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Potential role of giant marine diatoms in sequestration of atmospheric CO 2
during the Last Glacial Maximum: δ13C evidence from laminated Ethmodis-cus rex mats in tropical West Pacific
Zhifang Xiong, Tiegang Li, Xavier Crosta, Thomas Algeo, FengmingChang, Bin Zhai
PII: S0921-8181(13)00144-6DOI: doi: 10.1016/j.gloplacha.2013.06.003Reference: GLOBAL 1993
To appear in: Global and Planetary Change
Received date: 11 November 2012Accepted date: 7 June 2013
Please cite this article as: Xiong, Zhifang, Li, Tiegang, Crosta, Xavier, Algeo, Thomas,Chang, Fengming, Zhai, Bin, Potential role of giant marine diatoms in sequestrationof atmospheric CO2 during the Last Glacial Maximum: δ13C evidence from laminatedEthmodiscus rex mats in tropical West Pacific, Global and Planetary Change (2013), doi:10.1016/j.gloplacha.2013.06.003
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Potential role of giant marine diatoms in sequestration of
atmospheric CO2 during the Last Glacial Maximum: δ13
C
evidence from laminated Ethmodiscus rex mats in tropical
West Pacific
Zhifang Xiong a, Tiegang Li
a, *, Xavier Crosta
b, Thomas Algeo
c,
Fengming Chang a
, Bin Zhai d
a Key Laboratory of Marine Geology and Environment, Institute of Oceanology, Chinese
Academy of Sciences, Qingdao 266071, China
b UMR-CNRS 5805 EPOC, Université Bordeaux 1, Avenue des Facultés, 33405 Talence
Cedex, France
c Department of Geology, University of Cincinnati, Cincinnati, OH 45221-0013, U.S.A.
d Key Laboratory of Marine Hydrocarbon Resource and Geology, Qingdao Institute of
Marine Geology, Ministry of Land and Resources, Qingdao 266071, China
Abstract
Giant marine diatoms, blooming or aggregating within deep chlorophyll maxima under
stratified conditions, can generate substantial production and a large export flux of organic
carbon from surface waters. However, their role in regulating glacial-interglacial variation in
atmospheric pCO2 remains unclear. Here, we report the organic carbon isotopic compositions
of Ethmodiscus rex diatoms (δ13
CE. rex) and bulk sediments (δ13
Corg) from a sediment core in
* Corresponding author.
E-mail address: tgli@qdio.ac.cn (T. Li).
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the eastern Philippine Sea dated to ~19.5-31.0 kyr B.P. and consisting of (from youngest to
oldest) (1) laminated E. rex diatom mats (LDM), (2) diatomaceous clay (DC), and (3) pelagic
clay (PC). Our results suggest that δ13
CE. rex provides a better record of palaeoceanographic
processes during LDM and DC deposition than δ13
Corg because of reduced differential vital
effects in near-monospecific E. rex fractions. We used the isotopic composition of the coarse
E. rex fraction (δ13
CE. rex (>154 µm)) to calculate the CO2 partial pressure of eastern Philippine
Sea surface waters (pCO2-sw) during the Last Glacial Maximum (LGM). Our pCO2-sw records
suggest that the eastern Philippine Sea switched from being a strong CO2 source during DC
deposition to a weak CO2 sink during LDM deposition. The role of the eastern Philippine Sea
as a CO2 sink during the LGM was promoted by elevated primary production and, to a lesser
extent, intensified water-column stratification. These observations highlight the potential role
of giant marine diatoms in the sequestration of atmospheric CO2 during the LGM and, hence,
support changes in biogenic silica fluxes as a potential cause of lower glacial atmospheric
CO2. Our findings are consistent with an eolian source of silica, as proposed by the ‘silica
hypothesis’ and the ‘silicon-induced alkalinity pump hypothesis’ but not by the ‘silicic acid
leakage hypothesis.’
Keywords: vital effects; Ethmodiscus rex blooms; surface stratification; CO2 sink; marine
silica cycle; late Pleistocene; eastern Philippine Sea
1. Introduction
The atmospheric partial pressure of carbon dioxide (pCO2-atm) was ~80-100 ppm lower
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during the Last Glacial Maximum (LGM; ~24-18 kyr B.P.) than during the Holocene (Petit et
al., 1999), but the exact causes of pCO2-atm changes remain elusive. Among all the carbon
reservoirs on Earth, only the ocean has a carbon storage capacity sufficiently large and a rate
of carbon exchange with the atmosphere sufficiently rapid to account for the missing CO2
during the LGM (Sigman and Boyle, 2000). However, fundamental aspects regarding the role
of the ocean in sequestration of atmospheric CO2 remain uncertain. First, which ocean was
mainly responsible for the decrease in pCO2-atm during the LGM? Recent studies have
suggested that the Southern Ocean (Sigman et al., 2010), the North Pacific Ocean (Jaccard et
al., 2009), and the equatorial Pacific Ocean (Bradtmiller et al., 2010) each had the potential to
modulate glacial/interglacial variability in pCO2-atm. Second, what was the role of the surface
ocean in transferring CO2 to the deep ocean, which is widely believed to have sequestered
large amounts of respired CO2 during glacial stages (Anderson and Carr, 2010)? Third, were
glacial-interglacial pCO2-atm cycles driven primarily by biogenic (e.g., biological pump and
nutrient utilization), physical (e.g., ocean stratification and upwelling), or chemical (e.g.,
ocean alkalinity changes) processes?
Recent hypotheses regarding the cause of lower pCO2-atm during glacial stages have
focused on the role of increased productivity by siliceous marine plankton, e.g., the ‘silica
hypothesis’ or ‘silicon-induced alkalinity pump hypothesis’ (Harrison, 2000; Nozaki and
Yamamoto, 2001) and the ‘silicic acid leakage hypothesis’ (Brzezinski et al., 2002;
Matsumoto et al., 2002). Each hypothesis predicts that an increase in silicic acid levels in
seawater during glacial intervals shifted the composition of the plankton community in favor
of diatoms relative to coccoliths, resulting in a reduced rain ratio of carbonate carbon to
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organic carbon and, consequently, lower pCO2-atm levels (Archer et al., 2000). These models
differ in invoking different sources of dissolved silica in seawater, i.e., eolian dust for the
silica hypothesis and silicon-induced alkalinity pump hypothesis (Harrison, 2000; Nozaki and
Yamamoto, 2001) and export of silicic acid from Antarctic Intermediate Water (AAIW) and
Subantarctic Mode Water (SAMW) for the silicic acid leakage hypothesis (Brzezinski et al.,
2002; Matsumoto et al., 2002). Testing of these hypotheses has been hampered by a lack of
detailed sedimentary records relating biogenic silica fluxes to sediment δ13
Corg in key oceanic
regions.
