1.
Introduction
The marine calcium carbonate (CaCO
3
) cycle is integral to the global carbon cycle. The production of biogenic
CaCO
3
tends to raise atmospheric CO
2
due to consumption of surface ocean alkalinity, while the ballasting
of organic matter and export into the deep ocean provided by this material tends to lower CO
2
(De La Rocha
Abstract
The cycling of biologically produced calcium carbonate (CaCO
3
) in the ocean is a fundamental
component of the global carbon cycle. Here, we present experimental determinations of in situ coccolith and
foraminiferal calcite dissolution rates. We combine these rates with solid phase fluxes, dissolved tracers, and
historical data to constrain the alkalinity cycle in the shallow North Pacific Ocean. The in situ dissolution
rates of coccolithophores demonstrate a nonlinear dependence on saturation state. Dissolution rates of all
three major calcifying groups (coccoliths, foraminifera, and aragonitic pteropods) are too slow to explain the
patterns of both CaCO
3
sinking flux and alkalinity regeneration in the North Pacific. Using a combination of
dissolved and solid-phase tracers, we document a significant dissolution signal in seawater supersaturated for
calcite. Driving CaCO
3
dissolution with a combination of ambient saturation state and oxygen consumption
simultaneously explains solid-phase CaCO
3
flux profiles and patterns of alkalinity regeneration across the
entire N. Pacific basin. We do not need to invoke the presence of carbonate phases with higher solubilities.
Instead, biomineralization and metabolic processes intimately associate the acid (CO
2
) and the base (CaCO
3
)
in the same particles, driving the coupled shallow remineralization of organic carbon and CaCO
3
. The linkage
of these processes likely occurs through a combination of dissolution due to zooplankton grazing and microbial
aerobic respiration within degrading particle aggregates. The coupling of these cycles acts as a major filter on
the export of both organic and inorganic carbon to the deep ocean.
Plain Language Summary
The marine carbon cycle is made of organic carbon and calcium
carbonate (CaCO
3
) components. While the organic carbon cycle has received much attention, the CaCO
3
cycle
is relatively understudied. Through a dedicated research expedition to the North Pacific Ocean, we demonstrate
here a coupling of these two cycles, stemming from the fact that all organisms that produce CaCO
3
also produce
intimately associated organic carbon. We suggest that the mechanisms responsible for the formation and sinking
of organic carbon particles in the ocean are likely as important for CaCO
3
export, and that the respiration of
organic carbon is responsible for the dissolution of a substantial portion of CaCO
3
in the upper ocean.
SUBHAS ET AL.
© 2022. American Geophysical Union.
All Rights Reserved.
Shallow Calcium Carbonate Cycling in the North Pacific
Ocean
Adam V. Subhas
1
, Sijia Dong
2
, John D. Naviaux
3
, Nick E. Rollins
4
, Patrizia Ziveri
5,6
,
William Gray
7
, James W. B. Rae
8
, Xuewu Liu
9
,
Robert H. Byrne
9
, Sang Chen
10
, Christopher Moore
9
, Loraine Martell-Bonet
9
,
Zvi Steiner
11
, Gilad Antler
12
, Huanting Hu
10
, Abby Lunstrum
4
, Yi Hou
13
,
Nathaniel Kemnitz
4
, Johnny Stutsman
14
, Sven Pallacks
5
, Mathilde Dugenne
15
,
Paul D. Quay
14
, William M. Berelson
4
, and Jess F. Adkins
2
1
Department of Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole, MA, USA,
2
Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA, USA,
3
Four Twenty
Seven Inc., Sunnyvale, CA, USA,
4
Department of Earth Sciences, University of Southern California, Los Angeles, CA, USA,
5
Institut de Ciencia I Tecnologia Ambientals, Universitat Autonoma de Barcelona, Barcelona, Spain,
6
Catalan Institution
for Research and Advanced Studies (ICREA), Barcelona, Spain,
7
Laboratorie des Sciences du Climat et de l’Environnement
(LSCE/IPSL), Universite Paris-Saclay, Gif-sur-Yvette, France,
8
School of Earth and Environmental Sciences, University of St
Andrews, St Andrews, Scotland,
9
College of Marine Science, University of South Florida, St. Petersburg, FL, USA,
10
School
of Oceanography, Shanghai Jiao Tong University, Shanghai, China,
11
GEOMAR, Helmholtz Centre for Ocean Research Kiel,
Kiel, Germany,
12
Department of Earth and Environmental Sciences, Ben-Gurion University, Be'er Sheva, Israel,
13
Department
of Earth, Environmental, and Planetary Sciences, Rice University, Houston, TX, USA,
14
School of Oceanography, University
of Washington, Seattle, WA, USA,
15
School of Ocean and Earth Science and Technology, University of Hawaii, Honolulu, HI,
USA
Key Points:
•
High resolution carbonate chemistry,
δ
13
C-DIC, and particle flux
measurements in the NE Pacific sheds
light on the upper ocean calcium
carbonate and alkalinity cycles
•
Based on this sampling campaign,
there is evidence for substantial
CaCO
3
dissolution in the mesopelagic
zone above the saturation horizon
•
Dissolution experiments, observations,
and modeling suggest that shallow
CaCO
3
dissolution is coupled to
the consumption of organic carbon,
through a combination of zooplankton
grazing and oxic respiration within
particle microenvironments
Supporting Information:
Supporting Information may be found in
the online version of this article.
