
from in situ samplers. We present profiles, model gradients, and calculate diffusive fluxes from four
sites on the Cocos Ridge. We use three approaches to determine the magnitude and location of
carbonate dissolution fluxes within the sediment column depth: (1) alkalinity fluxes, (2) 𝛿
13
C of DIC
isotope mass balances, and (3) porewater calcium gradients. We relate dissolution fluxes to bottom
water saturation state and bottom water oxygen and discuss the implications of benthic dissolution
to carbonate mass balances in this region.

58
3.3 Study area and methods
3.3.1 Field site

Figure 3.1. Map of station locations on the Cocos Ridge. Thick contours denote every 1000 m;
thin contours denote every 200 m.

In December 2021, we conducted a cruise on the R/V Sally Ride on and adjacent to the
Cocos Ridge in the Eastern Equatorial Pacific (Figure 3.1). This region contains a wide range in
surface sediment carbonate content, <1-93 % CaCO3 (Moore et al., 1973; Cetiner et al., 2023).
Previous studies assign the sedimentation rate in this region as between 1-3 cm/kyr (Lea et al.,
2000; Liao and Lyle, 2014).
With the inherent challenges of collecting data from the deep seafloor, a common obstacle
is how to overcome sampling and measurement artifacts. To avoid sampling artifacts associated
with pressure and temperature changes through the water column that induce CaCO3
precipitation, we built an in situ porewater sampler (SIPR: Sampling In situ PorewateR) (Cetiner
Depth (m)
Station 1
Station 2
Station 3
Station 5
Costa Rica
4500
0
500
1000
1500
2000
2500
3000
3500
4000
59
et al., 2023), that was deployed at 4 stations at depths ranging from 1630– 3223 m. Station 1 was
located furthest from the ridge and deepest; Stations 5 and 2 were a similar distance from the ridge;
Station 3 was on the ridge (Table 3.1). Stations will be presented in this paper in order of deepest
to shallowest: 1, 5, 2, 3. Depth and distance to ridge scale with weight % CaCO3; stations that
were shallower and closer to the ridge had the highest sedimentary carbonate content.

Table 3.1. Station information of relevant bottom water and sediment properties.
Station #
Station
coordinates
Water
Depth (m)
Bottom water
Omega_calcite
Bottom water
oxygen (uM)
1
6° 47.08’ N
88° 15.65’ W
3223 0.84 105
5
6° 35.99’ N
86° 41.75’ W
2911 0.89 100
2
5° 57.48’ N
87° 57.40’ W
2650 0.94 100
3
5° 10.17’ N
86° 35.59’ W
1630 1.0 75

At each station, a CTD cast measured dissolved oxygen (Figure 3.2a) and Niskin samples
were collected for analyses of DIC, 𝛿
13
C of DIC, alkalinity, and pH. These parameters were all
measured to over-constrain the water column carbonate chemistry system and determine the
saturation state of calcite in bottom waters (Figure 3.2b) using the CO2SYS program.
60

Figure 3.2. Water column profiles of dissolved oxygen (a), and carbonate saturation state with
respect to calcite, with inset focused near omega = 1 (vertical line) (b).

3.3.2 Porewater sampling
Porewater samples were collected using SIPR, a device we built capable of filtering
porewater in situ from deep-sea sediments. The device design, field validation, and comparison to
traditional methods are described in detail in Cetiner et al. (2023). Briefly, two types of filtering
mechanisms, termed the “blades” and “needles,” are attached to a standard multicorer and
deployed off a ship in the traditional manner except that once landed, the multi-corer remains
unmoving on the seafloor for 6-10 h. The blades and needles have “windows” that act as sampling
ports, covered with filter paper (blades) or a Rhizon filter (needles). The blades and needles
penetrate the sediment and sampling ports are located between 0-35 cm, with high resolution (0.5
cm spacing) near the sediment-water interface. Spring-loaded syringes draw porewater through
the filters, where the samples are stored in 30 mL glass quartz coils designed to prevent gas loss
and contamination. Filtered surface seawater “fill water” with F
-
spike establishes mixing between
012345
calcite
0
1000
2000
3000
Water depth (m)
CDISP Water Column Alk/pH Omega
0 100 200
Oxygen ( mol/kg)
0
1000
2000
3000
Water depth (m)
CDISP Water Column Oxygen
0.9 1 1.1
calcite
1000
2000
3000
Water depth (m)
CDISP Water Column Alk/pH Omega
(a) (b)
Stn. 3
Stn. 2
Stn. 5
Stn. 1
Ω calcite = 1
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
61
porewater and fill water as a sample is drawn. In this paper, we report porewater samples that had
<1% fill water; corrections to measured parameters are therefore minimal. Sample volume drawn
was 30-60 mL. Before the multi-corer is retrieved, another mechanism pinches the tubing to
prevent further water draw during recovery. To assess penetration depth, a plate resting on the
sediment is locked in place to provide an estimation of the sediment-water interface relative to the
blades and needles. A camera also records the placement of the samplers and records the sample
draw rate. Once on board the ship, porewater samples are analyzed for pH, alkalinity and DIC
within 4 hours of recovery. Each blade or set of needles collects 8 samples. They are spaced around
a multicorer such that one set of samples could be 1 m distant from another set. The blades and
needles collected porewater from slightly different depths, so here we have combined samples from
all blades and needles, from two deployments per station, to create full porewater profiles from 0
to 35 cm at each station. Total in situ porewater samples recovered were 53, 54, 60, and 59, for
Stations 1, 5, 2, 3, respectively.
3.3.3 Chemical analyses
Dissolved inorganic carbon (DIC) and 𝛿
13
C of DIC were analyzed with a Picarro Cavity
Ring-Down Spectrometer with Liaison autosampler; the method for which is described in Subhas
et al. (2015). These measurements were made on-board ship, and standards included Dickson
CO2-reference seawater and pre-weighed calcite powder. Exetainer vials were pre-weighed and
pre-acidified in the lab prior to the cruise. After the vials were filled with 3-5 mL of pore water and
run on the cruise, the stored vials were weighed again in the lab to get the sample mass. Results
using this methodology were corrected using the Dickson seawater to standardize. Uncertainty (1σ)
for duplicate DIC and 𝛿
13
C are ±23 µmol/kg and ±0.15‰, respectively.
62
Total alkalinity and pH were measured with the mvMICA system, developed by the group
at U. South Florida and described in detail (X. Liu in prep., based on methods described in Liu et
al., 2006, 2011), The mvMICA is uniquely suitable for SIPR samples due to its small volume
requirement and high precision. The system has one channel each for alkalinity and pH. For
alkalinity, 1.5 mL was delivered to a semi-micro disposable cuvette, covered, and stored in a water
bath until it reached 20°C. For pH, 1.5 mL was directly collected in gas tight syringes made of
borosilicate glass and a triple seal plunger; 3 µL of 2 mM pure mCP dye was injected into the
sample through a thin Teflon needle. The samples were thermostated in a custom-made cell
warmer for the temperature to reach 20°C (Liu et al., 2006, 2011, 2013). pH values and profiles
reported here are calculated at in situ pressures and temperatures using CO2SYS. Uncertainty
(1σ) for duplicate alkalinity and pH measurements were ±2 µM and ±0.001, respectively.
Porewater dissolved manganese [Mn
2+
] was analyzed by inductively coupled plasma mass
spectrometry (ICP-MS) using an Agilent 8800 triple-quadrupole instrument. In-situ porewater
samples for metal analysis were collected in acid-cleaned HDPE bottles and were acidified using
6M distilled hydrochloric acid. After at least one week, 1 mL aliquots of porewater were diluted
10:1 in distilled 5% nitric acid for ICP-MS analysis. Concentrations of [Mn] were determined via
calibration to multi-element commercial ICP-MS standards matrix matched to samples by
addition to artificial seawater, diluted in 5% nitric acid identically to samples. Blanks were assessed
by analysis of both pure 5% nitric acid, and artificial seawater with no metal standard added. ICP-
MS measurements were made in MS/MS mode using helium as a carrier gas in the collision cell.
Detection limit was 0.027 μmol/kg [Mn].
Porewater calcium [Ca] and strontium [Sr] were measured at Northwestern University on
a Thermo iCap7600 ICP-OES (Inductively Coupled Plasma-Optical Emission Spectroscopy). In
order to capture signals from high to low abundance cations ([Na], [Sr], respectively), samples were
63
diluted 1:100, by mass, with metal-free 3% HNO3
-
. Each sample was weighed, diluted, and
measured in five replicates. Samples were run in a randomized order to account for any
instrumental drift not corrected through standards. Cation concentrations were based on counts
per second of Ca393.366, Ca396.847, Sr407.771, Sr421.552, Na588.995, and Na589.592 nm
wavelengths. The two wavelengths for each element produced virtually identical results, but were
averaged nonetheless. Measurements were taken in sample-standard bracketing. Calibration
standards were synthetic mixtures of 1000 mM cations ([Mg], [Sr], [Na], [Ca], [K]) diluted in
HNO3, matched to expected sample concentrations. Calibration standards were used to correct
for instrumental drift within a single run, by interpolating between standards run at the beginning
and end of the run. Blank measurements were made on HNO3
-
and were included in the
calibration curve to account for instrument blank. IAPSO seawater was used as a consistency
standard to account for drift across all runs. Between 3-8 IAPSO samples were measured every
run; the IAPSO average in a single run was normalized to the IAPSO average for all runs
(Appendix Figure 1). This normalization was applied to all porewater samples. In this manner,
samples across all runs were normalized to a single measured IAPSO average as a consistency
check. [Ca] and [Sr] were normalized to Na, via CTD-measured bottom water salinity. All stations
had salinity of 34.6 ppt; salinity was converted to Na using seawater Na:salinity of 10.781 g/kg:35
ppt (Pilson, 1998).
Multicores obtained during SIPR deployment provided solid phase sediment, which were
sampled for porosity, PIC, Total Carbon (TC), mineralogy, 14C analyses and uranium-series
isotopes. Porosity was obtained by weight of water loss upon drying, assuming water salinity (34.6
ppt) and sediment grain density (2.5 g/cm
3
). Bulk sediment samples were dried in an oven at 50°C
for 72 h, and then ground to be homogeneous for TC and mineralogy analysis. Approximately 2
– 10 mg (depending on the expected % TC) of the samples were used for PIC (% CaCO3),
64
analyzed on the Picarro following acidification. TC and d
13
C measurements were conducted on
an Elemental Analyzer (EA, Costech) coupled to a Picarro Cavity Ring-Down Spectrometer
(G2131-i). Mineralogy of the sediment samples was measured with a Bruker D8 Advance X-Ray
Diffraction (XRD). About 200 mg of each ground sediment sample was mounted onto a rotating
disc and scanned for ~1 h in XRD. Data analysis was conducted using the software Jade.
Radiocarbon values were measured by using accelerator mass spectrometry (AMS) at the
University of California, Irvine, (UCI) Keck Carbon Cycle Accelerator Mass Spectrometry
(KCCAMS) laboratory. The sediment was dried, washed over a 63 μm sieve with DI water, and
pelagic forams were selected from the >250 um fraction at different sediment horizons. They were
then leached with 10% HCl to remove any post-deposition carbonate and hydrolyzed under
vacuum using H3PO4. The CO2 released from this hydrolysis was then graphitized and counted
on the AMS.

3.4 Modeling, calibration, and uncertainty
3.4.1 Calibrating carbonate system measurements
We over-constrained the carbonate system by measuring pH, total alkalinity, and dissolved
inorganic carbon (DIC), providing an opportunity to compare different pairs and assess the overall
accuracy of our measurements. We also compared our water column alkalinity measurements to
those in the GLODAP database (Appendix Figure 2) and we see strong agreement. Water column
alkalinity and pH measurements defined water column DIC via CO2SYS (MATLAB v3.1.1)
(Lewis and Wallace, 1998), using the total pH scale and K1, K2 constants from Hansson (1972,
1973) & Mehrbach et al. (1973) refit by Dickson and Millero (1987). Comparing this calculated
DIC to our measured DIC, we found our measurements were 1.0051 times higher, so we applied
65
this correction to our measured values. We then compared our measured DIC in porewater (after
this correction) to calculated DIC (from porewater alkalinity and pH) and found strong agreement
(Figure 3.3). Hereafter, we show measured DIC (with this correction) in profiles and calculations
because our sampling procedure provided more porewater DIC than pH measurements.

Figure 3.3. Cross plot of DIC of water column (“WC” as squares and blue fit) and porewater (“PW”
as circles and red fit), calculated from alkalinity and pH, vs measured DIC.