Laminated diatom mats (LDMs) commonly contain high opal but low carbonate contents,
reflecting an ecosystem shift in favor of diatoms relative to coccoliths (Gingele and
Schmieder, 2001; De Deckker and Gingele, 2002). LDMs, formed by the accumulation of
giant and “shade flora” diatoms such as Ethmodiscus rex, Rhizosolenia spp. and
Thalassiothrix spp., have been found sporadically in abyssal environments of the global
ocean (Kemp et al., 2006; Romero and Schmieder, 2006; Romero et al., 2011). These giant
diatoms, with valve diameter or length of >50 µm (up to 2-5 mm maximally) (Goldman, 1993;
Smetacek, 2000), are adapted to survival in stable and stratified subsurface waters via a
specialized buoyancy strategy (Villareal, 1993; Villareal et al., 1999). They are capable of
vertical migration between the surface for photosynthesis and the nutricline to obtain
nutrients (Villareal and Carpenter, 1994). The giant diatoms mainly bloom or aggregate
within deep chlorophyll maxima (DCM) and generate substantial biomass and export flux.
Kemp et al. (2000) demonstrated that the export flux of these diatom species during the “fall
dump” could rival or exceed that of “spring bloom” species, highlighting the contribution of
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giant diatoms to aggregate surface production. In spite of their probable importance in the
global carbon cycle, these giant diatoms are understudied due to their large size coupled with
the constraints of conventional oceanographic survey schemes (Kemp et al., 2006). Thus, it is
unclear what role the giant diatoms play in global carbon production, export, and CO2
sequestration.
The Philippine Sea is the largest marginal sea in the West Pacific. Its surface waters are
oligotrophic at present, with silicate concentrations of just 2.5-3.5 μmol L-1
(Fig. 1A),
resulting in low primary and export productivity (Berger and Wefer, 1991). Because of low
biological utilization of CO2 and limited upwelling, its surface waters are in equilibrium with
atmospheric CO2 at present (Fig. 1B; Takahashi et al., 2009). E. rex LDMs were found in a
series of cores recovered in the eastern Philippine Sea (Fig. 1C) during the 2003-2004 cruise
of R/V Science No.1 (Zhai et al., 2009). E. rex LDMs and oozes of late Pleistocene age had
been reported previously from Stn. 03 on the Kyushu-Palau ridge (Shibamoto and Harada,
2010), Deep Sea Drilling Project (DSDP) sites 449 and 450 in the Parece Vela Basin (Martini,
1981), DSDP sites 451-456 and 458-460 (Martini, 1981), and an unnamed deep-sea site
(Wiseman and Hendey, 1953) in the Mariana Ridge-Trough-Trench region (Fig. 1C).
The widespread occurrence of E. rex LDMs of late Pleistocene age in the eastern
Philippine Sea (see Section 2.3) raises a number of intriguing questions. What was the trigger
for these E. rex blooms? What was the role of such blooms in the sequestration of atmosphere
CO2 during the LGM? Did these vast E. rex deposits cause the eastern Philippine Sea to serve
as a major sink for atmospheric CO2 at that time? In order to answer these questions, we
present diatom (E. rex)-bound and bulk organic carbon isotopic records from a core
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containing E. rex LDMs collected in the eastern Philippine Sea. The results of this study
provide new insights concerning the cause of glacial-interglacial variation in atmospheric
pCO2-atm and have significant implications for evaluating hypotheses invoking biogenic silica
fluxes in global climate cycles.
2. Materials and methods
2.1. Sediment core
A 405-cm-long gravity core (WPD-03) was collected from a water depth of 5250 m at
17°19.82′ N and 138°27.28′ E in the center of the Parece Vela Basin of the eastern Philippine
Sea (Fig. 1C). The study site lies below the calcium carbonate compensation depth (CCD)
and, consequently, sediments accumulating there contain little carbonate (<1%) (Xiong et al.,
2012b). Core WPD-03 comprises three discrete lithostratigraphic units. The upper unit (0-286
cm) is composed of olive-grey and grey laminated diatom mats (LDM), which are dominated
by fragmented valves of the mat-forming diatom E. rex in near-monospecific assemblages
(Fig. 2A-B) and contain extremely low abundances of other diatom species and radiolarians
(Zhai et al., 2009). The middle unit (286-334 cm) is characterized by grey diatomaceous clays
(DC) that are dominated by diatoms but lack lamination. The lower unit (334-405 cm)
comprises massive red pelagic clays (PC) that generally lack a microfossil component.
2.2. Analytical procedures
A series of physical separation and chemical oxidation steps were utilized to extract and
isolate diatoms, including the primary species of interest (E. rex), for carbon isotope
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measurements (Xiong et al., 2012c). Bulk wet samples (Fig. 2B) were treated with 10% H2O2
and 1 mol mL-1
HCl to remove excess organic matter and carbonate. The samples were then
wet-sieved to obtain the 63-154 µm and >154 µm size fractions using 63-µm and 154-µm
steel meshes. Pure diatom remains were obtained from the two size fractions via
centrifugation at 1500 rpm and heavy liquid flotation using 2.3 g mL-1
sodium polytungstate
(SPT) (Fig. 2C). Labile organic matter coating the diatom frustules (Fig. 2D) was removed
via a chemical oxidation step, in which samples were immersed in 30% H2O2 at 65°C for 2
hours. Sample purity was verified by visual inspection of treated samples using both standard
light and scanning electron (SEM) microscopy. However, our experiments indicate that
samples with opal concentrations of <10% cannot be totally purified by the physical
separation technique described here (Xiong et al., 2012c).
The 63-154 µm and >154 µm diatom fractions consist of nearly pure E. rex except for the
presence of trace radiolarians in the 63-154 µm fraction. The mass percentages of 63-154 µm
and >154 µm E. rex fractions were calculated by weighing the mass of 63-154 µm and >154
µm E. rex fractions after physical separation steps. The abundance of radiolarians in the
63-154 µm E. rex fraction was estimated following Xiong et al. (2012c), in which surface
areas of radiolarians and E. rex were calculated by the image processing software (Leica
Qwin pro) in the light microscope. The carbon isotopic measurements of organic matter in the
cleaned 63-154 µm and >154 µm E. rex fractions (Fig. 2E) were performed on a Euro Vector
EA3000 elemental analyzer in line with a GV IsoPrime isotope ratio mass spectrometer, with
an analytical precision (SD) less than 0.15‰. The carbon isotopes of 63-154 µm and >154
µm E. rex fractions were reported in the standard notation (δ13
CE. rex (63-154 µm) and δ13
CE. rex
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(>154 µm)), relative to the Pee Dee Belemnite standard, expressed in units of ‰. To compare
with δ13
CE. rex, the carbon isotopes (δ13
Corg) of bulk organic matter from samples of core
WPD-03 were also analyzed using the same instruments described above, with an analytical
precision (SD) better than 0.10‰.