Correspondence to:
A. V. Subhas,
asubhas@whoi.edu
Citation:
Subhas, A. V., Dong, S., Naviaux, J.
D., Rollins, N. E., Ziveri, P., Gray,
W., et al. (2022). Shallow calcium
carbonate cycling in the North Pacific
Ocean.
Global Biogeochemical Cycles
,
36
, e2022GB007388.
https://doi.
org/10.1029/2022GB007388
Received 14 MAR 2022
Accepted 28 APR 2022
Author Contributions:
Conceptualization:
Adam V. Subhas,
William M. Berelson, Jess F. Adkins
Data curation:
Adam V. Subhas, Sijia
Dong, John D. Naviaux, Nick E. Rollins,
Sang Chen
Formal analysis:
Adam V. Subhas, Sijia
Dong, John D. Naviaux, Patrizia Ziveri,
William Gray, James W. B. Rae, Zvi
Steiner, Paul D. Quay, Jess F. Adkins
Funding acquisition:
William M.
Berelson, Jess F. Adkins
Investigation:
Adam V. Subhas, Sijia
Dong, John D. Naviaux, Nick E. Rollins,
Patrizia Ziveri, William Gray, James W.
B. Rae, Xuewu Liu, Robert H. Byrne,
Sang Chen, Christopher Moore, Loraine
10.1029/2022GB007388
RESEARCH ARTICLE
1 of 22
Global Biogeochemical Cycles
SUBHAS ET AL.
10.1029/2022GB007388
2 of 22
et al.,
2008
; Klaas & Archer,
2002
; Passow & De La Rocha,
2006
). In addition to these roles, solid CaCO
3
is
crucial to the neutralization of CO
2
through its dissolution and associated production of ocean alkalinity (Archer
et al.,
1998
). Despite consensus on the general dynamics of marine CaCO
3
cycling, rates of CaCO
3
production
and dissolution are poorly constrained (Battaglia et al.,
2016
; Berelson et al.,
2007
; Dunne et al.,
2012
). If a
substantial amount of alkalinity is regenerated in the shallow ocean where precipitation is thermodynamically
favored, then the traditional relationship between water column CaCO
3
dissolution rates and mineral saturation
state must be reexamined.
The formation and dissolution of calcium carbonate minerals is canonically described as a function of seawater
saturation state
퐴퐴
(
Ω=
[
Ca
2+
][
CO
2−
3
]
∕K
′
sp
)
, where
퐴퐴
K
′
sp
is the in situ apparent solubility product of the CaCO
3
mineral of interest (e.g., calcite or aragonite). The dissolution rate
R
is empirically related to
Ω
using the equation
R = k
(1−
Ω
)
n
, where
k
is the specific dissolution rate constant, and
n
is a fit parameter describing the nonlinearity
of dissolution rate as a function of undersaturation (Keir,
1980
). In brief, calcite dissolution is a highly nonlinear
function of seawater undersaturation across a wide range of saturation states, an observation that dates back to
early experiments by Peterson (
1966
) and Berner and Morse (
1974
).
