Quaternary Science Reviews 39 (2012) 60e72
Contents lists available at SciVerse ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
The mystery of the missing deglacial carbonate preservation maximum
Figen A. Mekik a, *, Robert F. Anderson b, Paul Loubere c, Roger François d, Mathieu Richaud e
a
Department of Geology, Grand Valley State University, Allendale, MI 49401, USA
Lamont Doherty Earth Observatory, Columbia University, Palisades, NY 10964, USA
c
Department of Geology and Environmental Geosciences, Northern Illinois University, DeKalb, IL 60115, USA
d
Department of Earth and Ocean Sciences, University of British Columbia, Vancouver, Canada
e
Department of Earth & Environmental Sciences, California State University, Fresno, CA 93740, USA
b
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 30 March 2011
Received in revised form
18 January 2012
Accepted 27 January 2012
Available online xxx
A leading hypothesis for lower atmospheric CO2 levels during glacial periods invokes increased ocean
stratification with a corresponding shift of dissolved inorganic carbon and nutrients from intermediate
depths to deep waters. If the rapid deglacial rise in atmospheric CO2 (w17e10 ka) were caused by
a breakdown of this stratification and increased ventilation of deep water masses, then one consequence
would be increased CaCO3 preservation in deep sea sediments. We present down core records of CaCO3
preservation for the last 21,000 years from 31 cores in the tropical and subtropical Pacific, Atlantic and
Indian Oceans. Our preservation records are based on a multi-proxy approach involving a new CaCO3
dissolution proxy (the Globorotalia menardii Fragmentation Index), size normalized foraminifer shell
weights and 230Th-normalized CaCO3 accumulation rates. In some cores our proxy records add to the
growing body of evidence in support of the hypothesized breakdown of glacial stratification. However, in
most cores the expected deglacial increase in CaCO3 preservation is missing. Accepting that the deglacial
hypothesis is well supported by other evidence, here we explore processes and conditions that erased the
expected CaCO3 signal from our records including: (1) variations in the ratio of organic carbon to CaCO3
flux in the eastern equatorial Pacific, (2) very low sedimentation rates and bioturbation in the western
equatorial Pacific and (3) increased northward penetration of Antarctic Bottom Water in the equatorial
Atlantic.
Ó 2012 Elsevier Ltd. All rights reserved.
Keywords:
Deglaciation
Calcite preservation
Organic carbon to calcite rain ratio
Sediment focusing
1. Introduction
Paleoclimatologists have sought for nearly three decades to
identify the process(es) regulating the late-Pleistocene climaterelated changes in atmospheric CO2, the longest record of which
has been extracted from the EPICA Dome C ice core (Siegenthaler
et al., 2005; Luthi et al., 2008). Several hypotheses have been
proposed, and extensive testing of these ideas led to the conclusion
that no single mechanism accounted for the full amplitude of CO2
variability (Archer et al., 2000; Sigman and Boyle, 2000). Although
more than one mechanism may be required (Köhler et al., 2005;
Peacock et al., 2006), there is a convergence of views that lower
glacial CO2 levels require increased (relative to interglacials)
isolation of deep water masses from the atmosphere, for example
by increased stratification or reduced vertical mixing in the ocean
(e.g., Sigman and Boyle, 2000; Ridgwell et al., 2003; Köhler et al.,
* Corresponding author. Tel.: þ1 616 331 3020; fax: þ1 616 331 3740.
E-mail address: mekikf@gvsu.edu (F.A. Mekik).
0277-3791/$ e see front matter Ó 2012 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2012.01.024
2005; Peacock et al., 2006; Toggweiler et al., 2006; Watson et al.,
2006; Sigman et al., 2010).
If increased ocean stratification during glacial periods were
a dominant factor regulating atmospheric CO2 variability, then the
rapid rise in atmospheric CO2 during the last deglacial period
(w17e10 ka; thousand years before present) would have involved
the breakdown of this stratification and increased ventilation of
deep ocean water masses. Although direct evidence to identify the
location of, and the processes associated with, mixing and ventilation exist (Anderson et al., 2009), there is also substantial indirect
evidence to support this scenario. For example, the widespread
distribution of 13C-depleted carbon that invaded the upper ocean
and atmosphere during deglaciation (Smith et al., 1999; Spero and
Lea, 2002; Köhler et al., 2005), and the precipitous drop during
deglaciation in 14C activity of dissolved inorganic carbon in North
Pacific intermediate waters (Marchitto et al., 2007), have been
linked to a deglacial increase in deep overturning circulation.