3.4.2 Curve fitting
Alkalinity, DIC, 𝛿
13
C of DIC, and [Ca] pore water profiles were fit with an exponential
function in the following form (McManus et al., 1995):
Cz = Cd – (Cd – C0)exp(-β*z) (3.2)
where Cz is the fit concentration (μmol/kg) at each depth; Cd is the concentration when z
approaches infinity, i.e., where the profile approaches an asymptote; C0 is the bottom water
concentration, β is a fitting coefficient (cm
-1
), and z is depth in sediment (cm). Values of C0 come
1800 2000 2200 2400 2600 2800 3000
Measured DIC ( mol/kg)
1800
2000
2200
2400
2600
2800
3000
Calculated DIC ( mol/kg)
WC slope = 0.99
PW slope = 1.03
Water Column
WC Linear Fit
Porewater
PW Linear Fit
1:1 Line
66
from samples collected from Niskins triggered within 10 m of the bottom and assumed to be the
known and fixed bottom water value, from which profiles are “anchored” at the SWI. Blade and
needle samples collected from above the SWI confirmed the Niskin-obtained values represented
water concentrations very close to the interface.
Although Equation 3.2 was initially formatted for porewater dissolved silicate (Hurd, 1972;
McManus et al., 1995), the similarity in profile shapes of silicate and carbonate species, specifically
the strong gradient in the shallow sediments followed by an approach toward an asymptote,
suggests that this equation is also appropriate for porewater carbonate species. It is a fitting
equation, not intended to imply any mechanism. Other fitting equations were also applied to our
data, but there was no improvement in R
2
, so we use Eq. 3.2.
This curve fit was used to establish the gradient 𝜕C/𝜕z at the SWI and other depths. Two
adjustable parameters, β and Cd, provide adequate degrees of freedom to fit data with R
2
>0.92.
The following profiles were fit with this method: alkalinity, DIC, 𝛿
13
C of DIC, and calcium (Figures
3.4, 3.5, 3.6, 3.7).
3.4.3 Flux calculations
The concave-down curvature in the top few cm of each profile indicates reaction is
occurring in the shallow sediment column. These profile gradients were converted into fluxes via
Fick’s First Law:
Fz = -D0*𝜙
2.6
*𝜕C/𝜕z (3.3)
where F is flux (μmol cm
-2
s
-1
), D is the diffusivity coefficient of HCO3
-
(the predominant species in
alkalinity and DIC) ~ 5*10
-6
cm
2
s
-1
(exact value dependent on in situ bottom water temperature)
(Boudreau, 1997), 𝜙 is porosity raised to 2.6 to account for tortuosity and porosity in diffusive
transport (McManus et al., 1995), and 𝜕C/𝜕z is the change in concentration over change in depth
67
evaluated at z (μmol kg
-1
cm
-1
). By using Fick’s First Law, we are assuming transport is dominated
by diffusion, rather than by advection, which is an acceptable assumption in slowly accumulating
deep-sea sediments with few benthic infauna (McManus et al., 1995; Berelson et al., 2005).
3.4.4 Error propagation
Error was propagated through flux calculations using two statistical methods:
bootstrapping and Monte Carlo iterations. Because we combined data from different sampling
devices and from two deployments, we wanted to test the variability in profile fit. To do this, the
profiles were bootstrapped, meaning a random subsample (20% of the data for each station) was
fit to Equation 2 using a non-linear least squares fit. By using this approach, we are considering
that all porewater data represent the best approximation of the average seafloor conditions. The
random subsampling was repeated over 1000 iterations; the range of fits did not change when
more iterations were used. Each of the 1000 fits resulted in Cd and β values. The overall average
curve was calculated using the average of Cd and β; the overall uncertainty of the curve was
calculated by combining the uncertainties (1σ) for Cd and β. Also included in the curve fitting was
a depth uncertainty term. SIPR depths are determined by matching dissolved silica profiles from
shipboard-processed sediment cores to in situ porewater, as well as with video footage (see Cetiner
et al. (2023) for full description of depth assignment). Some error in depth assignment is possible,
so the depth term was allowed to vary ± 1 cm in the curve fitting, normally distributed, as 1 cm
was the average offset from the silica correction.
Once the profiles were fit, a Monte Carlo error propagation was used to get final
uncertainty on flux calculations. Monte Carlo error propagation involves randomly assigning error
to each variable, using its normally distributed standard deviation over 10,000 iterations, to get a
single uncertainty from an equation that incorporates uncertainties on each term (Equation 3.4).
68
Error was propagated through Equation 3.2 by including the uncertainty of porosity (σ𝜙 = 1%)
and ∂C/∂z. Error on ∂C/∂z comes from propagating the error through Equation 3.3 and including
the uncertainty of Cd and β that were obtained through bootstrapping. In this approach we assume
no uncertainty in diffusion coefficient. The uncertainty on flux can be propagated as follows:
σFlux = D0* 𝜙 (±σ)
2.6
* ∂C/∂z(±σ) (3.4)
Fluxes and their respective uncertainties are reported in Table 3.2. Estimation of
dissolution flux was derived using three approaches, described below. Included in Table 3.2 is the
weighted average and weighted standard deviation, respectively defined as:
x̅weighted = Σwixi / Σwi

(3.5)
σweighted = 1 / √(Σwi) (3.6)
where wi = 1/σi
2
(Taylor, 1997).

3.5 Results
3.5.1 Porewater profiles
All alkalinity (Figure 3.4) and DIC (Figure 3.5) porewater profiles increase from bottom
water values toward what appears to be a near-asymptotic value deeper in the sediment. When
sample depth resolution is very high in the top few cm (namely Stations 2 and 3), values in the top
0.5 cm approach the bottom water value. The profile curvature changes significantly between 0-
10 cm, where the gradient is much steeper; ∂C/∂z has a much smaller slope below 10 cm. 𝛿
13
C of
DIC profiles also exhibit steeper curvature in the top 0-10 cm, before approaching a near-
asymptotic value below 10 cm (Figure 3.6). Bottom water for all stations starts near 0 ‰, and
porewater DIC becomes lighter with depth, -1.5 to -2.5 ‰.
69

Figure 3.4. Porewater alkalinity. Arrows denote bottom water value. Solid black lines denote best
fit; dashed lines are 1σ of fit. Error bars (1σ of duplicates) are smaller than size of point.

Figure 3.5. Porewater dissolved inorganic carbon (DIC). Arrows denote bottom water value. Solid
black lines denote best fit of measured DIC; dashed lines are 1σ of fit. Error bars (1σ of duplicates)
are size of point.

2400 2600 2800 3000
Alkalinity ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 5
2400 2600 2800 3000
Alkalinity ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 3
2400 2600 2800 3000
Alkalinity ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 2
2400 2600 2800 3000
Alkalinity ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 1
Water depth:

3223 m
Water depth:

2911 m
Water depth:

2650 m
Water depth:

1630 m
2300 2500 2700 2900
DIC ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 5
2300 2500 2700 2900
DIC ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 1
2300 2500 2700 2900
DIC ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 2
2300 2500 2700 2900
DIC ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 3
Water depth:

3223 m
Water depth:

2911 m
Water depth:

2650 m
Water depth:

1630 m
70

Figure 3.6. Porewater 𝛿
13
C of DIC. Arrows denote bottom water value. Solid black lines denote
best fit; dashed lines are 1σ of fit. Error bars (1σ of duplicates) are size of point.

[Ca] was measured on porewater from all stations, resulting in the profiles in Figure 3.7
showing a general increase of 0.1-0.2 mmol/kg from the bottom water value with depth in
sediment and a curvature similar to that found in DIC and alkalinity. Because of the necessity of
measuring five replicates per sample and the water volume required to make this many [Ca]
measurements, there are fewer data defining these profiles.
-2 -1 0
13
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 5
-2 -1 0
13
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 3
-2 -1 0
13
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 2
-2 -1 0
13
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 1
Water depth:

3223 m
Water depth:

2911 m
Water depth:

2650 m
Water depth:

1630 m
71

Figure 3.7. Porewater Ca, normalized to Na. Arrows denote bottom water value. Solid black lines
denote best fit; dashed lines are 1σ of fit. Error bars are 1 standard error of five replicates.

Figure 3.8. pH of in situ collected porewater. Arrows denote bottom water value.

At all stations, pH decreases in porewater below overlying water values (Figure 3.8),
indicating a process that lowers pH is occurring in the shallow sediments. Profiles of pH were not
fit with a curve because an interpolated pH profile is not needed in further calculations.
10.2 10.4 10.6
Calcium (mmol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 5
10.2 10.4 10.6
Calcium (mmol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 3
10.2 10.4 10.6
Calcium (mmol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 2
10.2 10.4 10.6
Calcium (mmol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 1
Water depth:

3223 m
Water depth:

2911 m
Water depth:

2650 m
Water depth:

1630 m
7.4 7.5 7.6 7.7 7.8
pH
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 5
7.4 7.5 7.6 7.7 7.8
pH
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 3
7.4 7.5 7.6 7.7 7.8
pH
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 2
7.4 7.5 7.6 7.7 7.8
pH
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 1
Water depth:

3223 m
Water depth:

2911 m
Water depth:

2650 m
Water depth:

1630 m
72
Dissolved manganese [Mn] concentration is below detection (Figure 3.9) at z=0 and
remains low below the SWI until a depth where it increases substantially. The depth at which [Mn]
increases is interpreted as the depth of oxygen penetration (Froelich et al., 1979; Burdige, 1993;
Chong et al., 2018), based on the sequence of redox reactions that occur in sediments.

Figure 3.9. Porewater manganese profiles with interpreted depth of oxygen penetration shown by
arrows. Note difference in x-axes and the very low values of [Mn] at Station 3.

3.5.2 Solid phase
Porosity profiles (Figure 3.10) show canonical decreases in porosity with depth, although
there are some step-changes in the profiles. At Station 1, the step-wise porosity decrease coincides
with the % CaCO3 increase. Nevertheless, these step-changes occur deeper than 10 cm and
therefore do not affect our SWI flux results. Average 0-1 cm porosity was 94, 93, 88, and 83 % for
Stations 1, 5, 2, and 3, respectively. Particulate Inorganic Carbon (PIC), Total Carbon (TC), and
their 𝛿
13
C stable isotopes were measured through depth of core (Figure 3.10). % CaCO3 (PIC)
012345
Manganese ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 3
0 50 100 150
Manganese ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 5
0 5 10 15 20
Manganese ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 2
0 20 40 60 80
Manganese ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
Station 1
Water depth:

3223 m
Water depth:

2911 m
Water depth:

2650 m
Water depth:

1630 m
73
increases downcore at Stations 1 and 5 but remain consistent with depth at Stations 2 and 3.
Organic C content was higher at the deeper stations (0.7-1.5%) than at the shallower two stations
(0.2-0.5 %). 𝛿
13
C of the carbonate ranged from 0.4-1.1 ‰.

Figure 3.10. Solid phase properties: % CaCO3, 𝛿
13
C of CaCO3, % organic C, and porosity.