This high degree of nonlinearity can be explained by applying kinetic crystal growth models to calcite dissolu
-
tion, as summarized in Adkins et al. (
2021
) for inorganic calcite. As saturation state decreases, inorganic calcite
dissolution goes through distinct rate transitions at “critical” Omega values (
Ω
crit
, Dong, Berelson, Adkins,
et al.,
2020
; Naviaux, Subhas, Rollins, et al.,
2019
). At colder temperatures, a single break in slope in log(rate)
versus log(1−
Ω
) is observed at an
Ω
crit
∼0.75–0.8, marking the transition between dissolution at step-edges
and dissolution on the 2D mineral surface (Figure
1
black envelope, Dong et al.,
2019
; Naviaux, Subhas, Dong,
et al.,
2019
; Naviaux, Subhas, Rollins, et al.,
2019
; Peterson,
1966
). Practically, two dissolution regimes mean that
calcite dissolution in seawater is best described by two rate equations, uniquely defined between 1 >
Ω
>
Ω
crit
and
Ω
crit
>
Ω
> 0, each with distinct
k
−
n
pairs. Near equilibrium, dissolution rate is slow and relatively insensitive
to saturation state with a small
k
and
n
< 1. Far from equilibrium, dissolution rapidly increases as a function of
undersaturation, leading to a large
k
and
n
>> 1. Several in situ water column dissolution experiments with inor
-
ganic calcite have documented this sharp increase in dissolution rate below
Ω
crit
(Honjo & Erez,
1978
; Naviaux,
Subhas, Dong, et al.,
2019
; Peterson,
1966
).
Like inorganic calcite, biogenic calcites and aragonite demonstrate a rate transition at the same
Ω
crit
as inorganic
calcite (Dong et al.,
2019
; Subhas et al.,
2018
). Biogenic calcite dissolution rates demonstrate
k
and
n
values
that are distinct from the inorganic calcite curve (Figure
1
, red envelope). Coccolith dissolution rates appear
to behave similarly to calcite near equilibrium, but are less sensitive to saturation state far-from-equilibrium
(
n
= 2.2, Subhas et al.,
2018
). Foraminifera also demonstrate a far-from-equilibrium
n
∼ 2. They appear to
dissolve faster than coccoliths normalized by both mass and by surface area, suggesting that foraminifera require
their own specific
k
values (Gehlen et al.,
2005
; Subhas et al.,
2018
; Subhas, McCorkle, et al.,
2019
). Aragonite
dissolves more slowly than biogenic calcite, and also has the lowest
n
∼ 1.8 (Dong et al.,
2019
). Compared to
inorganic calcite, the lower sensitivity of biogenic calcite dissolution to saturation state, particularly farther from
equilibrium, is potentially due to the presence of organic templates and matrices within the calcite structure
(e.g., Subhas et al.,
2018
; Walker & Langer,
2021
). Differences in specific
k
and
n
values aside, it appears that
all calcium carbonates dissolve slowly near equilibrium, and only rapidly increase their dissolution rates once
Ω
drops below
Ω
crit
∼ 0.8 for the colder temperatures of the modern water column.
Contrary to these experiments, most oceanographic observations of CaCO
3
dissolution argue for a large dissolu-
tion flux in the shallow ocean, where waters are typically supersaturated, followed by relatively little dissolution
in the deep ocean, where waters are more deeply undersaturated. The mechanisms driving this large shallow
dissolution flux are confusing, because they cannot be explained by rate relationships between saturation state
and dissolution rate. This conundrum was posed explicitly by Milliman et al. (
1999
), who postulated that balanc
-
ing global CaCO
3
production and burial required a large shallow dissolution flux. This shallow dissolution flux
was then matched to the appearance of excess alkalinity in the Pacific (Feely et al.,
2002
) and Atlantic basins
(Chung et al.,
2003
). Later, some researchers argued that the appearance of excess alkalinity above the calcite
saturation horizon is complicated by water mass transport and mixing processes (Battaglia et al.,
2016
; Carter
et al.,
2021
; Friis et al.,
2006
). Recent modeling efforts have shown that a supersaturated alkalinity signal is due
to a combination of circulation and dissolution, lending support to the presence of a shallow CaCO
3
cycle (Carter
et al.,
2021
; Sulpis et al.,
2021
).
Martell-Bonet, Zvi Steiner, Gilad Antler,
Huanting Hu, Abby Lunstrum, Yi Hou,
Nathaniel Kemnitz, Johnny Stutsman,
Sven Pallacks, Mathilde Dugenne, Paul
D. Quay, William M. Berelson
Methodology:
Adam V. Subhas, John D.