Preservation of CaCO3 in deep-sea sediments offers another
indirect proxy for past changes in ventilation of deep water masses.
If other factors are held constant, the transfer of CO2 from deep
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
water to the atmosphere will raise the [CO2
3 ] of deep water
throughout the entire ocean, thereby creating a temporary
maximum in CaCO3 preservation in deep-sea sediments.
Evidence for a deglacial peak in CaCO3 preservation was
described more than three decades ago (Berger, 1977), and
corroborating results have appeared subsequently (e.g., Broecker
et al., 2001; Broecker and Clark, 2003; Jaccard et al., 2009, 2010),
seemingly in support of the “deglacial ventilation” hypothesis.
However, evidence is building from a growing number of sites
where the expected deglacial peak in CaCO3 preservation is not
found. This raises questions about the deglacial ventilation
hypothesis, as well as about the methods used to characterize
changes in CaCO3 preservation. Specifically:
1) Are there regional patterns to the presence or absence of the
deglacial preservation event in deep-sea sediments?
2) If there are such patterns, can they be linked to regional
geochemical and circulation responses to deglaciation, and can
their examination expand our understanding of those
responses?
3) Are there unrecognized artifacts in the proxies used to reconstruct past changes in CaCO3 preservation?
1.1. Why is a deglacial CaCO3 preservation peak expected?
Principles underlying the expected deglacial increase in CaCO3
preservation were described by Broecker and Peng (1987) and
Boyle (1988). They can be linked to the deglacial breakdown of
ocean stratification and rise of atmospheric CO2 as follows:
Under steady state conditions, removal of alkalinity from the
ocean by burial of CaCO3 must balance the supply of alkalinity by
continental weathering. Any perturbation of this balance will alter
the [CO2
3 ] of seawater and, therefore, the preservation and burial
of CaCO3, so as to restore the balance.
The glacial stratification scenario (e.g., Boyle, 1988; Sigman and
Boyle, 2000; Toggweiler, 2006) invokes reduced ventilation of the
deep ocean, for example by a reduction in the deep overturning
circulation of the ocean, accompanied by an increase in the efficiency of the biological pump to generate a net transfer of carbon
from the atmosphere and surface ocean into the deep sea. In
addition, there is a positive feedback from carbonate compensation
(Broecker and Peng, 1987). Specifically, adding respiratory CO2 to
the deep ocean lowers the [CO2
3 ] by driving the reaction shown in
Eq. (1) to the right:
H2 O þ CO2 þ CO2
3 42HCO3
(1)
Lowering the [CO2
3 ] reduces the fraction of CaCO3 that is
preserved and buried by driving the reaction in Eq. (2) to the right:
CaCO3 4Ca2þ þ CO2
3
(2)
[CO2
3 ]
Reduced CaCO3 burial causes the alkalinity and
of the
ocean to increase until the balance between alkalinity supply by
continental run-off and removal by CaCO3 burial is restored.
Although the [CO2
3 ] of the deep ocean returns approximately to its
original value under this scenario, vertical concentration gradients
are greater in a more stratified glacial ocean, causing the [CO2
3 ] of
surface waters to exceed those of interglacial periods. The increased
[CO2
3 ] of surface waters removes CO2 from the surface ocean and
atmosphere (Eq. (1)),contributing to the CO2 draw down (Broecker
and Peng, 1987; Boyle, 1988).
If increased ventilation of deep waters were responsible for the
rapid rises in CO2 after 18 ka (see Ahn and Brook, 2008; Anderson
et al., 2009), then this venting of CO2 would have generated
61
a sudden increase in the [CO2
3 ] of the deep ocean by driving the
reaction in Eq. (1) to the left. Concurrently, there would have been
an increase in the preservation and burial of CaCO3 in deep-sea
sediments as the increased [CO2
3 ] drove the reaction in Eq. (2) to
the left. Increased CaCO3 preservation would have lasted for
a period of several thousand years, long enough to restore the
balance between alkalinity supply and removal under conditions of
reduced ocean stratification.