2.3. Depositional age of LDM
Reliable dating of opal-rich marine sediments using the AMS14
C method is challenging
regardless of whether measurements are undertaken on carbonate (Broecker et al., 2000; De
Deckker and Gingele, 2002) or bulk organic matter (Zheng et al., 2002). Recent advances in
dating diatom-bound organic compounds represent a major step forward in the development
of age models for carbonate-free marine sediment cores (Ingalls et al., 2004; Hatté et al.,
2008). Radiometric dating of bulk organic matter in the WPD-03 study core yielded a mostly
orderly sequence of uncalibrated AMS14
C ages ranging between 27.6 kyr and 17.8 kyr (Fig.
3A; n.b., all reported ages are “before present” or B.P.). Radiocarbon ages younger than 20
kyr in core WPD-03 were converted to calendar ages using CALIB 5.0.2 program and Marine
04 calibration (Stuiver et al., 2005) and an assumed reservoir age of 400 years. AMS14
C ages
older than 20 kyr were calibrated to calendar years following Laj et al. (1996). The
construction of an age model for the study core was based on the linear regression of the
calibrated age data, excluding three anomalous samples at 228-286 cm (Fig. 3A).
The age model for the WPD-03 study core suggests deposition of the LDM and DC at
~29.4-19.5 kyr and ~31.0-29.4 kyr, respectively, yielding an average linear sedimentation rate
(LSR) of 27.4 cm kyr-1
for this interval. The age at the top of core WPD-03 (19.5 kyr) is
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consistent with non-deposition of E. rex during the last deglaciation and the Holocene (De
Deckker and Gingele, 2002). Rates of sedimentation of sub-recent pelagic clays in the eastern
Philippine Sea are quite low (~1.38 mm kyr-1
; Xu et al., 2008), so the thin fluff layer (~2.7
cm, based on a sedimentation rate of 1.38 mm kyr-1
) formed during the last deglaciation and
the Holocene is likely to have been easily removed by bottom currents and/or during core
collection. The age model developed here for core WPD-03 is robust enough to support
oceanic and climatic interpretations at approximately a millennial scale.
The WPD-03 study core contains three samples in the lower LDM (228-286 cm) yielding
anomalously young radiometric ages. This radiocarbon age reversal may have resulted from
one or more of the following causes: (1) input of exogenous organic matter owing to reactive
opal adsorption and/or lateral transport by oceanic currents, (2) adsorption of atmospheric
CO2 to opal surface and/or the fixation of atmospheric CO2 by chemosynthetic bacteria
during core storage (Zheng et al., 2002), (3) upsection vertical migration of old volatile
organic compounds derived from organic matter degradation during early diagenesis, or (4)
underestimation of reported dating errors (~65-140 yr) relative to actual errors. However,
physical overturn or inversion of the sedimentary sequence itself can be ruled out.
The LDM ages in the WPD-03 core are consistent with those obtained from fourteen
other sediment cores in the study area, in all of which the LDM deposits date to between 28.6
kyr and 16.05 kyr B.P. (Fig. 3B; Zhai et al., 2009; n.b., uncalibrated radiocarbon ages). These
ages indicate that LDMs formed primarily during marine isotope stage 2, approximately
corresponding to the LGM (Lisiecki and Raymo, 2005). This inference is corroborated by a
late Pleistocene age assignment for E. rex oozes from DSDP sites 449 and 450 on the basis of
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their radiolarian faunas (Martini, 1981). The formation age of LDMs in the study area is also
consistent with that of E. rex oozes in the equatorial Indian Ocean reported by Broecker et al.
(2000) and De Deckker and Gingele (2002).
3. Results
The concentration of opal in core WPD-03 ranges from 1.5% to 75.9%, with minimum
values in the PC and peak values in the middle to upper part of the LDM (Fig. 4A). The
upsection increase in opal is counterbalanced by a corresponding decline in clays, the second
most abundant constituent of the study core. The concentrations of total organic carbon (TOC)
(0.09%-0.35%) and total nitrogen (TN) (0.02%-0.08%) are low throughout the study core
(Fig. 4B) owing to severe degradation of organic matter and/or its strong dilution by biogenic
silica (Xiong et al., 2012b). The TOC profile is similar to that for opal, with minimum values
in the PC and maximum values in the LDM, although differences in TOC concentrations
between lithologic units are smaller than for opal. The TN profile shows strong positive
covariation with the TOC profile over short stratigraphic intervals, but differs from both the
TOC and opal profiles in exhibiting maximum values in the DC and minimum values in the
upper LDM, resulting in a progressive increase in TOC/TN molar ratios upsection (Fig. 4C).
Rising TOC/TN in the upper part of the WPD-03 core may have been due to organic matter
diagenesis under increasingly reducing conditions during LDM deposition (Xiong et al.,
2012b).
Bulk δ13
Corg values gradually increase from -21.48‰ at the base of the PC to -16.86‰ in
the lower LDM and then stabilize at around -18‰ in the middle to upper part of the LDM,
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exhibiting a two-step change pattern (Fig. 5A). The δ13
CE. rex profiles show distinct
differences from the bulk δ13
Corg record in the WPD-03 core. Whereas δ13
Corg exhibits a
modest ~2‰ increase at the transition from DC to LDM (Fig. 5A), δ13
CE. rex increases
abruptly at this contact by ~6‰ for the 63-154 m fraction and by ~7‰ for the >154 m
fraction (Fig. 5B-C). The subsequent depletion observed between 28.8 kyr and 27.0 kyr in the
δ13
CE. rex records is not evident in δ13
Corg. However, from 26.3 kyr to the core top, all three
profiles display similar δ13
C values and patterns of variation. These similarities are probably
due to the diatom E. rex being the dominant source of organic matter in this interval. We
suggest that the δ13
CE. rex profiles are superior to bulk δ13
Corg as a recorder of environmental
changes because of relatively constant vital effects in near-monospecific E. rex assemblages
(see Section 4.2). Differences in δ13
CE rex are evident between the two diatom size fractions
(Fig. 5D). Relative to the 63-154 µm fraction, the >154 µm fraction is systematically
13C-depleted (by 3-4‰) at 31-29 kyr but slightly
13C-enriched (by ~1‰) at 29-22 kyr.
The principal source of OM to the study site can be assessed through a combination of
δ13
Corg and TOC/TN data (Lamb et al., 2006). δ13
Corg varies from -16.86‰ to -22.31%
(average -19.56‰) and TOC/TN varies from 3.1 to 13.5 (average 6.7; Fig. 6). These ranges
are consistent with derivation of organic matter in the WPD-03 core entirely from marine
algae, with little or no contribution from terrigenous sources. A lack of measurable amounts
of terrigenous organic matter reflects the open pelagic deep-sea environment of the study
area.
4. Discussion
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4.1. Causes of size dependence of δ13
CE. rex
The size dependence of δ13
CE rex (Fig. 5D) is unlikely to be due to sample impurity or
stochastic factors. Although the 63-154 µm fraction contains a small amount of non-diatom
material (i.e., <3% radiolarians; Fig. 4E), this level of impurity cannot account for δ13
C
differences as large as 5.4‰ between the two size fractions. In addition, for 43 out of 48
samples, the two E. rex size fractions display a δ13
C offset larger than the analytical precision
(SD, 0.15‰). Consequently, assuming that the size of E. rex fragments exhibits a positive
correlation with the size of intact E. rex specimens, we infer that the δ13
CE. rex offsets shown
in Figure 5D resulted from size-related vital effects. This inference is consistent with culture
experiments showing the influence of cell size on carbon isotopic fractionation in diatoms
(Popp et al., 1998; Burkhardt et al., 1999).