Naviaux, Nick E. Rollins, Patrizia Ziveri,
Xuewu Liu, Robert H. Byrne, William M.
Berelson, Jess F. Adkins
Project Administration:
Adam V.
Subhas, William M. Berelson, Jess F.
Adkins
Resources:
Robert H. Byrne, William M.
Berelson
Supervision:
William M. Berelson
Validation:
Patrizia Ziveri
Visualization:
Adam V. Subhas
Writing – original draft:
Adam V.
Subhas
Writing – review & editing:
Adam V.
Subhas, Sijia Dong, John D. Naviaux,
Patrizia Ziveri, William Gray, James W.
B. Rae, Robert H. Byrne, Zvi Steiner,
Gilad Antler, Abby Lunstrum, Nathaniel
Kemnitz, Sven Pallacks, Mathilde
Dugenne, William M. Berelson, Jess F.
Adkins
19449224, 2022, 5, Downloaded from https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2022GB007388 by California Inst of Technology, Wiley Online Library on [06/10/2023]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Global Biogeochemical Cycles
SUBHAS ET AL.
10.1029/2022GB007388
3 of 22
Numerous observations of particulate CaCO
3
loss in the upper water
column serve as a complement to the appearance of dissolved alkalinity
(Barrett et al.,
2014
; Bishop & Wood,
2008
; Bishop et al.,
1980
,
1986
; Dong
et al.,
2019
; Milliman et al.,
1999
; Timothy et al.,
2013
; Troy et al.,
1997
;
Wong et al.,
1999
). Some authors have attempted to explain these observa
-
tions through the production and dissolution of more soluble CaCO
3
poly
-
morphs such as aragonite (Buitenhuis et al.,
2019
; Feely et al.,
2002
), and
pelagic fish-produced Mg-calcite, and amorphous carbonates (S. E. Wilson
et al.,
2008
; Woosley et al.,
2012
). The total production of these “missing”
highly soluble phases remains poorly constrained, and therefore their contri
-
bution to the shallow CaCO
3
cycle remains an open question.
An alternate hypothesis is that the aggregation of marine snow and the graz
-
ing activity of zooplankton create hotspots of calcium carbonate dissolution
(Bishop et al.,
1980
; Milliman et al.,
1999
). The interiors of marine snow
aggregates and zooplankton guts both exhibit lower pH, and therefore lower
Ω
, than ambient seawater, due to the production of respiratory CO
2
and
digestive acids (Alldredge & Cohen,
1987
; Pond et al.,
1995
). These graz
-
ing and aggregation processes are recognized as fundamental to total carbon
export from the euphotic zone through the mesopelagic (Boyd et al.,
2019
;
Grabowski et al.,
2019
; Henson et al.,
2019
), and they would allow more
abundant phases like coccolith and foraminiferal calcite to play a role in shal
-
low water column dissolution.
Here, we present new results and analyses from the CDisK-IV research expe
-
dition, occupying five stations in the North Pacific from Hawaii to Alaska in
the summer of 2017. We present new in situ dissolution rate measurements
of coccolith and foraminiferal calcite using a
13
C-tracer approach (Naviaux,
Subhas, Dong, et al.,
2019
; Subhas et al.,
2015
). The quantification of in situ
dissolution rates across a wide range of saturation states allows us to estimate
the magnitude of ambient
Ω
-driven water column CaCO
3
dissolution, with
explicit contributions from the three main calcifier groups (coccoliths, foraminifera, and pteropods) to the total
CaCO
3
dissolution flux. In order to assess the significance of dissolution above the saturation horizon, we present
new analyses of dissolved water column carbonate chemistry and oxygen data that complement the canonical
excess alkalinity approach. Finally, we attempt to reconcile these in situ dissolution rate measurements and our
own analyses with historical observations of a large, shallow dissolution cycle in the North Pacific Ocean.
2.
Materials and Methods
2.1.