These principles are well established. Evidence for increased
CaCO3 preservation during deglacial periods has been used to
support this scenario (see Section 1.3). However, although this
evidence provides compelling support for the hypothesis, deglacial
maxima in CaCO3 preservation are absent in records from a number
of sites where they would be expected, leading to the research
questions posed above.
1.2. Drivers of deep sea calcite dissolution
Calcium carbonate dissolution in deep-sea sediments is driven
by two independent factors: bottom water [CO2
3 ] and the release
of CO2 into sediment pore waters by respiration.
2
Bottom water [CO2
3 ] is often expressed as DCO3 , which is
defined as the difference between [CO2
]
and
[CO2
3 in situ
3 ] at
2
saturation. Where DCO3 is positive, the sediment is above the
calcite saturation horizon (¼water depth where [CO2
3 ] is at saturation with respect to calcite solubility) and calcite is likely to be
better preserved; where DCO2
3 is negative, the sediment is below
the calcite saturation horizon and calcite dissolution is thermodynamically favorable.
Independently of DCO2
3 , metabolism of organic carbon in
sediments drives additional CaCO3 dissolution by release of respiratory CO2 into pore waters (Emerson and Bender, 1981; Archer and
Maier-Reimer, 1994). If the rain of organic carbon is high enough,
then CaCO3 dissolution can occur even in sediments well above the
calcite saturation horizon, where bottom waters have positive
DCO2
3 values (Archer and Maier-Reimer, 1994). Whereas the rate of
CaCO3 dissolution depends on the supply of organic carbon, the
percent of CaCO3 that dissolves also depends on the rate of CaCO3
supply. Consequently, one often refers to the organic carbon/CaCO3
rain ratio as the variable that influences the percent of CaCO3 that is
ultimately preserved in the sediment record.
1.3. Summary of previously published deglacial CaCO3 preservation
data
The seminal work by Berger (1977), reconstructing aragonite
preservation trends in deep-sea cores over the last 20,000 years,
presents compelling evidence for a dramatic world-wide deepening in the aragonite compensation depth during the last deglacial
period, centered at around 14 ka. Berger (1977) constrained the
aragonite preservation peak to an interval spanning w1000 years.
Deglacial pteropod-rich layers were found in cores off northwest
Africa, Portugal, west India and the western equatorial Pacific
(WEP) lending support to the idea that the deglacial CaCO3 preservation maximum was a global event.
Broecker et al. (2001) and later Broecker and Clark (2003)
published size-normalized planktonic foraminifer whole shell
weight (SNSW) data in WEP cores showing heavier foraminifers
during the late deglaciation, which were interpreted to indicate
a period of enhanced CaCO3 preservation. Although the maximum
shell weights corresponded to an age of w10 ka, the apparent age
discrepancy compared to the pteropod event (w14 ka) may be
related to the effect of bioturbation as the WEP cores had low
sedimentation rates. Deglacial peaks in CaCO3 preservation have
also been observed in North Pacific sediments at 12e15 ka (Jaccard
62
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
et al., 2009), in a core from the Cape Basin at 10e17 ka (ODP 1089;
Hodell et al., 2001), and in several cores from the eastern equatorial Pacific (EEP) at 10e18 ka (Lalicata and Lea, 2011). Marchitto
et al. (2005) interpreted Zn/Ca ratios in benthic foraminifers as
a proxy for reconstructing bottom water [CO2
3 ] in two cores from
the EEP (RC13-114 at 3436 m water depth, and ODP 849 at 3851 m
water depth). Both cores show a prominent deglacial increase in
Zn/Ca ratios, from which the authors inferred an increase in [CO2
3 ]
during the last deglaciation (w10e14 ka). In RC13-114 the increase
in Zn/Ca is corroborated by an increase in Neogloboquadrina
dutertrei shell weights, and in ODP 849B lower fragmentation
index values corroborate the deglacial Zn/Ca spike. Further
corroborating evidence for the results presented by Marchitto et al.
(2005) came from the work of Yu et al. (2010) where a deglacial
increase of [CO2
3 ] by 15 mmol/kg, when compared to preceding
glacial and subsequent interglacial levels, was inferred from B/Ca
ratios in benthic foraminifera in five cores from three major ocean
basins.