There are several potential causes of the observed relationship of diatom size to δ13
CE. rex.
First, diatom size might vary with water depth, causing δ13
CE. rex to vary as a function of
vertical gradients in salinity or temperature. Although capable of migrating vertically, E. rex
is a photosynthetic organism that spends much of its time within the photic zone, which
extends to a depth of ~100 m in the clear open ocean. The upper 100 m of the eastern
Philippine Sea water column exhibit only modest salinity and temperature gradients,
amounting to ~0.3 psu and ~2.6 °C, respectively (Antonov et al., 2006; Locarnini et al., 2006).
While the salinity and temperature coefficients of carbon isotopic fractionation in diatoms are
not known in detail, it is unlikely that these relatively small oceanographic gradients could
account for offsets of δ13
C as large as 5.4‰ (Fig. 5D) between E. rex size fractions. Second,
previous work has suggested that diatom C-isotopic compositions can be influenced by
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changes in growth rates, with 13
C-enriched values at high growth rates (Laws et al., 1995;
Burkhardt et al., 1999). High growth rates generate small diatom cells due to high rates of
cell division (Korb et al., 1996), implying that the 63-154 µm fraction should be generally
13C-enriched relative to the >154 µm fraction. However, small diatom cells have a large
surface-to-volume ratio, which results in relatively 13
C-depleted values owing to reduced
carbon demand (Pancost et al., 1997). Therefore, growth rate and surface-to-volume ratio
impose counteracting controls on C-isotope fractionation for the 63-154 µm diatom fraction,
limiting the net effect of these processes.
Another factor is nutrient availability, which can have a significant impact on diatom
physiology (Ragueneau et al., 2000; Sarthou et al., 2005). In stratified ocean systems, the
biological pump results in a net export of nutrients from the surface layer, resulting in
oligotrophic conditions in the photic zone. Nutrients become sequestered below the nutricline,
which is approximately coextant with the thermocline at depths of 100-200 m in the tropical
West Pacific. Low nutrient availability limits the growth of small diatoms incapable of
vertical migration within the water column, but giant diatoms (e.g., E. rex) are capable of
buoyancy regulation and can descend to the nutricline to make use of the higher nutrient
levels there (Villareal et al., 1999). Therefore, nutrient availability should not be a key factor
in explaining size-dependent δ13
CE. rex offsets (Fig. 5D).
One further possibility is variations in seasonality of diatom growth. Large diatoms
(Coscinodiscus spp.) have been shown to have lower δ15
N compositions relative to smaller
diatoms (Studer et al., 2012), a relationship attributed to blooms and more rapid growth of
large diatoms at an early stage in the growing season, when the seawater nitrate pool is larger
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and not yet strongly 15
N enriched. Blooms of smaller diatoms later in the growing season,
following diminution and 15
N enrichment of the seawater nitrate pool, result in higher diatom
δ15
N values. This process might also account for relative 13
C enrichment of smaller diatoms,
as seen in the study units (Fig. 5D), although it has not been verified from modern marine
systems to date.
Size dependence of the δ13
C composition of an individual diatom species is documented
here for the first time, although size-related effects have been reported previously for the
oxygen (Swann et al., 2007, 2008), nitrogen (Studer et al., 2012), and interspecific carbon
isotopic compositions of diatoms (Jacot Des Combes et al., 2008). In the future,
paleoceanographic studies making use of diatom isotopic proxies may need to undertake
separation of monospecific assemblages in order to account for species-specific vital effects
and different life history strategies (De La Rocha, 2006; Crosta and Koç, 2007; Swann et al.,
2009). Our results suggest that even a single species of diatom may yield different δ13
C
signals when different size fractions are investigated. Consequently, considerable caution is
needed when interpreting isotopic data from even a single diatom species in relation to
palaeoceanographic changes. A clear need exists for further model assessment (Karsh et al.,
2003; Trull et al., 2008), sediment trap, core top, and culture studies of diatom-bound organic
carbon isotopes in modern marine systems.
4.2. Changes in pCO2-sw in eastern Philippine Sea during the LGM
Previous studies have suggested that the two main factors controlling bulk δ13
Corg in
marine environments are surface-water CO2 concentrations (CO2(aq)) and primary productivity,
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and, therefore, that δ13
Corg can be used to tentatively reconstruct surface-water CO2 partial
pressure (pCO2-sw) (see review in Royer et al., 2001). Variability in cell growth rate, size, and
shape (Popp et al., 1998; Burkhardt et al., 1999; Rosenthal et al., 2000), community structure
(Popp et al., 1999), biochemical metabolic pathway (Cassar et al., 2004) and carbon source
(Macko et al., 1987) can weaken the relationship between δ13
Corg and pCO2-sw (see review in
Crosta and Koç, 2007). We assume that δ13
CE. rex variations due to these latter factors are of
limited importance in the present study because all δ13
CE. rex analyses were carried out on
restricted diatom size fractions and a single diatom species (E. rex), and, therefore, that δ13
CE.
rex provides a more direct link than bulk δ13
Corg to pCO2-sw. Other potential influences include
local hydrodynamic conditions, terrigenous organic inputs, diagenesis and temperature effects.
In view of the fact that E. rex blooms are associated with stratified watermasses (Villareal,
1993; Villareal et al., 1999; Kemp et al., 2000), upwelling is unlikely to have been an
influence on the CO2 equilibrium between surface waters of the eastern Philippine Sea and
the atmosphere during LDM deposition (see Section 4.3). Although bulk organic matter
preserved in core WPD-03 suffered remineralization during export to the seafloor (Xiong et
al., 2012b), diatom-bound organic matter is comparatively well-protected from
remineralization and diagenesis by its siliceous matrix. Moreover, the lack of significant
correlations between sea-surface temperature in the eastern Philippine Sea and the C-isotopic
composition of the two size fractions of E. rex (r = 0.27 and r = 0.13, respectively; see below)
suggest that δ13
CE. rex values in the study core were not affected by thermal influences on
photosynthetic fractionation, as proposed by Fontugne and Duplessy (1981).
We converted δ13
CE. rex values to pCO2-sw estimates using two different approaches: the
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Rau et al. (1989) linear model and the Popp et al. (1989) exponential model. Although both of
these models were based on the carbon isotopic composition of total photosynthetic biomass
(see below), we used δ13
CE. rex instead because of the dominant role of E. rex as a primary
producer in the study units. For calculation of pCO2-sw in eastern Philippine Sea surface
waters during DC and LDM deposition, we chose to use δ13
CE. rex (>154 µm) rather than δ13
CE. rex
(63-154 µm). We used the coarser size fraction because (1) it represents the bulk of diatom mass
through most of the study section (Fig. 4D), and (2) the finer (63-154 µm) size fraction
contains a trace presence of radiolarians (Fig. 4E), which may bias its δ13
C composition.