Study Area
The North Pacific is a key region for constraining CaCO
3
cycling, due to the large portion of the water column
that is undersaturated for calcite and aragonite (Figure
2a
), and the foundational studies on changes in saturation
horizon and degree of carbonate dissolution performed there (Feely et al.,
2002
; Friis et al.,
2006
; Peterson,
1966
;
Sabine et al.,
2004
). We conducted the CDisK-IV cruise on the R/V Kilo Moana in the North Pacific Ocean
during August 2017, on a transect from Honolulu, HI to Seward, AK. At five survey stations (Figure
2
gray
points, Table
1
) we measured an extensive range in saturation states (
Ω
, defined as
퐴퐴
[
Ca
2+
][
CO
2−
3
]
∕K
′
sp
,
where
퐴퐴
K
′
sp
is the apparent solubility of the carbonate mineral; Figure
2a
), and established a spatial gradient in CaCO
3
production ecology. Quantitative CaCO
3
mineralogy, ecology, standing stock, production, and export flux meas
-
urements along this transect are presented elsewhere (Dong et al.,
2019
; Ziveri et al.,
2022
). We complemented
these solid-phase measurements by sampling the full water column for its dissolved inorganic carbon (DIC), total
alkalinity (Alk), pH, and δ
13
C of DIC; and deployed custom-built in situ incubators to measure biogenic calcite
dissolution rates as a function of water column chemistry. The dissolution rates of aragonite and inorganic calcite
are presented in detail elsewhere (Dong et al.,
2019
; Naviaux, Subhas, Dong, et al.,
2019
).
Figure 1.
A schematic of surface area-normalized log(rate) as a function
of log(1−
Ω
) at 5°C, for inorganic calcite (black) and aragonite and biogenic
calcites (red), adapted from Naviaux, Subhas, Dong, et al. (
2019
), Naviaux,
Subhas, Rollins, et al. (
2019
), Subhas et al. (
2018
), and Dong et al. (
2019
).
The dashed line denotes
Ω
crit
, the value of
Ω
at which the dissolution
mechanism shifts from near-equilibrium step-edge retreat to far-from-
equilibrium 2-dimensional, homogenous dissolution. All biogenic calcites, and
aragonite, demonstrate a rate transition at the same
Ω
crit
as inorganic calcite,
but demonstrate a shallower far-from-equilibrium slope compared to inorganic
calcite. The larger red envelope is intended to illustrate the larger range in
dissolution rates of biogenic materials. Logarithmically spaced ticks for 1−
Ω
are shown at the top of the plot for clarity.
log (1-
Ω
)
log (Rate)
Inorganic Calcite
Aragonite and
Biogenic Calcites
n~4.5
n~1.5-2.5
Ω
crit
2D Dissolution
Step-edge retreat
n~0.2
1 >
Ω
>
Ω
crit
Ω
crit
>
Ω
> 0
Increasing undersaturation
Increasing rate
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SUBHAS ET AL.
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4 of 22
2.2.
Sediment Traps
Sediment trap protocols can be found in companion papers by Subhas, Adkins, et al. (
2019
) and Dong et al. (
2019
)
and are outlined briefly here. Two sediment traps positioned at 100 and 200 m on a single line were deployed at
each station as a floating array. Each trap consisted of 12 polycarbonate particle interceptor tubes (70 cm long,
10 cm diameter) mounted on a circular frame, with a baffle grid to screen for macrofauna (Haskell et al.,
2016
,
Figure S1 in Supporting Information
S1
for picture). Sinking material was funneled into a 50 mL falcon tube
at the base of each tube. Falcon tubes were filled with 30 mL of buffer solution, prepared following US JGOFS
protocol. Six random tubes were combined, and “swimmers” (mostly large copepods, amphipods and larvae)
were picked out. The samples were filtered onto a pre-weighed glass fiber filter (Whatman glass GF/F, 1825-
047), which was then used to calculate mass flux and mineralogy via X-ray diffraction. Particulate inorganic
carbon (PIC) and total carbon fluxes were calculated using established methods (Dong et al.,
2019
).
Figure 2.
Carbonate chemistry of the upper 1,000 m of the North Pacific Ocean. (a) Depth-latitude zonal section of in situ
Ω
calculated using Alk and dissolved
inorganic carbon data from the GLODAPv2.2019 database. Color bar indicates
Ω
calcite
. Red and black dashed lines denote the aragonite and calcite saturation horizons
(defined as
Ω
= 1), respectively. Gray points indicate sampling locations during our 2017 research expedition (Table
1
). Yellow stars indicate the location of North
Pacific Intermediate Water (main text for details). (b) A calculation of
Ω
met
for calcite along the section in (a). The 0.75 calcite “isosat” is shown to demonstrate the
potential for deep undersaturation within confined environments in the upper water column.