The ideal location for investigating the deglacial carbonate
preservation peak requires special conditions: [1] a sediment
accumulation rate high enough that the deglacial signal is well
resolved despite the filtering effect of bioturbation, [2] cores from
depths ranging between the calcite saturation horizon and the
carbonate compensation depth, and [3] a relatively constant
focusing factor over time so that changes in CaCO3 dissolution can
be ascribed to changes in bottom water chemistry rather than to
changes in focusing. It is difficult to find a location that would fulfill
these criteria, particularly in low latitudes where there is abundant
carbonate in the sediments. Therefore, until ideal sites are discovered and cored, it is necessary to work with cores that are available,
keeping in mind the following caveats. The mere absence of the
preservation spike in any core does not require that DCO2
3 was
unchanged during the deglaciation. The fact that the deglacial
carbonate preservation peak is observed in some cores, especially
those from the North Pacific where this preservation event is most
clear (Jaccard et al., 2009, 2010), is evidence that the global ocean
had elevated bottom water [CO2
3 ] during the deglaciation. The
carbonate preservation record is a convolution of whole and
regional oceanic processes. Finding that the deglacial preservation
event is regionally absent provides evidence for changes in circulation and biogeochemical cycling which over-printed the whole
ocean signal. Our goal is to establish the regional patterns of
carbonate preservation and to use these to further define oceanographic responses to deglaciation.
2. Cores and proxies
2.1. Cores and data sources
We sought evidence for the deglacial CaCO3 preservation peak
in 31 deep-sea sediment cores combining newly generated data
with data compiled from the scientific literature. We chose our core
locations to provide wide geographic coverage. Also, we chose
specific cores on which previous work was done with various
proxies in order to compare those proxy results with those we
newly generated herein, such as cores used in Broecker et al. (2001)
and Marchitto et al. (2005). Tables 1 and 2 list all geographic, agemodeling and bottom water DCO2
3 information for each of our 31
cores. Bottom water [CO2
3 ] information for our cores is from
GLODAP bottle data (Key et al., 2004; Sabine et al., 2005) and DCO2
3
was calculated from this data using Ocean Data View software
(Schlitzer, 2008). Fig. 1 shows the locations of our cores overlain on
Table 1
Core information, age models and sedimentation rates. Italicized numbers denote reference citations: 1. Loubere et al. (2004), 2. Loubere and Richaud (2007), 3. Kienast et al.
(2007), 4. Loubere et al. (2003), 5. Pisias et al. (1990), 6. Marchitto et al. (2005), 7. Bradtmiller et al. (2007), 8. Martinson et al. (1987), 9. Charles et al. (1996), 10. Broecker et al.
(2001), 11. Berger and Killingley (1982), 12. Boltovskoy (1992), 13. Francois et al. (1990), 14. Kiefer et al. (2006).
Region
Cores
Latitude
Longitude
Water depth [m]
Age model
Age data
Average sedimentation rate (cm/ka)
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
CEP
CEP
WEP
WEP
WEP
WEP
WEP
WEP
WEP
WEP
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Indian
ODP 846B
ODP 849B
RC13-110
Y69-71
Y71-9-101
VNTR01-8
RC13-114
ME-24
ME-27
P7
T163-19P
V19-30
V21-40
RC13-140
RC11-238
TT013-PC18
TT013-PC72
MW91-9 36BC
MW91-9 38GGC
MW91-9 51GGC
MW91-9 56GGC
ERDC 125
ERDC 131
RC17-177