Plankton δ13
Corg is linearly related to CO2(aq) according to Rau et al. (1989):
CO2(aq) = (δ13
CE. rex (>154 µm) + 12.6) / -0.8 (1)
This relationship has been widely used to estimate CO2(aq) in local watermasses (i.e., pCO2-sw)
(Pedersen et al., 1991; Kienast et al., 2001). On the other hand, the Popp et al. (1989) model
is based on an exponential relationship determined by McCabe (1985):
CO2(aq) = exp10 [(εP – 3.4) / -17] (2)
where εP is the overall isotope effect associated with photosynthetic carbon fixation. Values
for εP are determined from:
εP = [(δ13
CE. rex (>154 µm) + 1000) / ( δ13
Cd + 1000) – 1] ×1000 (3)
where δ13
Cd represents the carbon isotopic composition of CO2(aq). δ13
Cd is determined from
the temperature-dependent carbon isotopic fractionation between CO2(aq) and precipitated
calcite (Jasper et al., 1994):
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δ13
Cd = δ13
Ccarb + 14.07 – 7050 / T (4)
where δ13
Ccarb and T are the isotopic composition of calcite carbon and the temperature in the
Kelvin scale, respectively. δ13
Ccarb is equal to the carbon isotopic composition of planktonic
foraminifera (δ13
Cforam) when assuming a negligible vital effect.
In view of our inability to generate paleo-temperature, paleo-salinity, and δ13
Cforam
records for the WPD-03 study core due to the absence of carbonate microfossils such as
foraminifera and coccoliths, we substituted paleoceanographic records from core MD98-2181
dating to the same 31.0-to-19.5-kyr interval (Stott et al., 2002). Use of temperature, salinity,
and δ13
CG. rubber records from core MD98-2181 is reasonable because (1) this core is located
in the western part of the eastern Philippine Sea and, thus, is relatively proximal to the
WPD-03 study site, (2) the two sites are within the same surface current system (i.e., North
Equatorial Current and/or its branch Mindanao Current), and (3) there are minimal salinity
(<0.5 psu) and temperature (<1°C) gradients between the sites today (Fig. 1C; Antonov et al.,
2006; Locarnini et al., 2006). On this basis, we infer similar δ13
Cd values for MD98-2181 and
WPD-03 and, therefore, rewrite formula (4) as:
δ13
Cd = δ13
CG. ruber + 14.07 – 7050 / SST (5)
where δ13
CG. ruber and SST (in the Kelvin scale) are the carbon isotopic composition of
Globigerinoides ruber and the sea-surface temperature from core MD98-2181, respectively.
In cases where δ13
CG. ruber and SST estimates are not available at the same age as the δ13
CE. rex
(>154 µm) data, the former were linearly interpolated to the age of the δ13
CE. rex (>154 µm) data.
For both models, calculated CO2(aq) values were then converted to pCO2-sw based on
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Henry’s Law:
pCO2-sw = CO2(aq) / α (6)
where α is a temperature- and salinity-dependent solubility constant. The solubility constant α
was calculated dependent on the temperature and salinity constructed from core MD98-2181
during the 31.0-to-19.5-kyr interval, according to data in Weiss (1974). The sea-surface
salinity (SSS) was calculated using the empirical relation of LeGrande and Schmidt (2006):
SSS = (δ18
Osw + 8.88) / 0.27 (7)
The sea-surface δ18
O composition (δ18
Osw in V-SMOW) was calculated per the equation of
Ravelo and Hillaire-Marcel (2007):
SST = 16.9 - 4.38 × (δ18
OG. ruber - δ18
Osw + 0.27) + 0.1 × (δ18
OG. ruber - δ18
Osw + 0.27)2
(8)
Finally, the solubility constant was calculated per the equation of Weiss (1974):
ln α = -58.0931 + 90.5069 × (100/SST) + 22.2940 × ln (SST/100) + SSS × [0.0278 –
0.0259 × (SST/100) + 0.0051 × (SST/100)2] (9)
In cases where SSS and SST values were not available at the same age as the δ13
CE. rex (>154 µm)
data, the former were linearly interpolated to the age of the δ13
CE. rex (>154 µm) data. One point
to note is that SSTs reported in the Celsius scale should be converted to those in the Kelvin
scale when using equation (9).
The results of the two methods for estimating pCO2-sw in surface waters of the eastern
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Philippine Sea during the 31.0-to-19.5-kyr study interval are shown in Figure 7A-B. Both
methods yield similar downcore changes in pCO2-sw, with extremely high values during DC
deposition. Despite modest uncertainty ranges for pCO2-sw estimates (average ~15-20 ppmv;
Fig. 7A-B), the two methods yield nearly identical absolute pCO2-sw values (Fig. 7C). These
considerations show that our pCO2-sw estimates are robust, supporting the conclusion of Rau
et al. (1991) that the different methods of Rau et al. (1989) and Popp et al. (1989) yield
comparable pCO2-sw estimates within the range of modern ocean CO2(aq).
4.3. Climatic-oceanographic controls on pCO2-sw during the LGM
The difference between pCO2-sw and pCO2-atm ( pCO2) is an indicator of whether ocean
surface waters in a given region are a net source or sink of carbon to the atmosphere. For the
eastern Philippine Sea during DC and LDM deposition, pCO2 can be used to evaluate the
potential role of giant diatom blooms in transferring CO2 from the atmosphere to ocean
sediment. Comparison with the pCO2-atm record from the Vostok ice core (Luthi et al., 2008)
reveals four distinct stages in the pCO2-sw of the eastern Philippine Sea (Fig. 7C). During the
first stage (DC, at 31.0-29.4 kyr), pCO2-sw was extremely high relative to pCO2-atm (average
pCO2 = 239 ppmv), indicating that the eastern Philippine Sea was a strong source of CO2 to
the atmosphere. At the beginning of the second stage (LDM1, at 29.4-27.4 kyr), pCO2-sw
dropped rapidly and then stabilized (average pCO2 = 98 ppmv), indicating that the eastern
Philippine Sea was a moderate source of CO2 to the atmosphere. At the beginning of the third
stage (LDM2, at 27.4-23.8 kyr), another rapid drop in pCO2-sw resulted in a near-equilibrium
with pCO2-atm (average pCO2 = 0 ppmv), indicating that the eastern Philippine Sea was
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neither a source nor a sink of CO2 to the atmosphere. During the fourth stage (LDM3, at
23.8-19.5 kyr), pCO2-sw was mostly somewhat lower than pCO2-atm (average pCO2 = -18
ppmv), indicating that the eastern Philippine Sea was a weak sink for atmospheric CO2.
The switch of the surface waters of the eastern Philippine Sea from carbon source to
carbon sink can be linked to changes in wind shear and upwelling during the LGM.