Ambient
Ω
A
Ω
Ca
= 1
Ω
Ar
= 1
“Metabolic”
Ω
met
B
Ω
Ca
= 1
Ω
Ca
= 0.75
Ω
Ar
= 1
Station
Latitude
Longitude
Date occupied
F
PIC,100 m
F
PIC,200 m
f
arag
f
cocco
1
a
22 45′
157 59′
2–4 August 2017
0.71
N.A.
0.44
0.98
2
27 45′
155 15′
6–10 August 2017
0.50
0.32
0.75
0.99
3
35 16′
150 59′
12–15 August 2017
0.13
N.A.
0.50
1.00
4
41 45′
148 16′
16–19 August 2017
2.23
0.82
0.10
0.83
5
b
49 50′
149 39′
21–26 August 2017
1.03
0.75
0.09
0.90
Note.
Floating sediment trap PIC fluxes (
F
) are in units of mmol m
−2
d
−1
, subscript is trap depth. PIC fluxes, and the fraction of total calcium carbonate as aragonite
(
f
arag
), are from Dong et al. (
2019
). The fraction of calcite is calculated as 1−
f
arag
. The mixed layer calcite inventory is further split into coccolith and foraminiferal
fractions, with
f
cocco
representing the coccolith calcite fraction (Ziveri et al.,
2022
).
a
Station 1 was located at Station ALOHA (22 45′N, 158′W).
b
Station 5 was located near Ocean Station Papa (OSP; 50 06′N, 144 54′W).
Table 1
Station Locations and Flux Data for the CDisK-IV Cruise
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5 of 22
2.3.
Dissolved Carbonate System and Nutrient Measurements
Methods for total alkalinity, pH, δ
13
C
DIC
, silica, and nutrient measurements are well established, and have been
reported previously (Dong et al.,
2019
; Naviaux, Subhas, Dong, et al.,
2019
). DIC and secondary measure
-
ments of δ
13
C were measured on Niskin rosette samples using a custom Picarro-based system developed at the
University of South Florida (USF) and at University of Southern California (USC). Samples were collected in
300 ml biological oxygen demand (BOD) bottles in a manner similar to that used for alkalinity samples, with
the exception that the bottle had a 5 ml headspace and was poisoned using a saturated HgCl
2
solution. The DIC
instrument was housed in a 15 gallon insulated cooler that was capable of accommodating 10 samples including
a certified reference material (CRM). After a brief tubing rinse, the sample was pumped from the BOD bottle
to a 20 ml glass bulb that was submerged in a thermostatted water bath. An acid pump injected 3 ml of 17%
H
3
PO
4
into the top of the glass bulb. A three-way valve was then switched, allowing carrier gas to push the acid
-
ified sample into the purging tube, and the CO
2
gas evolved was measured on a Picarro 2131-i instrument. The
software first recorded a baseline reading while N
2
was flowing in the system, and a threshold signal of 70 ppm
triggered the peak detection. Once a signal was detected, the peak integration continued until the signal returned
to the baseline. The DIC purging time was set to be longer than the required integration time to ensure all CO
2
was detected. Total counts of both
12
C and
13
C were obtained. The
13
C/
12
C ratio was calculated based on total
peak integrations. Since ratios are based on total
12
C and
13
C, potential fractionation of carbon during purging is
avoided. Sample
13
C/
12
C ratios and total DIC concentrations were linked to measured CRM solutions, and agree
within error to Niskin δ
13
C values measured by the CalTech/USC group (Naviaux, Subhas, Dong, et al.,
2019
).
We use the Caltech/USC δ
13
C and USF DIC values here.
2.4.