3.6 Discussion
3.6.1 Defining flux across the sediment-water interface
Various diagenetic reactions occurring with organic carbon will produce DIC and
alkalinity, including CaCO3 dissolution. In an effort to simplify modeling efforts, we define the
74
aerobic diagenetic zone as the upper 5 cm at each station. In this zone, we assume only the
following two reactions are occurring: CaCO3 dissolution and organic C respiration via oxygen
consumption. As a proxy for oxygen penetration, the dissolved [Mn] levels are below detection
(0.027 μmol/kg) within the upper 5 cm at three of our four stations: Stations 1, 5, and 2. For Station
3, we also use the 0-5 cm zone even though [Mn] appears within this depth. The concentration
levels of Mn at Station 3, however, are much lower than at other stations (max Station 3 = 5
μmol/kg; max at Stations 1, 5, 2 = 20-150 μmol/kg), so the [Mn] proxy for oxygen penetration
depth may not be definitive at Station 3.
We also note the transition in the gradients of alkalinity and DIC, from steeper gradients
to more gradual, occurring roughly at 10, 10, 7, and 7 cm for Stations 1, 5, 2, 3, respectively, which
we define as the “breaking points” of each profile, to be used in later analyses. This change in
gradient may indicate a transition from greater rates of DIC production via aerobic processes to
lower rates due to suboxic processes. Below we describe the three approaches to estimate CaCO3
dissolution fluxes.
3.6.2 Dissolution flux via alkalinity
The alkalinity flux from sediments can be converted to a carbonate dissolution flux
assuming a 2:1 alkalinity:CaCO3 ratio (Emerson et al., 1982; Berelson et al., 1990a). However,
there are other reactions that produce alkalinity below the oxygenated zone, so we also consider
that the deep production of alkalinity and its diffusion through the sediment column will result in
this assumption being an upper limit estimate. Assessment of the deeper gradients of alkalinity
indicate that 19 – 37% of the SWI alkalinity flux is coming from deeper than the breaking points,
so we subtract out this deep alkalinity flux to calculate the CaCO3 dissolution flux across the SWI.
By taking this approach, we are assuming that all the alkalinity from deeper in the sediment column
75
is coming from organic matter respiration. Applying this correction to the SWI alkalinity fluxes,
we determine dissolution fluxes to be 0.04±0.04, 0.10±0.04, 0.11±0.06, 0.09±0.04 mmol
CaCO3/m
2
/day, at Stations 1, 5, 2, 3, respectively (Table 3.2).
3.6.3 Dissolution flux via carbon isotopes
The chemical fingerprint of C isotopes of CaCO3 and organic C provide a second method
to determine the relative importance of each reaction to the total DIC flux. We use measured 𝛿
13
C
of DIC to determine the relative sources of DIC, analyzed through Keeling plots (Keeling, 1958).
The fitted 𝛿
13
C data was plotted against the inverse of the fitted DIC concentration, with the y-
intercept indicating the isotopic value of the added DIC. This is an approach similar to Martin et
al. (2000), in which the Keeling plot was broken up by depth horizon; here, we analyze the Keeling
curve for the top 5 cm. An isotopic mass balance was used to calculate the relative contributions
to DIC from a two-member mixing equation:
𝛿
13
C_DIC = CaCO3_fraction * CaCO3_ 𝛿
13
C + Corg_fraction *Corg_ 𝛿
13
C (3.10)
The average CaCO3_fraction from the top 5 cm is multiplied by the DIC flux (derived
from Fick’s Law) to get the CaCO3 dissolution flux. End-member values are assigned from
measured core top values at each station (Appendix Table 3.1). The sensitivity of the solution to
Equation 3.10 to the chosen end-member values of 𝛿
13
C for organic C and CaCO3 is small: a 𝛿
13
C
change of 0.15‰ (our analytical uncertainty) results in an 8% change in flux, smaller than the
overall flux uncertainty. This approach results in 20, 33, 41, and 47% of the DIC produced in 0-5
cm is added through CaCO3 dissolution. These fractions, applied to the DIC fluxes across the
SWI, result in dissolution fluxes of 0.04±0.01, 0.08±0.03, 0.12±0.05, 0.11±0.05 mmol
CaCO3/m
2
/day, at Stations 1, 5, 2, 3, respectively (Table 3.2).

76
3.6.4 Dissolution flux via calcium
The flux of calcium across the SWI is another estimate of net CaCO3 dissolution, assuming
the singular source of [Ca] is from dissolution. We used the model-fit gradient of ∂[Ca]/ ∂z at z =
0 to derive [Ca] fluxes across the SWI. Diffusivity of [Ca] was set to ~4*10
-6
cm
2
s
-1
(Boudreau,
1997); exact value dependent on bottom water temperature. The [Ca] increase above bottom
water with depth in the sediment column between 0-7 cm is consistent with the increase in
alkalinity through this same interval. In the upper ~10 cm, the relationship between porewater
[Ca] and alkalinity is roughly consistent with the expected 1:2 Ca:Alk ratio, but this ratio is
complicated by the coupled diffusivities of the ions. Below the breaking points in alkalinity profiles,
alkalinity tends to increase while [Ca] shows less or negligible change (Figure 3.11). This shift in
slope in [Ca] vs alkalinity space indicates a change in the reaction stoichiometry occurring deeper
in the sediment. Applying Fick’s Law to [Ca] profiles at the SWI, CaCO3 dissolution fluxes were
determined to be 0.04±0.03, 0.11±0.03, 0.07±0.03, 0.06±0.02 mmol CaCO3/m
2
/day, at Stations
1, 5, 2, 3, respectively (Table 3.2).
77

Figure 3.11. Porewater [Ca] vs alkalinity. Black points in bottom left are bottom water values. Solid
points are from SWI to “breaking points” in alkalinity profiles: 10, 10, 7, 7 cm for Stations 1, 5, 2,
3, respectively; hollow points are deeper than breaking points. The change in slope from shallow
to deep points represents where reaction stoichiometry changes. Error bars are standard error of
five replicates in [Ca] space; error bars in alkalinity space are smaller than size of point.

2400 2500 2600 2700 2800 2900 3000
Alkalinity ( mol/kg)
10.2
10.25
10.3
10.35
10.4
10.45
10.5
10.55
10.6
10.65
10.7
Calcium (mmol/kg)
Station 1
Station 5
Station 2
Station 3
2400 2500 2600 2700 2800 2900 3000
Alkalinity ( mol/kg)
10.2
10.25
10.3
10.35
10.4
10.45
10.5
10.55
10.6
10.65
10.7
Calcium (mmol/kg)
Station 1
Station 5
Station 2
Station 3
78
Table 3.2. Alkalinity, DIC, and CaCO3 dissolution fluxes in mmol/m2/day. CaCO3 dissolution
fluxes based on three approaches: shallow alkalinity flux, stable isotopes of DIC, and calcium flux.
Weighted average and uncertainty of three approaches is reported. 1σ in parentheses.

Alk flux
below 7-
10 cm
CaCO3 Dissolution Flux
Station DIC flux
at SWI
Alk flux
at SWI
Via shallow
Alk flux
Via 𝛿
13
C
of DIC
Via [Ca]
flux at SWI
Weighted
Average
1
0.18
(0.07)
0.15
(0.09)
0.06
(0.02)
0.04
(0.04)
0.04
(0.01)
0.04
(0.03)
0.04
(0.01)
5
0.24
(0.08)
0.25
(0.08)
0.05
(0.01)
0.10
(0.04)
0.08
(0.03)
0.11
(0.03)
0.10
(0.02)
2
0.30
(0.12)
0.26
(0.11)
0.05
(0.01)
0.11
(0.06)
0.12
(0.05)
0.07
(0.03)
0.09
(0.02)
3
0.24
(0.10)
0.21
(0.08)
0.04
(0.01)
0.09
(0.04)
0.11
(0.05)
0.06
(0.02)
0.07
(0.02)

3.6.5 Comparison of three approaches
The results of these three independent approaches provide CaCO3 dissolution fluxes for
sediments in the eastern equatorial deep Pacific Ocean that are internally consistent; the weighted
averages at each station have uncertainties <25% between the three approaches.
Station 1 exhibits the least amount of dissolution in the top 35 cm of sediment, while
Stations 5, 2, and 3 exhibit larger and comparable dissolution fluxes. There is consistent evidence
that dissolution is occurring at all sites including where bottom water is saturated for calcite (Station
3). The four stations range in bottom water Ωcalcite from 0.84 to 1, yet there is no correlation
between net dissolution flux and bottom water calcite saturation state, and hence no correlation
with water column depth. Stations 1, 5, and 2 have comparable bottom water O2 (100-105 μM),
while Station 3 has 75 μM O2, but dissolution flux also does not correlate with bottom water O2.
79
Station 1 has the lowest bottom water Ωcalcite, the lowest undersaturation within the
sediment column with omega values <0.55, and undersaturation persisting to 20 cm (Figure 3.12),
which reflects the lack of pH buffering in the upper sediment column. Station 5 bottom water and
the minimum porewater Ωcalcite are lower than at Stations 2 and 3, yet all three of these stations
have undersaturated pore waters persisting to 10-15 cm. Although we demonstrate an extensive
zone within the sediment column where bulk porewater is undersaturated, the degree of
undersaturation or the size of this zone are not the controlling factors of the net dissolution fluxes
measured. However, the depth of undersaturated porewater implies that some dissolution could
be occurring as deep as 10-20 cm, especially at Station 1 where high carbonate sediment is
encountered below 15 cm (Figure 3.10).
A multivariate analysis was performed to determine the best predictors of dissolution flux
out of the following variables: water column depth, bottom water Ωcalcite, bottom water oxygen,
oxygen penetration depth, and % CaCO3. Of these predictor variables, bottom water Ωcalcite and
oxygen penetration depth paired together were statistically the best predictor of dissolution flux
(Figure A.5). However, it is difficult to assign mechanistic understanding to this finding, given the
relatively few data points available for multivariate analysis.
3.6.6 Comparisons to existing flux data
Benthic fluxes in similar environments (water depth 3000-4500 m, Wcalcite range 0.83-1)
range from 0.15-0.7 mmol/m
2
/day in the abyssal Pacific (Berelson et al., 2007). Dissolution fluxes
along 140°W from the equator to 10°N were reported to around 0.4 mmol/m
2
/day at carbonate-
rich sites (Berelson et al., 1990a). Along the equatorial Pacific, dissolution fluxes were measured
between 0.2-0.7 mmol/m
2
/day via benthic flux chambers (Berelson et al., 1994). Dissolution fluxes
in western equatorial Atlantic range from 0.15-0.19 mmol/m
2
/day at a similar saturation range to
80
this study (Martin and Sayles, 1996). These previously published flux estimates are consistent with
our results. Although the compilation of Berelson et al. (2007) shows somewhat of a relationship
between Ωcalcite and net benthic dissolution of CaCO3, our data do not show such a relationship.
3.6.7 Respiration-driven dissolution
Our results confirm the hypothesis of Emerson and Bender (1981) and other studies, that
sedimentary carbonate dissolution can occur at and above the saturation horizon by way of acid
production during aerobic respiration. The evidence presented here includes:
1. Dissolution is observed at all four stations, including Station 3 where bottom water Ωcalcite
= 1.
2. At all stations, porewater Ωcalcite and pH decrease from bottom water values, indicating
some reaction is occurring in the top 5 cm that is lowering saturation state. Because sediments are
oxygenated in this zone, undersaturation here is most likely driven by acid production via aerobic
respiration.
3. The isotopic analysis indicates addition of DIC with
13
C contribution from both organic
C and CaCO3 at depths between 0-5 cm.
These observations and our interpretation suggest that bottom water oxygen, sediment
mixing processes, organic and inorganic C delivery to the seafloor control net dissolution fluxes,
and that within the sediment, dissolution is occurring even where bottom water is saturated with
respect to calcite. Our results speak to the debate regarding precisely where seafloor dissolution is
occurring, whether within the sediment or in the boundary layer between water column and
sediment (Boudreau, 2013). We show, using well-constrained carbonate chemistry profiles, that
there is alkalinity and DIC production occurring below the SWI to at least 5 cm and likely deeper,
indicating carbonate dissolution within sediments. The profiles of porewater omega also point to a
81
driving force for dissolution occurring within the sediment. We think these data offer a convincing
argument that within-sediment processes are more important than surficial-sediment processes, at
least in this region.
3.6.8 Porewater omega
The saturation state (Ω) of a particular mineral phase, in this case calcite or aragonite, is a
thermodynamic threshold controlling the onset of dissolution. However, there are examples of
dissolution occurring where Ω ≥ 1, for both aragonite and calcite in both the water column and in
sediments (Milliman et al., 1999; Lunstrum and Berelson, 2022; Subhas et al., 2022). We believe
we report here the first porewater Ω at cm-scale resolution in deep-sea sediments that span
overlying water to 35 cm depth (Figure 3.12), calculated from porewater alkalinity and pH.
At all stations, omega initially decreases from bottom water values before increasing deeper
in the sediment. Minimum omega values are 0.48, 0.72, 0.87, and 0.89 for Stations 1, 5, 2, 3,
respectively. This minimum appears between 3-8 cm below the SWI. Stations 1, 2, and 3 approach
and appear to stabilize around Ωcalcite = 1, indicating the dominant phase controlling carbonate
equilibria in these sediments is likely calcite (with thermodynamic properties as defined by
CO2SYS). Station 5, however, surpasses calcite saturation and approaches a higher Omega value.
Additionally, Ωcalcite at two sites (Stations 1 and 5) dips below two critical omega thresholds (0.9
and 0.75) where the mechanism of dissolution is known to change between step-wise, defect-
assisted etch pit, and finally non-defect-assisted etch pits (Naviaux et al., 2019a; Adkins et al., 2021).

82

Figure 3.12. Porewater profile of saturation state with respect to calcite, calculated with alkalinity
and pH. Arrows denote bottom water omega to their respective stations. Vertical lines show Ω =
1 for calcite and aragonite. Error bars are smaller than size of point; 1σ of duplicate measurements
range from 0.001 – 0.003 Ω.