Clay-mineral ratios have been used as a proxy for wind shear, with a higher proportion of
illite representing a relative increase in wind shear and Asian eolian dust input to the
Philippine Sea (Xiong et al., 2010; Wan et al., 2012; Xu et al., 2012). Illite/smectite ratios
have been shown to covary with the intensity of the East Asian winter monsoon (EAWM)
(Wan et al., 2012; Xu et al., 2012). In the WPD-03 core, illite/smectite ratios rise substantially
within the DC interval, with an average value ~2X greater than those for the PC or LDM
intervals (Fig. 7D), suggesting an increase in EAWM intensity at 31.0-29.4 kyr. The peak in
illite/smectite ratios matches well that in pCO2-sw (Fig. 7A-C), consistent with wind-driven
intensification of upwelling within the eastern Philippine Sea during the DC interval (Xiong
et al., 2010). The SST record (Fig. 7E) shows a minimum coinciding with the peaks in
illite/smectite ratios and pCO2-sw, consistent with surface cooling driven by upwelling during
the DC interval. We infer that intensified upwelling brought high concentrations of deep CO2
to the surface, causing the eastern Philippine Sea to act as a CO2 source to the atmosphere
during the DC interval (Fig. 8A). Upwelling also may have stimulated modest levels of small
“spring bloom” diatom productivity during this interval.
An abrupt decrease in illite/smectite ratios at the DC/LDM transition (Fig. 7D) indicates
that wind-driven upwelling terminated abruptly at ~29.4 kyr. The scenario is also supported
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by the SST estimates, which increase rapidly at the DC/LDM transition (Fig. 7E). The
development of LDMs in tropical ocean settings is generally associated with the onset of
water-column stratification (Gingele and Schmieder, 2001; De Deckker and Gingele, 2002).
Whereas small “spring bloom” diatom species are dominant in turbulent fertile surface waters,
E. rex prefers stably stratified oligotrophic surface waters owing to its ability to descend to
the deep nutrient pool for nitrate and silicic acid. The onset of LDM formation at ~29.4 kyr is
thus consistent with an inferred reduction in upwelling intensity and development of stable
water-column stratification in the eastern Philippine Sea (Fig. 8B). These conditions resulted
in a rapid reduction of pCO2-sw from ~550 ppmv to ~200-300 ppmv (Fig. 7C). As a
consequence, the rate of deep-water CO2 outgassing diminished, but the eastern Philippine
Sea remained a moderate source of CO2 to the atmosphere during the LDM1 interval (Fig.
8B).
During the LDM2 interval, upwelling was completely terminated in the eastern Philippine
Sea, and the degree of water-column stratification intensified. Under these stable
surface-water conditions, CO2 exchange between the surface ocean and the atmosphere
resulted in an equilibrium between pCO2-sw and pCO2-atm (Fig. 8C). During the LDM3 interval,
surface stratification persisted, but the level of primary productivity associated with E. rex
blooms increased, as reflected in profiles for biogenic components such as %opal (Fig. 4A),
TOC/Ti, and Ba/Ti (Xiong et al., 2012b). These observations suggest that the blooms of E.
rex became sufficiently vigorous that they induced a net uptake of CO2 by surface waters of
the eastern Philippine Sea, as reflected in negative pCO2 values (Fig. 7C). Thus, the surface
waters of the eastern Philippine Sea became a weak sink for atmospheric CO2 during the
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LDM3 interval (Fig. 8D). Although large fluctuations in SSTs occurred during LDM
deposition, SSTs were generally higher than for the DC interval (Fig. 7E). These observations
are consistent with surface warming resulting from persistent stratification during LDM
deposition.
4.4. Role of marine silica cycle in pCO2-atm drawdown during the LGM
The transition of eastern Philippine Sea surface waters from carbon source to carbon sink
during LDM deposition suggests a potential role for giant marine diatom blooms in
sequestration of atmospheric CO2 during the LGM. The mass flux (MF) of organic carbon to
the sediment during LDM deposition can be calculated from the burial flux (BF) of organic
carbon and the geographic extent of LDM deposits, which covered an area approximately
from 15 to 21 °N latitude and from 136 to 140 °E longitude:
BF = 27.4 cm kyr-1
× 0.23% × 2.37 g cm-3
× 10 = 1.49 g m-2
yr-1
(10)
where 27.4 cm kyr-1
, 0.23% and 2.37 g cm-3
are average LSR (see Section 2.3), TOC
concentration and sediment density (Xiong et al., 2012b) during LDM deposition,
respectively. Xiong et al. (2012b) estimated an LDM organic carbon burial flux of 5.27 g m-2
yr-1
based on a LSR of 92.9 cm kyr-1
, but reanalysis of the AMS14
C dates for eastern
Philippine Sea cores in the present study yields a revised average LSR of 27.4 cm kyr-1
for
the WPD-03 core (Fig. 3A).
MF = 1.49 g m-2
yr-1
× 6.6 × 105 m × 4.4 × 10
5 m = 4.33 × 10
11 g yr
-1 (11)
where 6.6×105 m and 4.4×10
5 m are the approximate length and width of the area of LDM
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deposition, respectively.
It has been estimated that, relative to the preindustrial Holocene, an additional ~2×1017
g
of atmospheric carbon was transferred to the ocean during the LGM (Yu et al., 2010).
However, it is not yet known how much of this mass of atmospheric carbon was buried in the
sediment as opposed to being retained as dissolved inorganic carbon in seawater. If 100% of
this carbon mass was buried (an unlikely possibility), the maximum mass flux (MFmax) of
atmospheric carbon to the seafloor during the LGM was:
MFmax = 2 × 1017
g / ((24-18) × 103 yr) = 3.33 × 10
13 g yr
-1 (12)
The proportion of this mass flux of organic matter represented by LDM deposits in the
eastern Philippine Sea is thus:
MF/MFmax = 4.33 × 1011
g yr-1
/ 3.33 × 1013
g yr-1
× 100% = 1.3% (13)
These results indicate that LDM deposition in the eastern Philippine Sea accounts for a
minimum of 1.3% of the atmospheric carbon removed to the ocean during the LGM, and
possibly more if some CO2 was retained as dissolved inorganic carbon in seawater. Although
1.3% is only a fraction of the total oceanic uptake of atmospheric CO2, it is nonetheless
significant considering that the area of LDM deposition occupied just ~0.08% of the ocean
floor. The full areal extent of E. rex blooms during the LGM is not known at present; if
widespread, they may have played a more important role in atmospheric CO2 sequestration
than presently perceived (De Deckker and Gingele, 2002). Our calculations suggest a
potentially important role for giant diatom productivity in atmospheric CO2 drawdown during
glacial times.
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Also noteworthy is the role of water-column stratification, which prevents outgassing of
deep water CO2, in sequestering atmospheric CO2 in the ocean during glacial times.