Biogenic CaCO
3
Dissolution Experiments
Dissolution experiments were conducted in situ using modified Niskin incubators, described in detail by Naviaux,
Subhas, Dong, et al. (
2019
) and Dong et al. (
2019
). In this study, we report dissolution rates of bleached,
13
C-la-
beled
E. huxleyi
liths. A total of 20 coccolith dissolution experiments were conducted at depths between 240
and 1,000 m at Stations 2–5 with temperatures ranging from 2.4 to 4.8°C and
Ω
calcite
from 0.96 to 0.67. We
conducted one experiment with a planktic foraminiferal assemblage, cultured and
13
C-labeled as described in
Subhas et al. (
2018
). Roughly 0.5–1.5 mg of labeled biogenic calcite was sealed in between 47 mm diameter
“Nuclepore” polycarbonate membrane filters (0.8 μm pore size). This preparation has no effect on net disso
-
lution rates, as documented by Naviaux, Subhas, Dong, et al. (
2019
). These packets were then mounted inside
the Niskin incubators. The incubators were hung on a hydrowire, sent down to depth, and triggered closed. As a
closed system, a pump continuously circulated water from the housing holding the packets into the Niskin. The
Niskin reactors remained closed at depth for 24–58 hr and were sampled for silica, soluble reactive phosphorus
(SRP), nitrate, alkalinity, pH, and δ
13
C-DIC upon recovery. Niskin data were quality checked by comparing SRP,
silica, and nitrate to ambient water-column values obtained via CTD/rosette deployments on the same cruise.
Saturation states in the Niskin reactors were determined from Alk-pH pairs, input into CO2SYS along with the
temperature, salinity, depth, SRP, and silica concentrations at which the reactor was deployed. Dissolution rates
were calculated by taking the difference between the final incubator and ambient water column
13
C/
12
C ratios,
multiplied by the [DIC] and the mass of seawater (1.126 kg) inside the incubators, and divided by the incubation
time (Naviaux, Subhas, Dong, et al.,
2019
). Biogenic materials were not 100% labeled and rates were scaled for
the extent of isotope labeling following Subhas et al. (
2018
). Briefly, when the amount of
13
C in the dissolving
material is greatly enriched above natural abundance (
13
C/
12
C ∼ 0.01), isotope ratio differences can be multiplied
by a correction factor of (
R
s
+ 1)/
R
s
, where
R
s
is the
13
C/
12
C ratio of the dissolving solid (i.e., a reduced form of
Equation 3 from Subhas et al.,
2018
). One batch of
13
C-labeled
E. huxleyi
was used for all dissolution rates shown
here (
R
s
= 0.928), except for one experiment that used an older batch (
R
s
= 20, Subhas et al.,
2018
). The planktic
foraminifera assemblage from Subhas et al. (
2018
) was used (
R
s
= 1.6). Adjusting the rates using these
R
s
values
gives mass-normalized rates in units of g CaCO
3
g
−1
d
−1
. Mass-normalized dissolution rates were further divided
by the specific surface areas of
E. huxlyei
liths (105,000 cm
2
g
−1
) and planktic foraminifera (43,000 cm
2
g
−1
,
Subhas et al.,
2018
) and the molar mass of calcium carbonate (100 g mol
−1
) to generate specific dissolution rates
for each calcite type in units of mol CaCO
3
cm
−2
d
−1
.
19449224, 2022, 5, Downloaded from https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2022GB007388 by California Inst of Technology, Wiley Online Library on [06/10/2023]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License
Global Biogeochemical Cycles
SUBHAS ET AL.
10.1029/2022GB007388
6 of 22
2.5.
1-D CaCO
3
Flux and Dissolution Calculations
We constructed a one-dimensional model of particulate CaCO
3
sinking and dissolution, combining our measure
-
ments of CaCO
3
export fluxes and export mineralogy (Table
1
); ambient seawater
Ω
; and aragonite, foraminiferal,
and coccolith dissolution kinetics (Table
2
). This model extends the analysis of Dong et al. (
2019
) to coccolith,
foraminifera, and aragonite dissolution at all five survey stations. The model was initiated with our measured
100 m sinking PIC fluxes, partitioned into aragonite and calcite:
퐹퐹
arag
∼
퐹퐹
tot
푓푓
arag
;
퐹퐹
δ −δ
∼
퐹퐹
tot
(
1−
푓푓
arag
)
,
where
F
tot
is the total PIC export flux at 100 m, and
f
arag
is the fraction of total CaCO
3
as aragonite measured in
the sinking flux (Table
1
, Dong et al.,
2019
). The dissolution flux for each phase was then calculated at all depths
for which we acquired water column dissolution data (at least 24 vertical points, usually more) in a similar way
to Dong et al. (
2019
):
퐹
푧푖
=
퐹
푧푖
−1
1−
푅
diss
(
푓
(
Ω
푧푖
−1
))
푧
푖
−
푧
푖
−1
푤
,
where
F
zi
is the mineral flux at depth
z
i
,
R
diss
is the dissolution rate of the mineral and is a function of the meas
-
ured mineral saturation state at depth
z
i
−1
, and
w
is the particle sinking rate. For aragonite fluxes (
F
arag
) and
dissolution (
R
arag
), we use the aragonite dissolution kinetic rate law generated from in situ data on the same cruise
(Dong et al.,
2019
, Table
2
). For calcite fluxes (
F
calc
), we assume that the calcite rain is exclusively foraminifera
and coccoliths (
f
cocco
+ f
foram
= 1, Table
1
). The dissolution rate of coccoliths,
R
cocco
, uses the parameters derived
from the data in Figure
3
, described in Table
2
. The dissolution rate of foraminifera,
R
foram
, is assumed to follow
the same functional form as the dissolution rate of coccoliths,
R
cocco
, but dissolves a factor of 1.6 faster, when
normalized by mass (
R
foram
=
1.6
R
cocco
; Figure
3
, Table
2
, Section
3.2
). This assumption yields a calcite dissolu
-
tion rate of:
푅
calc
=
푓
foram
푅
foram
+
푓
cocco
푅
cocco
;
=
(
1−
푓
cocco
)
1
.