To further investigate the mineral solubility control on omega and the reason for porewater
Ωcalcite increasing >1 at Station 5, we used two methods to determine if aragonite was present in
sediment. X-ray diffraction was used to identify the mineralogy of the sediment. The top 1 cm of
sediment was measured at each station, plus four additional depth horizons at Station 5 up to 35
cm. No aragonite was detected; only calcite and quartz were detected (clay minerals were not
assessed). It is possible that aragonite was present below detection limits (<1%). It is also possible
that at Station 5, aragonite is present below 35 cm, the maximum depth of the collected sediment
core. Such a case could occur if an extreme sedimentation event delivered aragonite to the seafloor
that was buried and preserved faster than it could dissolve. We do not dismiss the possibility that a
0.5 0.75 1 1.25 1.5
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
Ω aragonite = 1
1
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
Ω calcite = 1
83
carbonate mineral with a solubility other than calcite or aragonite as defined by CO2SYS (Mucci,
1983) is present in these sediments, although it would be unlikely that it was only present at Station
5.
A second method to identify the presence of aragonite was through measurements of
dissolved strontium (Sr), as aragonite contains more [Sr] than calcite (Dodd, 1967). Stations 1, 2,
and 3 show minimal to no increase in [Sr] from bottom water to 35 cm (Figure 3.13). However,
Station 5 shows increasing [Sr] with depth and does not appear to reach an asymptotic value at
depth. This [Sr] profile is consistent with the interpretation that aragonite is buried >35 cm,
potentially supplied by coral and platform carbonate input from nearby Cocos Ridge. A small
increase in [Sr] below bottom water values was also detected at Station 3, which sits on Cocos
Ridge close to sources of aragonitic shallow sediments. Regarding the proposal that aragonite is
buffering calcite dissolution in sediments (Sulpis et al, 2022), we do not see evidence of this at
Stations 1, 2, or 3, but perhaps see this phenomenon at Station 5 where [Sr] increases deeper in
the sediment.
84

Figure 3.13. Porewater [Sr] (μmol/kg) (following normalization to [Na]). Arrows represent bottom
water. Error bars are standard error of five replicates.

3.6.9 Sedimentary mass balance and paleo-implications
Mass accumulation rates were determined using C14 ages of foraminifera (Figure A.4).
Assuming linear sedimentation rate, a linear curve was fit through the points of the profile to get
sedimentation rate. The sedimentation rate was then converted to mass accumulation rate using
the dry bulk density of sediment based on porosity. From these mass accumulation rates and the
dissolution fluxes described above, we built a sedimentary CaCO3 budget for the Cocos Ridge
region (Table 3.3). While dissolution flux is not correlated with water depth and saturation state,
CaCO3 burial efficiency (defined as PIC burial rate/PIC rain rate to the seafloor) is well-correlated
with omega. Burial efficiency increases seven-fold in moving only 300 m shallower from Station 1
to 5. This implies the sensitivity of the sedimentary lysocline is more dependent on aerobic
respiration and PIC delivery to the sea floor than to bottom water omega. We also show that ~35%
84 86 88 90
Strontium ( mol/kg)
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Strontium
Station 1
Station 2
Station 3
Station 5
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
0.4 0.6 0.8 1 1.2 1.4 1.6
calcite
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Porewater Alk/pH Omega
Station 1
Station 2
Station 3
Station 5
85
of the carbonate falling to the seafloor at 1600 m dissolves before burial, even though this station
sits at the calcite saturation horizon.
Surface ocean biological productivity can control deep-sea dissolution, as the relative rain
rates of CaCO3 and organic C impact dissolution and preservation (Emerson and Bender, 1981;
Archer and Maier-Reimer, 1994). However, biological productivity at these four stations is equally
low as suggested by satellite-derived mean ocean chlorophyll (~0.1 mg Chl/m
3
) in this region
(Pennington et al., 2006) and the assumption that productivity may not vary much between these
closely spaced stations. The input rate of organic C into these sediments is not part of the scope of
this study, but it is clearly coupled to the overall dissolution fluxes we have established. Sediment
mixing and incorporation of organic C into the sediment column and organic C reactivity are also
likely factors in the preservation of CaCO3.

Table 3.3. Mass balance of Cocos Ridge surface sediments (0-1 cm), including mass accumulation
rates via Carbon-14, CaCO3 accumulation rate, CaCO3 rain rate (a derived parameter based on
the sum of dissolution and accumulation), and CaCO3 burial efficiency, assuming a steady state.
Rates and fluxes are in g/cm2/kyr. Dissolution fluxes (units converted from Table 3.2) and
subsequent calculations are averages based on the three approaches to calculating dissolution rate
(Table 3.2).
Station
Water
depth
(m)
Mass
accumulation
rate
%
CaCO3
CaCO3
accumulation
rate
CaCO3
dissolution
flux
CaCO3
rain
rate
CaCO3
burial
efficiency
(%)
1 3223 0.49 0.8 0.00 0.15 0.15 3
86
5 2911 0.56 13 0.07 0.37 0.44 16
2 2650 1.04 63 0.65 0.33 0.98 66
3 1630 0.97 78 0.75 0.26 1.0 75

The mass balance described in Table 3.3 makes the assumption that the system is balanced,
i.e., rain to the seafloor = sum of burial and dissolution. But if net chemical erosion of deeper
carbonate sediments is occurring at Station 1 (or elsewhere), then this part of the ocean is not a
balanced system. Because the % CaCO3 in the upper 10 cm is < 2% (Figure 3.10) and low omega
porewater values persist to ~20 cm (Figure 3.12), it is likely that Station 1 is out of balance and is
undergoing net erosion of deeper carbonate sediment. Given that % CaCO3 also increases at
Station 5, and undersaturation persists below 10 cm, it is possible that deep dissolution is occurring
at this and possibly at other stations as well. However, the porewater alkalinity, [Ca] and DIC
profiles point toward reactions occurring at the highest magnitude in the top ~10 cm.
Though we have not defined the biogenic constituents that are dissolving to produce the
flux we observe, we assume that forams and other carbonate-secreting organisms undergo
dissolution not just on top of the seabed, but within the sediment as well. This has implications for
paleoceanographic studies that make assumptions about post-depositional dissolution of
carbonate-secreting organisms.

3.7 Summary
We present the results and interpretations of in situ collected porewater profiles from the
Cocos Ridge in the eastern equatorial Pacific on a transect spanning the saturation horizon. We
87
measured alkalinity, pH, DIC, 𝛿
13
C of DIC, calcium, strontium, and manganese in porewater
from 0-35 cm and solid phases % CaCO3, 𝛿
13
C of PIC, % organic C, and 𝛿
13
C of organic C.
Through profile curve fitting and applying Fick’s first law, we calculated fluxes of alkalinity, DIC,
and [Ca] from the sediment to water column. Thus, we have three independent estimates for
dissolution fluxes: 1) alkalinity fluxes, 2) stable isotopes of DIC, and 3) calcium fluxes. Dissolution
fluxes among these three approaches have a weighted mean with uncertainty of <25%: 0.04±0.01,
0.10±0.02, 0.09±0.02, and 0.07±0.02 mmol/m
2
/day at sites 3200, 2900, 2700, 1600 m deep,
respectively. We see lowest CaCO3 dissolution flux at the deepest and most undersaturated site
(Ω= 0.84), and higher but comparable dissolution fluxes at the remaining sites ranging from 2900-
1600 m and Ωcalcite values of 0.89 - 1. Respiration-driven CaCO3 dissolution is occurring at all sites,
including at Station 3 where bottom water is saturated with respect to calcite. These dissolution
fluxes are not a simple function of water depth, bottom water omega, bottom water oxygen, or
surface ocean biological productivity. Therefore, dissolution must be attributed to diagenetic
reactions occurring in the sediments rather than reflecting bottom water properties alone.
Bottom water oxygen content and the lability and distribution of organic C within the
sediments may be important co-factors of dissolution. Based on the approaches above that provide
estimates of the relative fractions of aerobic respiration and dissolution, we attribute 20-50% of the
observed DIC flux across the SWI to respiration-driven dissolution. When making global estimates
of the potential for seafloor dissolution and in trying to interpret the shape of the lysocline, bottom
water carbonate saturation state alone is not enough to predict the presence or magnitude of
dissolution in a given region.

3.8 Acknowledgements
88
This work was supported by NSF Ocean Acidification (OCE-1834475). J.E.P.C. thanks
USC Earth Sciences for PhD funding for this work. The authors thank Aaron Celestian (Natural
History Museum of Los Angeles) and Menglong Zhang (Nanjing University) for assistance with
XRD mineralogy. J.E.P.C thanks her dissertation committee members, Professors Doug
Hammond and Naomi Levine, for helpful comments on the manuscript. The science party,
captain, and crew of R/V Sally Ride SR2113 cruise were instrumental in collecting the data
presented here.