Intensified stratification during LDM deposition contributed to a rapid decline in pCO2-sw in
the eastern Philippine Sea at ~29.4-27.4 kyr (Fig. 7C). This inference is also supported by
evidence for a deep respired carbon pool, reflecting increased carbon storage in the deep
eastern Philippine Sea during the LGM (Xiong et al., 2012a).
Although earlier studies focused on the influence of eolian Fe fluxes and N fixation in
Pleistocene climate transitions (e.g., Martin, 1990; Falkowski, 1997), more recent studies
have proposed hypotheses linking the marine silica cycle to glacial-interglacial changes in
atmospheric pCO2. The ‘silica hypothesis’ (Harrison, 2000) suggests that elevated dust flux to
the ocean mixed layer increased the availability of dissolved silicon for diatom uptake during
glacial times, inducing a shift from coccolith to diatom primary production in marine systems.
Decreased coccolith productivity reduced the flux of calcite to marine sediments, lowering
glacial pCO2-atm levels (Fig. 9A). The ‘silicon-induced alkalinity pump hypothesis’ (Nozaki
and Yamamoto, 2001) proposes a similar mechanism except for raising the significance of
eolian iron in diatom blooms. This hypothesis alleges that both eolian Si and Fe enhance
diatom production but inhibit coccolith growth, a change of phytoplankton accounting for the
lowered pCO2-atm levels during the glacial period (Fig. 9B). In contrast to the two hypotheses
above, the ‘silicic acid leakage hypothesis’ (Brzezinski et al., 2002; Matsumoto et al., 2002)
invokes leakage of silicic acid from the Southern Ocean via AAIW and SAMW as the source
of silica stimulating diatom productivity in the tropical ocean, although it infers similar
changes in the composition of marine plankton communities (Fig. 9C). All three hypotheses
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are in agreement that diminished coccolith productivity resulted in a reduced rain ratio of
carbonate carbon to organic carbon (Fig. 9). According to seawater carbonate chemistry
(Zeebe and Wolf-Gladrow, 2001), 1 mol carbonate carbon production removes 1 mol of
dissolved inorganic carbon (DIC) and 2 equivalent moles of alkalinity, whereas 1 mol organic
carbon generation removes 1 mol DIC but adds a fraction of alkalinity to seawater. Because
pCO2-atm correlates positively with DIC but negatively with alkalinity, the reduced rain ratio
of carbonate carbon to organic carbon can lower pCO2-atm levels (Archer et al., 2000).
These silica-based models yield testable predictions that may allow evaluation of their
viability as a driver of Pleistocene climate change. All three models predict a decline in the
sinking flux of calcium carbonate concurrent with an increase in the sinking flux of biogenic
silica. We have calculated changes in fluxes in the eastern Philippine Sea during the LGM,
and our results suggest that fluxes of TOC and opal are comparable to primary productivity
rates of some modern continent-margin upwelling systems (Xiong et al., 2012b). Fluxes of
this magnitude are likely to have influenced the transition of eastern Philippine Sea surface
waters from carbon source to carbon sink during the LGM (see Section 4.3). These results are
evidence of the importance of the marine silica cycle in modulating glacial-interglacial
transitions during the Pleistocene. The source of silicic acid for E. rex blooms in the eastern
Philippine Sea can be assessed based on silicon isotopic compositions. In cores WPD-03 and
WPD-12, the >154 µm E. rex fraction yielded δ30
Si values of -0.9‰ to 0‰ (relative to the
NBS28 standard; Z. Xiong et al., in preparation). This range of values is similar to that for
silicon in Asian dust but different from that of silicic acid in advected AAIW and SAMW,
indicating that the former is the more likely source of silica for E. rex blooms in the eastern
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Philippine Sea (Z. Xiong et al., in preparation).
The flux of Asian dust to the eastern Philippine Sea, as proxied by illite/smectite ratios
(Fig. 7D), varied considerably during the LGM. The Asian dust flux increased sharply during
DC deposition, subsequently promoting blooms of E. rex during LDM deposition (Fig. 4A).
The ~1.6-kyr lag between the increase in eolian dust flux (during DC deposition) and E. rex
blooms (during LDM deposition) can be explained with regard to insights from giant diatom
ecology and dissolved silicon biogeochemistry (Xiong et al., 2010). First, although
strengthening of the EAWM during DC deposition supplied plenty of eolian silicon for
diatom growth, wind-driven upwelling inhibited blooms of giant diatoms such as E. rex
owing to their requirement for watermass stratification. Upwelling may have stimulated
modest levels of small “spring bloom” diatom productivity, which is supported by an opal
peak (Fig. 4A) and illite/smectite ratio maximum (Fig. 7D) during DC deposition. Second,
only ~5% of eolian silicon is converted to silicic acid in seawater (Duce et al., 1991), which
would have been easily exhausted by small “spring bloom” diatoms during DC deposition.
New silicic acid for E. rex blooms must have been generated from biogeochemical recycling
of eolian silicon, which takes place over some period of time in the water column and,
consequently, can account for the lag of E. rex blooms relative to increased eolian inputs. The
residence time of silicon in seawater is approximately 15 kyr (Tréguer et al., 1995),
suggesting a lag time of 1.6 kyr is reasonable for generation of silicic acid through
biogeochemical recycling of eolian silicon. Thus, our observations are consistent with the
hypotheses that changes in wind shear controlled the availability of the silicic acid in eastern
Philippine Sea surface waters during the LGM, and that giant diatom E. rex blooms
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sequestered atmospheric carbon and lowered glacial pCO2-atm, as proposed by the silica
hypothesis (Harrison, 2000) and silicon-induced alkalinity pump hypothesis (Nozaki and
Yamamoto, 2001).
5. Conclusions
Through organic carbon isotopic analysis of bulk sediments and cleaned individual E. rex
frustules from a sediment core in the eastern Philippine Sea containing LDMs, we draw the
following conclusions regarding the significance of giant marine diatom blooms during the
LGM:
(1) δ13
CE. rex is a more robust palaeoceanographic proxy than δ13
Corg due to reduced
differential vital effects in near-monospecific E. rex fractions. However, δ13
CE. rex exhibits
significant size-dependent variation that may be due to a vital effect. Further study will be
needed to understand the source of this variation and the limits it imposes on use of diatom
stable isotopes in palaeoceanographic analyses.
(2) Conversion of δ13
C E. rex (>154 µm) to pCO2-sw estimates demonstrated that the eastern
Philippine Sea switched from a source to a sink of atmospheric CO2 during the LGM.
Wind-driven upwelling resulted in the extremely high pCO2-sw during DC deposition.
Subsequently, a combination of upwelling collapse and initiation of stratification lowered
pCO2-sw during LDM1 deposition. Development of stable surface stratification during LDM2
resulted in an equilibrium of pCO2-sw with pCO2-atm. Finally, relatively high primary
productivity associated with blooms of the giant marine diatom E. rex caused the eastern
Philippine Sea to become a weak CO2 sink during LDM3 deposition. These observations
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demonstrate a potentially significant role for giant marine diatom blooms in the sequestration
of atmospheric CO2 during the LGM and reflect the general importance of diatom
productivity in the global carbon cycle.