6
푅
cocco
+
푓
cocco
푅
cocco
;
=
(
1
.
6−0
.
6
푓
cocco
)
푅
cocco
.
Errors in regression parameters were propagated through the dissolution model (shaded envelopes in Figure
3a
)
by using the linearized log-log fits as “linear model” variables in MATLAB and the “predict” function. Aragonite
and calcite fluxes were then summed together to calculate the total PIC flux at each depth.
The appearance of regenerated alkalinity, normalized to the mass of seawater (e.g., Battaglia et al.,
2016
; Carter
et al.,
2014
,
2021
; Feely et al.,
2002
) is linked to the dissolution of particulate CaCO
3
. Although Feely et al. (
2002
)
divided their TA* tracer by 2 to convert into moles of CaCO
3
dissolved, we leave all tracer results in units of
1 >
Ω
> 0.78
0.78 >
Ω
> 0
log
10
k
mass
g g
−1
d
−1
log
10
k
sa
mol cm
−2
s
−1
n
log
10
k
mass
g g
−1
d
−1
log
10
k
sa
mol cm
−2
s
−1
n
E. hux
lab 21°C
a
−1.8 ± 0.1
−13.7 ± 0.1
0.33 ± 0.09
−0.31 ± 0.1
−12.4 ± 0.1
2.2 ± 0.1
E. hux
in situ
−2.5 ± 0.1
−14.5 ± 0.1
0.18 ± 0.16
−1.3 ± 0.3
−13.3 ± 0.3
2.1 ± 0.5
Foram assemblage lab 21°C
a
–
–
–
−0.35 ± 0.05
−11.9 ± 0.2
1.7 ± 0.1
E. hux
in situ × 1.6
−2.32 ± 0.14
−14.28 ± 0.14
0.18 ± 0.16
−1.1 ± 0.3
−12.8 ± 0.3
2.1 ± 0.5
Aragonite in situ
b
−2.74 ± 0.19
−14.02 ± 0.19
0.24 ± 0.19
−1.75 ± 0.15
−13.03 ± 0.15
1.76 ± 0.36
Note.
We report
k
mass
(dissolution rate constant normalized to mass),
k
sa
(
k
mass
further normalized to specific surface area) and
n
(reaction exponent) for the regions near
equilibrium above
Ω
crit
(=0.78), and far from equilibrium below
Ω
crit
. To estimate foraminiferal dissolution rates, the in situ mass-normalized
E. hux
data was multiplied
by a factor of 1.6 (see text for details).
a
Laboratory data are from Subhas et al. (
2018
).
b
Aragonite field data are from Dong et al. (
2019
).
Table 2
Regression Parameters for the Rate Equation Rate = k(1−Ω)
n
for the Dissolution Rate Data Shown in Figure
3
, and Used in the Dissolution Rate Model (Section
2.4
)
19449224, 2022, 5, Downloaded from https://agupubs.onlinelibrary.wiley.com/doi/10.1029/2022GB007388 by California Inst of Technology, Wiley Online Library on [06/10/2023]. See the Terms and Conditions (https://onlinelibrary.wiley.com/terms-and-conditions) on Wiley Online Library for rules of use; OA articles are governed by the applicable Creative Commons License