89
3.9 References
Adkins J. F., Naviaux J. D., Subhas A. V., Dong S. and Berelson W. M. (2021) The Dissolution
Rate of CaCO3 in the Ocean. Ann Rev Mar Sci 13, 1–24.
Archer D. (1991) Modeling the calcite lysocline. J Geophys Res 96.
Archer D. E., Morford J. L. and Emerson S. R. (2002) A model of suboxic sedimentary diagenesis
suitable for automatic tuning and gridded global domains. Global Biogeochem Cycles 16, 17-1-17–
21.
Archer D., Eby M., Brovkin V., Ridgwell A., Cao L., Mikolajewicz U., Caldeira K., Matsumoto
K., Munhoven G., Montenegro A. and Tokos K. (2009) Atmospheric Lifetime of Fossil Fuel
Carbon Dioxide. Annu Rev Earth Planet Sci 37, 117–134.
Archer D., Emerson S. and Reimers C. (1989) Dissolution of calcite in deep-sea sediments: pH
and O2 microelectrode results. Geochim Cosmochim Acta 53, 2831–2845.
Archer D. and Maier-Reimer E. (1994) Effect of deep-sea sedimentary calcite preservation on
atmospheric CO2 concentration. Nature 367, 260–263.
Balch W., Drapeau D., Bowler B. and Booth E. (2007) Prediction of pelagic calcification rates using
satellite measurements. Deep Sea Res 2 Top Stud Oceanogr 54, 478–495.
Berelson W. M., Balch W. M., Najjar R., Feely R. A., Sabine C. and Lee K. (2007) Relating
estimates of CaCO3 production, export, and dissolution in the water column to measurements
of CaCO3 rain into sediment traps and dissolution on the sea floor:A revised global carbonate
budget. Global Biogeochem Cycles 21, 1–15.
Berelson W. M., Hammond D. E. and Cutter G. A. (1990a) In situ measurements of calcium
carbonate dissolution rates in deep-sea sediments. Geochim Cosmochim Acta 54, 3013–3020.
Berelson W. M., Hammond D. E., Mcmanus J. and Kilgore T. E. (1994) Dissolution kinetics of
calcium carbonate in equatorial Pacific sediments. Global Biogeochem Cycles 8, 219–235.
Berelson W. M., Hammond D. E., O’Neill D., Xu X. m., Chin C. and Zukin J. (1990b) Benthic
fluxes and pore water studies from sediments of the central equatorial north Pacific: Nutrient
diagenesis. Geochim Cosmochim Acta 54, 3001–3012.
Berelson W. M., Prokopenko M., Sansone F. J., Graham A. W., McManus J. and Bernhard J. M.
(2005) Anaerobic diagenesis of silica and carbon in continental margin sediments: Discrete
zones of TCO2 production. Geochim Cosmochim Acta 69, 4611–4629.
Boudreau B. P. (2013) Carbonate dissolution rates at the deep ocean floor. Geophys Res Lett 40, 744–
748.
Boudreau B. P. (1997) Diagenetic Models and Their Implementation: Modelling Transport and Reactions in
Aquatic Sediments., Springer, Halifax, NS.
90
Boudreau B. P., Sulpis O. and Mucci A. (2020) Control of CaCO3 dissolution at the deep seafloor
and its consequences. Geochim Cosmochim Acta 268, 90–106.
Broecker W., Mix A., Andree M. and Oeschger H. (1984) Radiocarbon measurements on
coexisting benthic and planktic foraminifera shells: potential for reconstructing ocean
ventilation times over the past 20,000 years. Nuclear Instruments and Methods in Physics Research,
331–339.
Broecker W. S. and Broecker S. (1974) Carbonate dissolution on the western flank of the East
Pacific Rise. Studies in Paleo-Oceanography, 44–57.
Burdige D. J. (1993) The biogeochemistry of manganese and iron reduction in marine sediments.
Earth Sci Rev 35, 249–284.
Cetiner J. E. P., Berelson W. M., Rollins N. E., Barnhart H. A., Liu X., Dong S., Byrne R. H. and
Adkins J. F. (2023) Novel device to collect deep-sea porewater in situ: A focus on benthic
carbonate chemistry. Limnol Oceanogr Methods 21, 82–97.
Chong L. S., Berelson W. M., McManus J. and Rollins N. E. (2018) Meter-scale early diagenesis
of organic matter buried within deep-sea sediments beneath the Amazon River plume. Front
Mar Sci 5, 1–15.
Dickson A. G. and Millero F. J. (1987) A comparison of the equilibrium constants for the
dissociation of carbonic acid in seawater media. Deep-Sea Research 34, 1733-1743.
Dodd J. R. (1967) Magnesium and Strontium in Calcareous Skeletons: A Review. Journal of
Paleontology 41, 1313–1329.
Doney S. C., Fabry V. J., Feely R. A. and Kleypas J. A. (2009) Ocean Acidification: The Other
CO2 Problem. Ann Rev Mar Sci 1, 169–192.
Dong S., Subhas A. V., Rollins N. E., Naviaux J. D., Adkins J. F. and Berelson W. M. (2018) A
kinetic pressure effect on calcite dissolution in seawater. Geochim Cosmochim Acta 238, 411–423.
Emerson S. and Bender M. (1981) Carbon fluxes at the sediment-water interface of the deep-sea:
calcium carbonate preservation. J Mar Res 39, 139–162.
Emerson S., Grundmanis V. and Graham D. (1982) Carbonate chemistry in marine pore waters:
MANOP sites C and S. Earth Planet Sci Lett 61, 220–232.
Feely R. A., Sabine C. L., Lee K., Millero F. J., Lamb M. F., Greeley D., Bullister J. L., Key R.
M., Peng T.-H., Kozyr A., Ono T. and Wong C. S. (2002) In situ calcium carbonate dissolution
in the Pacific Ocean. Global Biogeochem Cycles 16, 91-1-91–12.
Froelich P. N., Klinkhammer G. P., Bender M. L., Luedtke N. A., Heath G. R., Cullen D.,
Dauphin P., Hammond D., Hartman B. and Maynard V. (1979) Early oxidation of organic
matter in pelagic sediments of the eastern equatorial Atlantic: suboxic diagenesis. Geochim
Cosmochim Acta 43, 1075–1090.
91
Green M. A., Aller R. C. and Aller J. Y. (1998) Influence of carbonate dissolution on survival of
shell-bearing meiobenthos in nearshore sediments. Limnol Oceanogr 43, 18–28.
Hales B. and Emerson S. (1997) Calcite dissolution in sediments of the Ceara Rise: In situ
measurements of porewater O2, pH, and CO2(aq). Geochim Cosmochim Acta 61, 501–514.
Hales B., Emerson S. and Archer D. (1994) Respiration and dissolution in the sediments of the
western North Atlantic: estimates from models of in situ microelectrode measurements of
porewater oxygen and pH. Deep-Sea Research Part I 41, 695–719.
Hansson I. (1973) A new set of acidity constants for carbonic acid and boric acid in sea water. Deep
Sea Research and Oceanographic Abstracts 20, 461–478.
Hansson I. (1972) An analytical approach to the carbonate system in sea water. Göteborgs
universitet.
Hurd D. C. (1972) Factors affecting solution rate of biogenic opal in seawater. Earth Planet Sci Lett
15, 411–417.
Jahnke R. A., Craven D. B., McCorkle D. C. and Reimers C. E. (1997) CaCO3 dissolution in
California continental margin sediments: The influence of organic matter remineralization.
Geochim Cosmochim Acta 61, 3587–3604.
Jahnke R. A. and Jahnke D. B. (2004) Calcium carbonate dissolution in deep sea sediments:
Reconciling microelectrode, pore water and benthic flux chamber results. Geochim Cosmochim
Acta 68, 47–59.
Keeling C. D. (1958) The concentration and isotopic abundances of atmospheric carbon dioxide
in rural areas. Geochim Cosmochim Acta 13, 322–334.
Keir R. S. (1980) The dissolution kinetics of biogenic calcium carbonates in seawater. Geochim
Cosmochim Acta 44, 241–252.
Lea D. W., Pak D. K. and Spero H. J. (2000) Climate impact of late quaternary equatorial Pacific
sea surface temperature variations. Science (1979) 289, 1719–1724.
Lewis E. and Wallace D. (1998) Program developed for CO2 system calculations. Oak Ridge
National Laboratory/Carbon Dioxide Information Analysis Center.
Liao Y. and Lyle M. (2014) Late Miocene to Pleistocene sedimentation and sediment transport on
the Cocos Ridge, eastern tropical Pacific Ocean. Mar Geol 355, 1–14.
Liu X., Byrne R. H., Adornato L., Yates K. K., Kaltenbacher E., Ding X. and Yang B. (2013) In
situ spectrophotometric measurement of dissolved inorganic carbon in seawater. Environ Sci
Technol 47, 11106–11114.
Liu X., Patsavas M. C. and Byrne R. H. (2011) Purification and characterization of meta-cresol
purple for spectrophotometric seawater ph measurements. Environ Sci Technol 45, 4862–4868.
92
Liu X., Wang Z. A., Byrne R. H., Kaltenbacher E. A. and Bernstein R. E. (2006)
Spectrophotometric measurements of pH in-situ: Laboratory and field evaluations of
instrumental performance. Environ Sci Technol 40, 5036–5044.
Lunstrum A. and Berelson W. (2022) CaCO3 dissolution in carbonate-poor shelf sands increases
with ocean acidification and porewater residence time. Geochim Cosmochim Acta 329, 168–184.
Martin W. R., McNichol A. P. and McCorkle D. C. (2000) The radiocarbon age of calcite
dissolving at the sea floor: Estimates from pore water data. Geochim Cosmochim Acta 64, 1391–
1404.
Martin W. R. and Sayles F. L. (1996) CaCO3 dissolution in sediments of the Ceara Rise, western
equatorial Atlantic. Geochim Cosmochim Acta 60, 243–263.
Martin W. R. and Sayles F. L. (2006) Organic matter oxidation in deep-sea sediments: Distribution
in the sediment column and implications for calcite dissolution. Deep Sea Res 2 Top Stud Oceanogr
53, 771–792.
McManus J., Hammond D. E., Berelson W. M., Kilgore T. E., Demaster D. J., Ragueneau O. G.
and Collier R. W. (1995) Early diagenesis of biogenic opal: Dissolution rates, kinetics, and
paleoceanographic implications. Deep-Sea Research Part II 42, 871–903.
Mehrbach C., Culberson C. H., Hawley J. E. and Pytkowicx R. M. (1973) Measurement of the
Apparent Dissociation Constants of Carbonic Acid in Seawater at Atmospheric Pressure.
Limnol Oceanogr 18, 897–907.
Milliman J. D., Troy P. J., Balch W. M., Adams A. K., Li Y. H. and Mackenzie F. T. (1999)
Biologically mediated dissolution of calcium carbonate above the chemical lysocline? Deep Sea
Res 1 Oceanogr Res Pap 46, 1653–1669.
Moore T. C., Heath G. R. and Kowsmann R. O. (1973) Biogenic Sediments of the Panama Basin.
J Geol 81, 458–472.
Mucci A. (1983) The solubility of calcite and aragonite in seawater at various salinities,
temperatures, and one atmosphere total pressure. Am J Sci 283, 780–799.
Murray J. W., Johnson E. and Garside C. (1995) A U.S. JGOFS process study in the equatorial
Pacific (EqPac): Introduction. Deep-Sea Research Part II 42, 275–293.
Murray J. W., Leborgne R. and Dandonneau Y. (1997) JGOFS studies in the equatorial pacific.
Deep Sea Res 2 Top Stud Oceanogr 44, 1759–1763.
Naviaux J. D., Subhas A. V., Rollins N. E., Dong S., Berelson W. M. and Adkins J. F. (2019)
Temperature dependence of calcite dissolution kinetics in seawater. Geochim Cosmochim Acta
246, 363–384.
Pennington J. T., Mahoney K. L., Kuwahara V. S., Kolber D. D., Calienes R. and Chavez F. P.
(2006) Primary production in the eastern tropical Pacific: A review. Prog Oceanogr 69, 285–317.
93
Pilson M. E. Q. (1998) An Introduction to the Chemistry of the Sea., Prentice Hall.
Reimers C., Jahnke R. and McCorkle D. (1992) Carbon fluxes and burial rates over the continental
slope and rise off central California with implications for the global carbon cycle. Global
Biogeochem Cycles 6, 199–224.
Sabine C. L., Feely R. A., Gruber N., Key R. M., Lee K., Bullister J. L., Wanninkhof R., Wong
C. S., Wallace D., Tilbrook B., Millero F. J., Peng T.-H., Kozyr A., Ono T. and Rios A. F.
(2004) The Oceanic Sink for Anthropogenic CO2. Science (1979) 305, 367–371.
Sayles F. L. and Dickinson W. H. (1991) The ROLAI^2D lander: A benthic lander for the study
of exchange across the sediment-water interface. Deep Sea Research Part A, Oceanographic Research
Papers 38, 505–529.
Sayles F. L., Martin W. R., Chase Z. and Anderson R. F. (2001) Benthic remineralization and
burial of biogenic SiO2, CaCO3, organic carbon, and detrital material in the Southern Ocean
along a transect at 170° West. Deep Sea Res 2 Top Stud Oceanogr 48, 4323–4383.
Steiner Z., Rae J. W. B., Berelson W. M., Adkins J. F., Hou Y., Dong S., Lampronti G. I., Liu X.,
Achterberg E. P., Subhas A. V. and Turchyn A. V. (2022) Authigenic Formation of Clay
Minerals in the Abyssal North Pacific. Global Biogeochem Cycles 36.
Steiner Z., Sarkar A., Liu X., Berelson W. M., Adkins J. F., Achterberg E. P., Sabu P., Prakash S.,
Vinaychandran P. N., Byrne R. H. and Turchyn A. V. (2021) On calcium-to-alkalinity
anomalies in the North Pacific, Red Sea, Indian Ocean and Southern Ocean. Geochim
Cosmochim Acta 303, 1–14.
Stuiver M. and Polach H. A. (1977) Discussion Reporting of 14 C Data . Radiocarbon 19, 355–363.
Subhas A. V, Dong S., Naviaux J. D., Rollins N. E., Ziveri P., Gray W., Rae J. W. B., Liu X.,
Byrne R. H., Chen S., Moore C., Martell-bonet L., Steiner Z., Antler G., Hu H., Lunstrum
A., Hou Y., Kemnitz N., Stutsman J., Pallacks S., Dugenne M., Quay P. D., Berelson W. M.
and Adkins J. F. (2022) Shallow Calcium Carbonate Cycling in the North Pacific Ocean Global
Biogeochemical Cycles. Global Biogeochem Cycles 2, 1–22.
Subhas A. V., Rollins N. E., Berelson W. M., Dong S., Erez J. and Adkins J. F. (2015) A novel
determination of calcite dissolution kinetics in seawater. Geochim Cosmochim Acta 170, 51–68.
Sulpis O., Boudreau B. P., Mucci A., Jenkins C., Trossman D. S., Arbic B. K. and Key R. M.
(2018) Current CaCO3 dissolution at the seafloor caused by anthropogenic CO2. Proceedings of
the National Academy of Sciences, 1-6.
Sulpis O., Lix C., Mucci A. and Boudreau B. P. (2017) Calcite dissolution kinetics at the sediment-
water interface in natural seawater. Mar Chem 195, 70–83.
Tans P. P., Fung I. Y. and Takahashi T. (1990) Observational constraints on the global
atmospheric CO2 budget. Science (1979) 247, 1431–1438.
94
Taylor J. R. (1997) An Introduction to Error Analysis: The Study of Uncertainties in Physical Measurements.
2nd ed., University Science Books.