MD2138
KNR110 82 GGC
KNR110 58GGC
KNR110 55GGC
EN066 38GGC
EN066 21GGC
ENO66 29GGC
GS7309-6PC
WIND 28 KA
3.095
0.183
0.1
0.1
6.383
0.183
1.65
0.022
1.853
2.604
3.6
3.383
5.517
2.867
1.517
1.84
0.11
0
0
0
0
0.003
0.026
1.45
1.25
4.34
4.79
4.95
4.918
4.233
2.46
2.533
10.154
90.818
110.517
95.65
86.48
106.94
110.517
103.63
86.463
82.787
83.986
83.95
83.517
106.767
87.75
85.817
139.71
139.4
158
158
158
158
160.986
162.702
159.45
146.24
43.49
43.04
0.43.89
20.498
20.625
19.762
12.993
51.769
3307
3851
3231
2741
3175
3791
3436
2941
2203
3085
3209
3091
3182
2246
2573
4354
4298
2310
2456
3430
4041
3368
4441
2600
1900
2816
4341
4556
2931
3995
5104
3310
4157
1 and 2
1 and 2
1
1 and 3
4
5
6
3
3
3
3
7
7
7
7
7
7
10
8 and 9
10
10
11 and 12
11 and 12
7
7
13
13
13
13
13
13
New
14
d18O and14C
d18O and14C
d18O
d18O and14C
d18O
d18O
d18O and14C
d18O and14C
d18O and14C
4
3.7
2.7
8.9
3.2
2.7
3.2
15.2
6.1
3.5
3.1
9.2
4.3
5.4
5.2
1.6
2.5
3.1
2
2.7
1.8
2.4
1.2
2
9.6
4
3.5
3.3
1.5
1.8
2.4
5.4
4.1
14
C
d18O and14C
d18O
14
C
C
C
18
d O
18
d O
14
C
d18O
14
C
14
C
18
d O
18
d O
d18O
d18O
d18O
d18O
d18O
d18O
d18O
d18O
d18O
d18O
14
14
and14C
and14C
and14C
and14C
and14C
and14C
and14C
and14C
and14C
and14C
and14C
and14C
and14C
and14C
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
63
Table 2
DCO2
3 and data sources for various proxies from each core. Italicized numbers denote reference citations: 1. Loubere et al. (2004), 2. Kienast et al. (2007), 3. Marchitto et al.
(2005), 4. Bradtmiller et al. (2007), 5. Francois et al. (1990), 6. Loubere et al. (2003), 7. Broecker et al. (2001).
Region
Cores
Foraminifer weight
MFI
230
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
EEP
CEP
CEP
WEP
WEP
WEP
WEP
WEP
WEP
WEP
WEP
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Eq. Atl.
Indian
ODP 846B
ODP 849B
RC13-110
Y69-71
Y71-9-101
VNTR01-8
RC13-114
ME-24
ME-27
P7
T163-19P
V19-30
V21-40
RC13-140
RC11-238
TT013-PC18
TT013-PC72
MW91-9 36BC
MW91-9 38GGC
MW91-9 51GGC
MW91-9 56GGC
ERDC 125
ERDC 131
RC17-177
MD2138
KNR110 82 GGC
KNR110 58GGC
KNR110 55GGC
EN066 38GGC
EN066 21GGC
ENO66 29GGC
GS7309-6PC
WIND 28 KA
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
7
e
7
7
e
e
e
e
e
e
e
e
e
e
e
e
1
1
1
1
New
New
New
New
New
e
e
e
e
e
e
e
e
New
New
New
New
e
e
e
e
e
e
e
e
e
e
New
New
1
1
1
1
New
e
4
2
e
e
e
4
4
4
4
4
4
4
e
4
e
New
New
4
4
5
5
5
5
5
5
e
e
Th-normalized CAR
top of a map of bottom water DCO2
3 (Archer, 1996). All new data
presented herein are available from the lead author and they will be
submitted to the National Climatic Data Center upon publication.
2.2. CaCO3 preservation proxies
2.2.1. Globorotalia menardii fragmentation index
The G. menardii fragmentation index (MFI) was developed by
Mekik et al. (2002, 2010) and is based on laboratory experiments of
Ku and Oba (1978), which showed that dissolution damage in
G. menardii shells is quantifiable. MFI is the ratio of the number of
damaged G. menardii specimens (D) to the number of whole (W)
plus damaged specimens of this species within a sediment aliquot,
DCO2
3 [Zn/Ca]
DCO2
3 core top GLODAP mmol/kg
Organic Carbon flux
e
e
e
e
e
e
3
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
6.62
13.05
6.49
4.09
4.7
11.77
8.78
4.59
1.59
7.47
6.07
4.49
4.63
0.5
1.09
20.83
20.53
5.79
5.69
5.16
15.82
6.38
24.57
2.75
13.54
36.74
0.76
0.96
30.1
8.72
21.25
24.94
14.52
6
6
6
2 and 6
e
e
e
2
2
2
2
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
e
New
e
such that MFI ¼ D/(D þ W). The number of damaged specimens per
sample is calculated with Eq. (3).