(3) Our geochemical results served to test the validity of published hypotheses concerning
the role of biogenic silica fluxes as a potential cause of lower glacial atmospheric CO2. The
silicon isotopic composition of E. rex is consistent with an enhanced flux of Asian dust prior
to and during giant diatom blooms, which is consistent with the ‘silica hypothesis’ or
‘silicon-induced alkalinity pump hypothesis’, but not with the ‘silicic acid leakage
hypothesis’.
Acknowledgments
Special thanks are owed to Patrick De Deckker for comments on an early draft, Taro
Takahashi and Xufeng Zheng for assistance with the generation of Fig. 1B, and Rebecca
Robinson and George Swann for discussions on the diatom-cleaning method. This study was
supported by the National Natural Science Foundation of China (grant nos. 41230959,
41106042, 40776031, and 41006032) and National Fundamental Research and Development
Planning Project (grant no. 2007CB815903).
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Figure captions
Fig. 1. Maps of eastern Philippine Sea region showing (A) silicate concentrations (μmol L-1
),
(B) mean annual sea-air pCO2 differences (μatm) for the reference year 2000 (non-El Niño
conditions) (data from Takahashi et al., 2009), and (C) SSTs (℃). Shown in (C) are the
location of sediment core WPD-03 (red circle), the general area of occurrence of laminated
Ethmodiscus rex diatom mats (LDM) (open black rectangle), and the positions of cores and
sites discussed in this study (labeled symbols): core MD98-2181 (Stott et al., 2002), Stn. 03
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(Shibamoto and Harada, 2010), DSDP sites 449-450, 451-456 and 458-460 (Martini, 1981),
and an unnamed site (blue diamond) of E. rex oozes near the Marianas Trench (Wiseman and
Hendey, 1953). Arrows show regional circulation (modified from Fine et al., 1994): NEC is
North Equatorial Current, NECC is North Equatorial Counter Current, KC is Kuroshio
Current, and MC is Mindanao Current. The base maps of Fig. 1A and 1C were drawn using
the Ocean Data View (ODV) software package.
Fig. 2. Scanning electron microscope photos showing samples at each stage of the procedure
for physical separation and chemical oxidation prior to organic carbon isotopic analysis. (A)
An intact E. rex; (B) Bulk sediment sample before physical separation; (C) E. rex diatom
frustules after physical separation; (D) Diatom frustules coated by labile organic matter
before chemical oxidation; (E) Pure diatom frustules after chemical oxidation.
Fig. 3. AMS14
C ages of bulk organic matter in the LDM samples from core WPD-03 (A) and
from other sediment cores in the study area (B) (Zhai et al., 2009). In (A), the solid line
represents a linear regression of the calibrated age data, excluding three anomalous samples
(enclosed by a dashed ellipse) that yielded chronologically out-of-sequence dates. LDM:
laminated E. rex diatom mats; DC: diatomaceous clay; PC: pelagic clay. In (B), the dashed
horizontal line shows the average AMS14
C age of all LDMs in the study area.
Fig. 4. Age/depth profiles (WPD-03) for (A) opal content, (B) TOC and TN, (C) TOC/TN
molar ratio, (D) mass percentage (>154 µm vs. 63-154 µm E. rex fraction) and (E) radiolarian
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abundance in the 63-154 µm E. rex fraction. Data in profiles (A-C) are from Xiong et al.
(2012b). See Fig. 3 for abbreviations.
Fig. 5. Age/depth profiles (WPD-03) for (A) bulk δ13
Corg, (B) δ13
CE. rex (>154 µm), (C) δ13
CE. rex
(63-154 µm), and (D) δ13
CE. rex offset between the two diatom size fractions (>154 µm minus
63-154 µm). δ13
Corg is repeated in profiles (B) and (C) for comparative purposes. See Fig. 3
for abbreviations.
Fig. 6. Organic discriminant diagram, showing TOC/TN molar ratios versus δ13
Corg; modified
from Lamb et al. (2006). For comparative purposes, the weight ratios of TOC/TN in Lamb et
al. (2006) were converted to molar ratios in this figure. Composition of bulk organic matter
from core WPD-03 is shown by symbols. See Fig. 3 for abbreviations.
Fig. 7. Age/depth profiles (WPD-03) for pCO2-sw of the eastern Philippine Sea based on the
Rau et al. (1989) linear model (A) and the Popp et al. (1989) exponential model (B). The
errors in PCO2-sw are calculated from the propagated error of the individual δ13
CE. rex (>154 µm)
(0.15‰, SD), δ13
CG. rubber (0.13‰, SD) (L. Stott, personal communication), δ18
OG. rubber
(0.18‰, SD) and SST (1˚C, SD) (Stott et al., 2002) uncertainties. (C) Direct comparison of
the two models as well as the atmospheric pCO2 record from the Vostok ice core (Luthi et al.,
2008). The pCO2-sw records were divided into four intervals (DC, LDM1, LDM2 and LDM3)
shown by dotted horizontal lines; see text for detailed explanation. (D) Illite/smectite ratios
and (E) SSTs (estimated from core MD98-2181; Stott et al., 2002). See Fig. 3 for
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abbreviations.
Fig. 8. Cartoon showing relationships of pCO2-sw variations to oceanographic conditions in
the eastern Philippine Sea during (A) DC, (B) LDM1, (C) LDM2, and (D) LDM3 intervals.
The size or number of arrows shows relative magnitudes of deep CO2 upwelling flux, organic
carbon (Corg) export flux, or net CO2 sea-air flux. See text for discussion.
Fig. 9. Cartoon illustrating the (A) ‘silica hypothesis’ (Harrison, 2000), (B) ‘silicon-induced
alkalinity pump hypothesis’ (Nozaki and Yamamoto, 2001) and (C) ‘silicic acid leakage
hypothesis’ (Brzezinski et al., 2002; Matsumoto et al., 2002). In A-C, the number of diatoms
and coccoliths shows relative magnitudes of their production. Abbreviations: OC, organic
carbon; CC, carbonate carbon; DIC, dissolved inorganic carbon; ALK, alkalinity; AAIW,
Antarctic Intermediate Water; SAMW, Subantarctic Mode Water.
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Fig. 1
B A
Silicate
NEC NEC NEC NEC
KC
KC
KC
MC
MC
NECC NECC NECC NECC
C
SST
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Research Highlights
The giant diatom Ethmodiscus rex shows size-dependent vital effects on δ13C.
δ13CE. rex is a more reliable palaeoceanographic proxy than bulk δ13Corg.
The eastern Philippine Sea (EPS) switched from being a source to a sink of
atmospheric CO2 during the Last Glacial Maximum (LGM).
The shift of the EPS from CO2 source to sink was due to intensified
surface-water stratification, promoting E. rex blooms.
This study highlights the potential role of giant marine diatoms in sequestration
of atmospheric CO2 during the LGM.