95
Chapter 4: Conclusions
In December 2021, along with 16 other scientists, I cruised to the Cocos Ridge in the
eastern equatorial Pacific to study carbonate dissolution fluxes in the sediment. The goal of this
cruise was to study carbonate dissolution at and below the calcite saturation horizon in order to
better understand the role of respiration-driven dissolution in carbonate diagenesis. Chapter 2
describes the SIPR device I built to filter porewater from deep-sea sediments in situ, avoid sampling
artifacts associated with traditional core recovery methods, and obtain high quality samples for
accurately characterizing porewater chemistry. This chapter also includes a detailed evaluation of
the sampling artifacts that are present through ex situ collected porewater, demonstrating the
necessity of in situ sampling, particularly with regards to carbonate chemistry parameters.
Chapter 3 analyzes the data collected from SIPR on a cruise transect across the Cocos
Ridge in the eastern Equatorial Pacific. From carbonate parameters and modeling, I calculated
dissolution fluxes from sediments and constructed a sedimentary mass budget which included
accumulation rates from Carbon-14 measurements. Dissolution fluxes across the SWI were
derived by three independent methods: alkalinity fluxes, stable isotopes of DIC, and calcium fluxes.
Particularly from the DIC isotopic approach, it is clear that respiration-driven carbonate
dissolution is occurring at all stations, with up to 50% of the DIC flux originating from carbonate
dissolution. This explains how dissolution is occurring even where bottom water is saturated with
respect to calcite. Thus, the rain of organic carbon to the seafloor and the penetration of oxygen
into the sediment column appear to be controlling seafloor carbonate dissolution.
A novel observation from the porewater data is the shape of the alkalinity and DIC profiles,
and the extent of calcite undersaturation with depth. The curvature in alkalinity and DIC profiles
changes from site to site but all sites show concave-down profiles indicative of production of DIC
96
and alkalinity in the upper 10 cm. All sites decrease in saturation state from bottom water values
between 0.5-0.85 between 0-10 cm before increasing to or exceeding saturation. In combination,
these data confirm sediment dissolution occurring well below the SWI. At one site, Station 1, it
appears that there is net dissolution of carbonate occurring deeper than 15 cm, that is, carbonate
aged 5,000 – 10,000 years old is undergoing net dissolution.
The question also arises of whether this dissolving carbonate is primarily calcite or
aragonite. From the measurements made on the solid phase of these sediments, no aragonite was
detected. However, there is possible evidence of aragonite buried deeper than 30 cm at Station 5
based on porewater Wcalcite increasing past 1, as well as dissolved strontium increasing with depth.
These data suggest aragonite is buried deeper in the sediment that can be providing some of the
alkalinity flux observed, potentially buffering calcite dissolution.
Another debate that this study addresses is the sediment-side vs water-side control of
dissolution, that is, whether carbonate is dissolving primarily in the sediment or in the diffusive
boundary layer in the overlying water. The shape of the alkalinity and DIC profiles analyzed in
this study show clear evidence of sediment-side control and that carbonate is dissolving within the
sediments themselves, not just on the surface of sediments.
This dissertation sheds light on the controls of carbonate dissolution in the deep seafloor,
which is an essential piece to understanding the relationship between the ocean and climate
regulation. This work shows that carbonate can dissolve in unexpected regions and that it is not
necessarily correlated to the most canonical driver, that is, saturation state of the bottom water.
Indeed, I found dissolution fluxes in sediments to be lowest at the site that is most undersaturated,
which is where dissolution would be expected to be highest if bottom water saturation state was
the most important factor. Instead, dissolution appears to be driven more by the relative amounts
of carbonate and organic matter rain arriving at the seafloor. The early diagenetic reactions that
97
happen in the recently deposited sediments, namely aerobic respiration, create acid that in turn
can dissolve carbonate, but only if there are meaningful amounts of carbonate present on the
seafloor to begin with. These reactions, ubiquitously occurring in ocean sediments as long as there
is an oxygen source and an organic matter source, can dissolve more carbonate than previously
estimated, which would neutralize more oceanic CO2 and, in turn, allow the ocean to draw down
more atmospheric CO2.

98
Appendices
Supplementary Material to Chapter 3

0 5 10 15 20
IAPSO #
450
455
460
465
470
475
480
485
490
495
500
Sodium (mmol/kg)
0 5 10 15 20
IAPSO #
80
81
82
83
84
85
86
87
88
89
90
Strontium (umol/kg)
0 5 10 15 20
IAPSO #
10
10.1
10.2
10.3
10.4
10.5
10.6
10.7
Calcium (mmol/kg)
Ca
Sr Na
0 5 10 15 20
IAPSO #
450
455
460
465
470
475
480
485
490
495
500
Sodium (mmol/kg)
0 5 10 15 20
IAPSO #
80
81
82
83
84
85
86
87
88
89
90
Strontium (umol/kg)
0 5 10 15 20
IAPSO #
10
10.1
10.2
10.3
10.4
10.5
10.6
10.7
Calcium (mmol/kg)
Ca
Sr Na
99

Figure A.1. ICP-OES IAPSO consistency standards across all runs, by which calcium and
strontium samples were normalized.

Figure A.2. Water column alkalinity and DIC comparing measurements from this study vs
GLODAP data, including 1:1 lines.
0 5 10 15 20
IAPSO #
450
455
460
465
470
475
480
485
490
495
500
Sodium (mmol/kg)
0 5 10 15 20
IAPSO #
80
81
82
83
84
85
86
87
88
89
90
Strontium (umol/kg)
0 5 10 15 20
IAPSO #
10
10.1
10.2
10.3
10.4
10.5
10.6
10.7
Calcium (mmol/kg)
Ca
Sr Na
2200
2300
2400
2200 2300 2400
GLODAP Alkalinity (µmol/kg)
Cocos Ridge Alkalinity (µmol/kg)
2200
2300
2400
2200 2300 2400
GLODAP Alkalinity (µmol/kg)
Cocos Ridge Alkalinity (µmol/kg) Measured
2200
2300
2400
2200 2300 2400
GLODAP Alkalinity (µmol/kg)
Cocos Ridge Alkalinity (µmol/kg)
2200
2300
2400
2200 2300 2400
GLODAP Alkalinity (µmol/kg)
Cocos Ridge Alkalinity (µmol/kg)
2200
2300
2400
2200 2300 2400
GLODAP Alkalinity (µmol/kg)
Cocos Ridge Alkalinity (µmol/kg)
2200
2300
2400
2200 2300 2400
GLODAP Alkalinity (µmol/kg)
Cocos Ridge Alkalinity (µmol/kg)
100
Table A.1. Core top (0-1 cm) solid phase properties. PIC 𝛿
13
C and POC 𝛿
13
C values were used as
end-members in Keeling analysis.
Station
% Total
Carbon
% CaCO3
% Organic
Carbon
Total
Carbon
𝛿
13
C (‰)
Particulate
Inorganic
Carbon
𝛿
13
C (‰)
Particulate
Organic
Carbon
𝛿
13
C (‰)
1 1.2 0.83 1.11 -17.8 -2.1* -19.2
5 2.4 12.8 0.91 -7.13 0.39 -19.8
2 7.8 62.7 0.41 -0.19 1.1 -23.5**
3 9.8 77.9 0.47 0.13 1.1 -19.6
Footnotes: *Suspect due to low %CaCO3, approaching instrument detection limit. **Suspect due
to profile scatter and difference from other stations.
101

Figure A.3. Keeling plots with intercepts listed for every 1 cm sediment horizon.

102

Figure A.4. Foraminifera ages determined by Carbon-14.

0 0.5 1 1.5 2 2.5
C14 Age (yr) 10
4
0
5
10
15
20
25
30
35
Sediment depth (cm)
CDISP Foram Ages
Station 1
Station 2
Station 3
Station 5
2400 2500 2600 2700 2800 2900 3000
Alkalinity ( mol/kg)
10.2
10.25
10.3
10.35
10.4
10.45
10.5
10.55
10.6
10.65
10.7
Calcium (mmol/kg)
Station 1
Station 5
Station 2
Station 3
0 0.05 0.1 0.15
Actual dissolution flux (mmol/m2/day)
0
0.05
0.1
0.15
Predicted dissolution flux (mmol/m2/day)
Predictor variables: BW omega & O
2
penetration depth
r
2
= 0.974
103

Figure A.5 Multivariate analysis determining the best predictors of dissolution flux. The predictor
variables were water column depth, bottom water Ωcalcite, bottom water oxygen, oxygen
penetration depth, and % CaCO3. Bottom water Ωcalcite and oxygen penetration depth paired
together were statistically the best predictor of dissolution flux (top plot). The second and third best
predictor pairs are also shown in the middle and bottom plots.

0 0.05 0.1 0.15
Actual dissolution flux (mmol/m2/day)
0
0.05
0.1
0.15
Predicted dissolution flux (mmol/m2/day)
Predictor variables: Water column depth & O
2
penetration depth
r
2
= 0.823
0 0.05 0.1 0.15
Actual dissolution flux (mmol/m2/day)
0
0.05
0.1
0.15
Predicted dissolution flux (mmol/m2/day)
Predictor variables: % carbonate & O
2
penetration depth
r
2
= 0.574
104
Ex situ processed porewater data
Appended below are the data from sediment cores processed ex situ, on board the ship,
with Rhizon filters from Cocos Ridge sediments. These data are used in Chapter 2 to compare ex
situ and in situ porewater collection to highlight the sampling artifacts associated with ex situ
porewater filtration. Descriptions of table columns:
No data: “n.d.” stands for “no data.” This is either because that parameter was not measured for
that sample, or because it was measured but eliminated due to a flag on the analytical measurement
or considerable ( > 3 standard deviations) difference from profile trend.
Sample IDs: Sample IDs are named CR (for “Cocos Ridge”) - Core # - Sample #.
Depth: For ex situ filtered porewater, depth (cm) is determined by measuring from sediment-
water interface.
Alkalinity: Total Alkalinity (μmol/kg) measured on mvMICA. See Chapter 2 for full description
of methods.
Silica: Dissolved silicate (μM) measured on spectrophotometer. See Chapter 2 for full description
of methods.
DIC: Dissolved Inorganic Carbon (μmol/kg) measured on Picarro Cavity Ring-Down
Spectrometer. See Chapter 2 for full description of methods.
𝛿
13
C of DIC: 𝛿
13
C of DIC (‰) measured on Picarro Cavity Ring-Down Spectrometer. See
Chapter 2 for full description of methods.

105
Table A.2. Ex situ porewater data from Rhizon-filtered multi-cores.
Sample ID Depth (cm)
Alkalinity
(μmol/kg) Silica (μM)
DIC
(μmol/kg)
𝛿
13
C of
DIC (‰)
CR1-13-1 1 2408 204 n.d. n.d.
CR1-13-2 2 2454 220 n.d. n.d.
CR1-13-3 3 2493 244 2551 -1.92
CR1-13-4 5 2484 251 n.d. n.d.
CR1-13-5 7 2545 310 2620 -1.99
CR1-13-6 9 2612 322 n.d. n.d.
CR1-13-7 11 2626 331 2655 -2.17
CR1-13-8 13 2663 344 n.d. n.d.
CR1-13-9 16 2658 351 n.d. n.d.
CR1-13-10 20 2691 358 n.d. n.d.
CR1-13-11 24 2732 362 2667 -2.33
CR1-13-12 27 2732 359 n.d. n.d.
CR1-15-1 2 2371 212 2386 -1.01
CR1-15-2 5 2460 247 2473 -1.15
CR1-15-3 9 2551 298 2576 -1.69
CR1-15-4 14 2640 335 2647 -1.98
CR1-15-5 18 2712 344 n.d. n.d.
CR1-15-6 22 2745 350 2653 -2.11
CR1-15-7 25 2749 353 n.d. n.d.
CR1-15-8 30 2814 356 2757 -2.27
CR1-14-1 1.5 2441 199 n.d. n.d.
CR1-14-2 3.5 2424 218 2485 -1.60
CR1-14-3 4.5 2457 239 n.d. n.d.
CR1-14-4 6.5 n.d. 283 2585 -2.08
CR1-14-5 8.5 2562 305 n.d. n.d.
CR1-14-6 11.5 2624 321 2681 -2.36
CR1-14-7 15.5 2644 321 n.d. n.d.
CR1-14-8 19.5 2733 354 2765 -2.77
CR1-14-9 23.5 2764 364 n.d. n.d.
CR1-14-10 28.5 2819 371 2811 -2.61
CR1-14-11 33.5 2814 372 n.d. n.d.
CR2-3-1 1.5 2335 200 n.d. n.d.
CR2-3-2 2.5 2357 223 n.d. n.d.
CR2-3-3 3.5 2375 242 2437 -1.71
CR2-3-4 5.5 2421 268 n.d. n.d.
CR2-3-5 7.5 n.d. 286 2504 -1.85
CR2-3-6 9.5 2505 303 n.d. n.d.
106
CR2-3-7 11.5 2539 315 2548 -1.78
CR2-3-8 14.5 2545 323 n.d. n.d.
CR2-3-9 18.5 2547 334 2583 -1.99
CR2-3-10 22.5 2609 347 n.d. n.d.
CR2-3-11 27.5 2674 354 2717 -2.23
CR2-3-12 32.5 2700 360 n.d. n.d.
CR2-15-1 0.5 2477 186 n.d. n.d.
CR2-15-2 1.5 2409 214 n.d. n.d.
CR2-15-3 3.5 2418 249 2419 -1.60
CR2-15-4 6.5 2451 282 2295 -1.07
CR2-15-5 9.5 2503 322 n.d. n.d.
CR2-15-6 12.5 2533 332 2568 -2.00
CR2-15-7 15.5 2570 334 n.d. n.d.
CR2-15-8 18.5 2594 342 2617 -2.04
CR2-15-9 21.5 2616 348 n.d. n.d.
CR2-15-10 25.5 2644 353 2656 -2.07
CR2-15-11 29.5 2655 358 n.d. n.d.
CR2-15-12 34.5 2684 365 2706 -2.26
CR2-15-13 38.5 2701 369 n.d. n.d.
CR2-3.2-1 1 2478 191 n.d. n.d.
CR2-3.2-2 2 2437 233 2451 -1.73
CR2-3.2-3 4 2432 259 2451 -1.73
CR2-3.2-4 6 2439 298 n.d. n.d.
CR2-3.2-5 9 2498 310 2429 -1.90
CR2-3.2-6 12 2570 326 2545 -2.09
CR2-3.2-7 15 2554 342 n.d. n.d.
CR2-3.2-8 18 2593 345 2568 -1.87
CR2-3.2-9 22 2611 355 n.d. n.d.
CR2-3.2-10 26 2661 355 2497 -2.02
CR2-3.2-11 31 2641 344 n.d. n.d.
CR2-3.2-12 36 2733 367 2727 -2.20
CR3-15-1 2 2428 196 n.d. n.d.
CR3-15-2 3 2450 225 2478 -1.32
CR3-15-3 5 2472 250 2510 -1.43
CR3-15-4 8 2549 267 n.d. n.d.
CR3-15-5 10 2444 272 2577 -1.87
CR3-15-6 12 2558 278 n.d. n.d.
CR3-15-7 15 2562 281 2584 -1.74
CR3-15-8 18 2595 282 n.d. n.d.
CR3-15-9 22 2608 283 2618 -1.81
107
CR3-15-10 25 2612 290 n.d. n.d.
CR3-15-11 29.5 2599 291 2646 -1.86
CR3-15-12 33.5 2632 296 n.d. n.d.
CR3-3-1 0.5 2494 176 n.d. n.d.
CR3-3-2 2.5 2498 229 2520 -1.54
CR3-3-3 4.5 2553 258 n.d. n.d.
CR3-3-4 7.5 2575 267 2580 -1.85
CR3-3-5 10.5 2577 267 2615 -1.86
CR3-3-6 13.5 2601 278 2622 -1.88
CR3-3-7 16.5 2621 276 n.d. n.d.
CR3-3-8 20.5 2560 275 2585 -1.76
CR3-3-9 23.5 2501 279 n.d. n.d.
CR3-3-10 26.5 2540 280 2633 -1.96
CR3-3-11 29.5 2588 282 n.d. n.d.
CR3-3-12 34.5 2634 286 2627 -2.06
CR3-13-1 0.5 2389 141 n.d. n.d.
CR3-13-2 1.5 2377 178 2447 -1.30
CR3-13-3 2.5 2464 227 n.d. n.d.
CR3-13-4 3.5 2485 250 2470 -1.72
CR3-13-5 5.5 2485 263 2510 -1.80
CR3-13-6 7.5 2518 272 n.d. n.d.
CR3-13-7 9.5 2530 277 n.d. n.d.
CR3-13-8 11.5 2518 280 2571 -1.79
CR3-13-9 14.5 2542 280 n.d. n.d.
CR3-13-10 17.5 2538 282 2523 -1.80
CR3-13-11 23 2538 287 n.d. n.d.
CR4-13-1 1 2401 128 n.d. n.d.
CR4-13-2 2 2370 127 2383 -0.60
CR4-13-3 3 2360 135 2323 -0.60
CR4-13-4 5 2460 n.d. 2513 -1.47
CR4-13-5 7 2478 156 2517 -1.15
CR4-13-6 9 2486 156 2477 -1.10
CR5-7-1 1 2504 212 n.d. n.d.
CR5-7-2 2 2545 231 2544 -1.53
CR5-7-3 3 2502 223 n.d. n.d.
CR5-7-4 5 2539 252 2540 -1.39
CR5-7-5 7 2534 263 2498 -1.27
CR5-7-6 9 2628 316 2592 -1.97
CR5-7-7 13 2710 347 2687 -2.05
CR5-7-8 17 2737 348 n.d. n.d.
108
CR5-7-9 22 2842 354 2727 -2.45
CR5-7-10 26 2881 356 n.d. n.d.
CR5-7-11 30 2963 355 2839 -2.64
CR5-7-12 34 2878 357 n.d. n.d.
CR5-14-1 1.5 2522 190 2641 -2.24
CR5-14-2 2.5 2539 217 2487 -1.52
CR5-14-3 3.5 2475 207 n.d. n.d.
CR5-14-4 5.5 2584 264 2678 -2.43
CR5-14-5 7.5 2631 301 2667 -2.35
CR5-14-6 9.5 2694 318 n.d. n.d.
CR5-14-7 12.5 2731 329 2764 -2.81
CR5-14-8 16.5 2755 341 2787 -2.64
CR5-14-9 19.5 2788 340 n.d. n.d.
CR5-14-10 22.5 2829 346 2893 -3.10
CR5-14-11 26.5 2852 337 n.d. n.d.
CR5-14-12 30.5 2837 338 2889 -3.02
CR5-7.2-1 1.5 2507 230 2525 -1.31
CR5-7.2-2 2.5 2611 252 n.d. n.d.
CR5-7.2-3 3.5 2530 275 2579 -1.57
CR5-7.2-4 5.5 2595 308 2650 -2.03
CR5-7.2-5 7.5 2679 320 n.d. n.d.
CR5-7.2-6 10.5 2732 347 2750 -2.28
CR5-7.2-7 12.5 2785 359 n.d. n.d.
CR5-7.2-8 15.5 2798 360 2773 -2.36
CR5-7.2-9 18.5 2791 361 n.d. n.d.
CR5-7.2-10 21.5 2863 363 2782 -2.65
CR5-7.2-11 24.5 2875 365 n.d. n.d.
CR5-7.2-12 27.5 2973 360 2826 -2.78
CR5-7.2-13 30.5 2892 360 n.d. n.d.
CR5-7.2-14 34.5 2897 350 2809 -2.60
CR5-7.2-15 38.5 2920 352 n.d. n.d.

109
In situ porewater data
Appended below is the primary data set used in this dissertation, that is, the in situ collected
porewater from Cocos Ridge sediments. Descriptions of table columns:
No data: “n.d.” stands for “no data.” This is either because that parameter was not measured for
that sample, or because it was measured but eliminated due to a flag on the analytical measurement
or considerable ( > 3 standard deviations) difference from profile trend.
Sample IDs: Sample IDs are named CR (for “Cocos Ridge”) - SIPR deployment # (either
deployment 1 or 2) - window # (from 0-62). Window # 0 indicated bottom water value measured
on a Niskin sample. Window numbers 1-9 are increasing window depths on Blade 1; 10-18 on
Blade 2; 19-27 on Blade 3; 28-36 on Blade 4; 37-45 on Blade 5; 46-54 on Blade 6; 55-62 on
Needles.
Depth: Silica -adjusted depth (cm). See Chapter 2 for full description of depth assignment. Bottom
water samples (indicated with Sample ID ending in 0) are assigned a depth of 0 cm.
Alkalinity: Total Alkalinity (μmol/kg) measured on mvMICA. See Chapter 2 for full description
of methods.
pH: pH measured on mvMICA, calculated at in situ conditions in CO2SYS. See Chapter 2 for
full description of methods.
Silica: Dissolved silicate (μM) measured on spectrophotometer. See Chapter 2 for full description
of methods.
DIC: Dissolved Inorganic Carbon (μmol/kg) measured on Picarro Cavity Ring-Down
Spectrometer. See Chapter 2 for full description of methods.
𝛿
13
C of DIC: 𝛿
13
C of DIC (‰) measured on Picarro Cavity Ring-Down Spectrometer. See
Chapter 2 for full description of methods.
110
Mn
2+
: Dissolved Manganese (μmol/kg) measured on ICP-MS. See Chapter 3 for full description
of methods.
Ca
2+
: Dissolved Calcium (mmol/kg) measured on ICP-OES. See Chapter 3 for full description of
methods.
Sr
2+
: Dissolved Strontium (μmol/kg) measured on ICP-OES. See Chapter 3 for full description of
methods.

111
Table A.3. In situ porewater data collected from SIPR.
Sample
ID
Depth
(cm)
Alkalinity
(μmol/kg) pH
Silica
(μM)
DIC
(μmol/kg)
𝛿
13
C of
DIC (‰)
Mn
2+

(μmol/kg)
Ca
2+

(mmol/kg)
Sr
+2

(μmol/kg)
CR1-1-0 0 2437 7.724 159 2370 -0.02 n.d. 10.31 85.51
CR1-1-02 7.0 2604 n.d. 310 2607 -1.27 0.2 n.d. n.d.
CR1-1-03 8.0 2555 7.521 314 2573 -1.25 2.7 n.d. n.d.
CR1-1-05 10.0 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR1-1-06 18.0 2779 7.571 356 2666 -1.66 28.2 n.d. n.d.
CR1-1-08 28.0 2713 7.636 369 2706 n.d. 35.1 n.d. n.d.
CR1-1-11 6.9 2684 7.613 300 2627 -1.39 0.0 n.d. n.d.
CR1-1-12 7.9 2713 7.607 314 2634 -1.44 0.1 n.d. n.d.
CR1-1-14 10.0 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR1-1-15 17.9 2848 7.697 368 2744 -1.80 33.2 n.d. n.d.
CR1-1-17 27.9 2827 7.721 381 2768 -1.88 49.8 n.d. n.d.
CR1-1-37 3.6 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR1-1-38 4.1 2543 7.531 255 2531 -1.11 0.0 n.d. n.d.
CR1-1-40 5.1 2503 7.533 262 2542 -1.14 0.0 n.d. n.d.
CR1-1-42 10.6 2662 7.585 330 2635 -1.56 14.6 n.d. n.d.
CR1-1-44 12.6 2692 7.619 320 2638 -1.51 17.1 n.d. n.d.
CR1-1-55 1.0 2530 7.592 232 2488 -0.76 0.0 n.d. n.d.
CR1-1-56 3.0 2508 7.582 234 2480 -0.63 0.0 n.d. n.d.
CR1-1-57 5.0 2516 7.628 214 2452 -0.54 0.1 n.d. n.d.
CR1-1-58 7.0 2640 7.698 261 2571 -1.04 24.3 n.d. n.d.
CR1-1-59 9.0 2562 7.727 240 2526 -0.88 21.0 n.d. n.d.
CR1-1-60 11.0 2770 7.688 354 2719 -1.62 37.2 n.d. n.d.
CR1-1-61 15.0 2796 7.641 369 2759 -1.74 28.0 n.d. n.d.
CR1-1-62 19.0 2786 7.570 366 2729 -1.77 15.2 n.d. n.d.
CR1-2-19 7.2 2480 7.489 331 2641 -1.46 0.1 10.43 85.53
CR1-2-21 9.2 2696 7.531 334 2666 -1.43 0.0 10.43 85.78
CR1-2-24 19.2 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR1-2-25 24.2 2862 7.686 379 2741 -1.83 59.3 n.d. n.d.
CR1-2-27 34.2 2813 7.619 390 2759 -2.07 59.6 10.44 85.81
CR1-2-32 8.8 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR1-2-34 14.8 2816 7.674 389 2693 -1.97 71.4 10.39 84.86
CR1-2-46 6.5 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR1-2-49 8.0 2534 7.448 324 2595 -1.44 0.0 10.37 84.79
CR1-2-51 13.5 2749 7.632 374 2661 -1.59 38.9 n.d. n.d.
CR1-2-55 6.1 2548 7.496 305 2559 -1.22 0.0 10.34 84.88
CR1-2-58 12.1 2767 7.651 370 2682 -1.73 48.5 10.41 85.75
112
CR1-2-61 20.1 2814 7.727 384 2722 -1.88 72.2 10.43 85.83
CR1-2-62 24.1 2721 7.709 390 2729 -1.84 75.3 10.40 85.06
CR2-1-0 0 2436 7.742 157 2364 0.09 0.0 10.34 85.58
CR2-1-14 10.8 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR2-1-16 23.8 2791 7.674 375 2714 -1.62 12.0 10.47 85.58
CR2-1-20 7.2 2754 7.668 321 2658 -1.31 0.0 n.d. n.d.
CR2-1-22 9.2 2688 7.674 321 2604 -1.10 n.d. n.d. n.d.
CR2-1-24 18.2 2801 7.690 378 2725 -1.60 15.0 n.d. n.d.
CR2-1-26 28.2 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR2-1-27 33.2 2836 7.657 386 2766 -1.74 16.1 10.51 85.95
CR2-1-29 0.7 2534 7.690 232 2466 -0.66 0.0 n.d. n.d.
CR2-1-32 2.2 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR2-1-35 9.2 2763 7.673 347 2686 -1.44 6.7 n.d. n.d.
CR2-1-55 1.0 2469 n.d. 166 2354 -0.08 0.0 n.d. n.d.
CR2-1-56 3.0 2533 7.727 219 2466 -0.51 0.0 n.d. n.d.
CR2-1-57 5.0 2644 7.692 285 2590 -1.08 0.0 10.49 85.60
CR2-1-58 7.0 2710 7.695 327 2587 -1.35 2.6 n.d. n.d.
CR2-1-59 9.0 2714 7.701 342 2666 -1.34 7.6 n.d. n.d.
CR2-1-60 14.0 2793 7.701 356 2701 -1.45 12.9 n.d. n.d.
CR2-1-62 -3.0 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR2-2-19 4.3 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR2-2-23 8.3 2688 7.662 334 2653 0.14 n.d. 10.50 85.77
CR2-2-26 26.3 2844 7.677 388 2744 -0.33 n.d. n.d. n.d.
CR2-2-37 1.3 2484 7.724 194 2375 1.55 n.d. n.d. n.d.
CR2-2-41 3.3 2647 7.674 283 2626 0.28 n.d. n.d. n.d.
CR2-2-43 9.3 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR2-2-46 2.1 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR2-2-47 2.6 2535 7.704 208 2444 1.30 n.d. 10.38 85.18
CR2-2-48 3.1 2568 7.689 241 2522 0.88 n.d. n.d. n.d.
CR2-2-49 3.6 2505 7.690 275 2606 0.27 n.d. n.d. n.d.
CR2-2-52 10.1 2750 7.666 355 2703 -0.07 n.d. n.d. n.d.
CR2-2-53 11.1 2784 7.684 354 2707 -0.12 n.d. n.d. n.d.
CR2-2-54 12.0 2782 7.673 358 2726 -0.13 n.d. 10.54 86.08
CR2-2-55 0.6 2445 7.732 164 2368 1.73 n.d. 10.26 84.61
CR2-2-56 3.8 2635 7.682 281 2604 0.51 n.d. 10.45 85.64
CR2-2-57 0.8 2440 7.730 165 2373 1.71 n.d. n.d. n.d.
CR2-2-58 1.8 2497 7.743 164 2359 1.75 n.d. n.d. n.d.
CR2-2-59 4.8 2715 7.678 288 2629 0.36 n.d. n.d. n.d.
CR2-2-60 6.8 2666 7.678 327 2686 0.13 n.d. n.d. n.d.
CR2-2-61 2.8 2594 7.691 250 2525 0.68 n.d. n.d. n.d.
113
CR2-2-62 10.8 2862 7.713 354 2714 -0.10 n.d. n.d. n.d.
CR3-1-0 0 2399 7.697 132 2353 -0.09 0.0 10.32 84.96
CR3-1-02 4.4 2617 7.617 275 2583 0.12 1.8 n.d. n.d.
CR3-1-03 5.4 2641 7.628 276 2600 0.03 3.0 10.43 85.38
CR3-1-04 6.4 2666 7.630 291 2625 -0.24 0.0 n.d. n.d.
CR3-1-05 7.4 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR3-1-06 15.4 2742 7.632 312 2677 -0.32 4.0 n.d. n.d.
CR3-1-07 20.4 2748 7.637 312 2700 -0.34 4.3 10.48 86.61
CR3-1-09 30.4 2786 7.647 321 2717 -0.45 3.8 10.51 86.64
CR3-1-12 5.8 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR3-1-13 6.8 2685 7.628 292 2631 -0.06 2.8 n.d. n.d.
CR3-1-17 25.8 2787 7.670 314 2714 -0.28 4.7 n.d. n.d.
CR3-1-28 1.4 2429 7.658 160 2388 0.87 0.0 n.d. n.d.
CR3-1-32 3.4 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR3-1-33 8.4 2687 7.629 287 2626 -0.26 3.8 n.d. n.d.
CR3-1-55 0.0 2418 7.695 138 2358 1.26 0.0 n.d. n.d.
CR3-1-56 0.0 2433 7.695 138 2369 0.99 0.0 n.d. n.d.
CR3-1-57 1.8 2598 7.646 238 2525 0.08 0.2 n.d. n.d.
CR3-1-59 5.8 2620 7.650 258 2576 -0.13 2.0 n.d. n.d.
CR3-1-60 7.8 2699 7.639 280 2642 -0.16 2.3 n.d. n.d.
CR3-1-61 11.8 2754 7.620 289 2671 -0.38 n.d. n.d. n.d.
CR3-1-62 15.8 2726 7.637 290 2648 -0.21 3.4 n.d. n.d.
CR3-2-21 6.2 2664 7.614 294 2665 -1.21 n.d. 10.43 85.31
CR3-2-23 8.2 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR3-2-26 26.2 2757 7.643 314 2730 -1.51 n.d. n.d. n.d.
CR3-2-37 0.9 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR3-2-39 1.9 2511 7.629 214 2516 -0.73 n.d. 10.40 85.23
CR3-2-42 7.9 2664 7.625 298 2649 -1.20 n.d. n.d. n.d.
CR3-2-46 2.1 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR3-2-47 2.6 2501 7.642 206 2496 -0.57 n.d. n.d. n.d.
CR3-2-48 3.1 2524 7.626 220 2514 -0.77 n.d. 10.40 84.75
CR3-2-49 3.5 2562 7.620 235 2545 -0.85 n.d. n.d. n.d.
CR3-2-51 9.1 2684 n.d. 315 2651 -1.20 n.d. n.d. n.d.
CR3-2-52 10.1 2693 n.d. 315 2657 -1.17 n.d. 10.48 85.95
CR3-2-53 11.0 2686 n.d. 317 2643 -1.17 n.d. n.d. n.d.
CR3-2-54 12.0 2687 7.629 316 2647 -1.22 n.d. n.d. n.d.
CR3-2-55 0.0 2434 7.664 137 2346 0.17 n.d. n.d. n.d.
CR3-2-56 0.7 2425 7.689 137 2354 0.30 n.d. n.d. n.d.
CR3-2-57 2.7 2529 7.639 215 2510 -0.65 n.d. n.d. n.d.
CR3-2-58 4.7 2575 7.640 237 2578 -0.86 n.d. n.d. n.d.
114
CR3-2-59 6.7 2619 n.d. n.d. 2602 -1.01 n.d. n.d. n.d.
CR3-2-60 8.7 2642 7.638 281 2619 -1.07 n.d. n.d. n.d.
CR3-2-61 12.7 2657 7.638 288 2629 -1.14 n.d. n.d. n.d.
CR3-2-62 17.2 2688 7.648 293 2655 -1.19 n.d. n.d. n.d.
CR5-1-0 0 2435 7.730 164 2368 0.14 0.0 10.35 85.65
CR5-1-02 8.4 2792 7.596 325 2697 -1.37 0.1 n.d. n.d.
CR5-1-04 10.4 2822 7.647 352 2738 -1.74 23.8 n.d. n.d.
CR5-1-5 11.4 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR5-1-06 19.4 2923 7.759 373 2784 -1.94 71.4 10.63 88.72
CR5-1-07 24.4 2936 n.d. 380 2814 -2.25 100.3 10.61 88.13
CR5-1-08 29.4 2980 7.794 379 2827 -2.34 n.d. 10.64 89.14
CR5-1-09 34.4 3003 7.846 383 2859 -2.43 123.6 10.65 89.38
CR5-1-21 9.5 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR5-1-22 10.5 2790 7.568 336 2727 -1.56 0.1 n.d. n.d.
CR5-1-26 29.5 2969 7.844 378 2836 -2.36 113.6 n.d. n.d.
CR5-1-38 6.1 2689 n.d. 293 2618 -1.18 0.0 n.d. n.d.
CR5-1-40 7.1 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR5-1-43 13.6 2863 7.727 354 2759 -1.76 43.8 n.d. n.d.
CR5-1-55 3.3 2613 7.661 246 2536 -0.86 0.0 10.46 85.98
CR5-1-56 5.3 2697 7.624 313 2636 -1.23 1.6 10.58 86.94
CR5-1-58 9.3 2832 7.745 341 2721 -1.57 45.2 10.54 86.95
CR5-1-60 13.3 2844 7.756 353 2738 -1.69 51.3 n.d. n.d.
CR5-1-61 17.3 2901 n.d. 374 2801 -1.93 68.6 n.d. n.d.
CR5-1-62 21.5 2959 7.764 377 2846 -2.16 82.5 n.d. n.d.
CR5-2-03 7.4 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR5-2-05 9.4 2696 7.677 307 2611 -1.20 n.d. n.d. n.d.
CR5-2-28 6.0 2658 7.660 281 2549 -0.92 n.d. n.d. n.d.
CR5-2-29 6.5 2669 7.639 294 2581 -1.12 n.d. n.d. n.d.
CR5-2-30 7.0 2729 7.630 304 2628 -1.19 n.d. n.d. n.d.
CR5-2-31 7.5 2763 7.628 334 2684 -1.44 n.d. n.d. n.d.
CR5-2-32 8.0 2813 7.606 305 2695 -1.56 n.d. n.d. n.d.
CR5-2-35 15.0 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR5-2-39 6.2 2763 7.631 332 2703 -1.46 n.d. n.d. n.d.
CR5-2-44 14.2 n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d.
CR5-2-55 4.7 n.d. n.d. 257 2560 -0.96 n.d. 10.52 86.66
CR5-2-56 6.7 2730 7.649 316 2678 -1.28 n.d. 10.55 86.36
CR5-2-57 8.7 2771 7.630 334 2715 -1.56 n.d. 10.57 86.53
CR5-2-58 10.7 n.d. n.d. 357 2746 -1.63 n.d. n.d. n.d.
CR5-2-59 12.7 2785 7.730 334 2692 -1.54 n.d. n.d. n.d.
CR5-2-60 14.7 2875 7.782 379 2794 -1.93 n.d. 10.55 87.01
115
CR5-2-61 18.7 2877 7.761 395 2775 -1.93 n.d. n.d. n.d.
CR5-2-62 22.9 2889 7.775 365 2777 -2.04 n.d. n.d. n.d.

Abstract (if available)

Abstract Calcium carbonate plays an unusual role in the global C cycle. Its formation produces CO2, yet the dissolution of calcium carbonate (CaCO3) in the ocean is one of Earth’s natural mechanisms for neutralizing high concentrations of atmospheric carbon dioxide (CO2). Despite the importance of CaCO3 dissolution in long-term climate regulation, there remain large uncertainties in the rate of seafloor dissolution with respect to depth, water column saturation state, surface productivity, and sediment composition. Specifically, the role and magnitude of respiration-driven dissolution is not well understood, in which acid production via aerobic respiration of organic carbon fuels CaCO3 dissolution. Here, I present results from a novel in situ porewater sampler that I led in constructing and testing. It is capable of extracting deep-sea porewater with cm-scale resolution, and my first paper describes its design and provides examples of ex situ sampling artifacts. I then present a study of CaCO3 dissolution on the Cocos Ridge seafloor in the eastern equatorial Pacific through in situ porewater measurements. Dissolution fluxes were calculated with three independent approaches: alkalinity fluxes, stable isotopes of dissolved inorganic carbon (DIC) combined with DIC fluxes, and dissolved calcium fluxes. Although the saturation state of calcite thermodynamically controls whether dissolution is favored, these results show dissolution flux is not correlated with the saturation state of overlying bottom water. Additionally, dissolution is observed in sediments where the water column is saturated with respect to calcite, indicating a process other than water column chemistry is responsible for dissolution. I conclude that the driver of sedimentary dissolution in this region is respiration-driven dissolution. That dissolution occurs within the sediment column, not only at the sediment-water interface, contradicts a paradigm in the field (rapid, water-side control) and will alter paleoceanographic proxy signals. These interpretations underscore the necessity to consider factors other than saturation state, such as organic matter rain and oxygenation when estimating global CaCO3 budgets and assessing the oceans’ response to the current pulse in atmospheric CO2. Via aerobic respiration of organic carbon, sedimentary CaCO3 dissolution has the potential to neutralize anthropogenic CO2 at larger magnitudes than previously estimated.

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University of Southern California Dissertations and Theses

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Asset Metadata

Creator Pittman, Jaclyn Elise (author)

Core Title Carbonate dissolution at the seafloor: fluxes and drivers from a novel in situ porewater sampler

School College of Letters, Arts and Sciences

Degree Doctor of Philosophy

Degree Program Geological Sciences

Degree Conferral Date 2023-08

Publication Date 08/11/2023

Defense Date 05/18/2023

Publisher University of Southern California. Libraries (digital)

Tag alkalinity,carbonate,deep sea,dissolution,dissolved inorganic carbon,OAI-PMH Harvest,oceanography,ph.,porewater,saturation

Language English

Contributor Electronically uploaded by the author (provenance)

Advisor Berelson, William (committee chair), Hammond, Douglas (committee member), Levine, Naomi (committee member)

Creator Email jepittma@usc.edu;jaclynepittman@gmail.com

Permanent Link (DOI) https://doi.org/10.25549/usctheses-oUC113298048

Unique identifier UC113298048

Identifier etd-PittmanJac-12243.pdf (filename)

Legacy Identifier etd-PittmanJac-12243

Document Type Dissertation

Rights Pittman, Jaclyn Elise

Internet Media Type application/pdf

Type texts

Source 20230814-usctheses-batch-1084 (batch), University of Southern California (contributing entity), University of Southern California Dissertations and Theses (collection)

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Repository Name University of Southern California Digital Library

Repository Location USC Digital Library, University of Southern California, University Park Campus, Los Angeles, California 90089, USA

Repository Email cisadmin@lib.usc.edu