D ¼ #with holes þ #>half þ ð# < half=3Þ þ ð#keels=5Þ
(3)
The keel is a thick calcareous rim at the edge of the foraminifer
shell.
The MFI transfer function relates the fragmentation trend of
G. menardii shells in core tops of deep marine sediments from
tropical and subtropical regions of three ocean basins (Pacific,
Atlantic and Indian) to model-derived estimates of percent CaCO3
dissolved with the calibration relationship shown in Eq. (4)
(R2 ¼ 0.84; Mekik et al., 2002, 2010). See Mekik et al. (2010) for
details of the modeling and calibration of the MFI transfer function.
Fig. 1. Core locations shown over map of bottom water DCO2
3 , the difference between the in situ carbonate ion concentration and the carbonate ion concentration that would exist
at saturation with calcite, expressed in mmol/kg (Archer, 1996). Black dots are cores with MFI data, purple dots show cores for which we compiled data from literature. (For
interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
64
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
%CaCO3 Dissolved ¼ 4:0081 þ ðMFI*113:87Þ ðMFI2*37:879Þ
(4)
MFI is unique among available CaCO3 dissolution proxies in
multiple ways: (1) G. menardiis have a quantifiable fragmentation
trend with increasing dissolution; (2) MFI is the only dissolution
proxy calibrated with model-derived estimates of percent CaCO3
dissolved per sample location (Mekik et al., 2002, 2010); (3) MFI is
efficient (w20e30 min per sample); (4) G. menardii fragments are
easy to identify; (5) MFI has demonstrated success in tracing CaCO3
dissolution in places where the surface ocean has a strong productivity gradient (Mekik et al., 2002, 2007a, 2007b); (6) MFI has some
independent corroboration from Mg/Ca and Mg/Sr in multiple
species of planktonic foraminifers (Mekik and François, 2006); and
(7) Mekik and Raterink (2008) showed that MFI-based percent
CaCO3 dissolved estimates are mostly insensitive to surface ocean
environmental parameters in G. menardii’s calcification waters in
the EEP (such as surface ocean temperature or [CO2
3 ]).
The MFI transfer function has a predictive error of 10e15%
calcite dissolved in its core top calibration (Mekik et al., 2010),
which includes errors introduced into modeling from organic
carbon and calcite flux data. This defines the accuracy of the MFI
proxy. Its precision can be ascertained by repeated measurements
of MFI from the same sediment aliquot by multiple researchers.
Reproducibility of MFI data among three researchers was presented
by Mekik et al. (2010) to yield a precision of 0.04 MFI units per
measurement, which corresponds to w2e4% calcite dissolved.
While MFI’s core top calibration is relatively well established, its
down core applicability has yet to be demonstrated. Fig. 2 illustrates a down core comparison of MFI-based % CaCO3 preserved in
core RC13-110 from EEP with % CaCO3 in ODP 1089 from the Cape
Basin. By comparing CaCO3 records from different depths in the
South Atlantic, Hodell et al. (2001) concluded that the %CaCO3 at
site 1089 mainly reflects changes in CaCO3 preservation due to
varying [CO2
3 ] in bottom water. Hodell et al. (2001) further
concluded that the pattern of CaCO3 abundance at Site 1089 reflects
the widespread pattern of changes in chemistry of Indo-Pacific
bottom water, so it is to be expected that changes in CaCO3 preservation in the deep Cape Basin should parallel changes in the deep
Pacific. This conclusion was supported by the equatorial Pacific
results of Anderson et al. (2008), providing a basis for demonstrating the performance of the MFI proxy (Fig. 2). When age dating
uncertainties between the two cores are taken into account, the
correlation between the two cores and two proxies is striking and
provides evidence supporting MFI’s down core reliability.
2.2.2. Size-normalized foraminifer shell weight (SNSW)
The main assumption behind the SNSW method is that foraminifer tests within a specified size range become lighter with
increased dissolution (Lohmann, 1995; Broecker and Clark, 2001a,
2001b). This has been well-established for several species of
planktonic foraminifers including N. dutertrei, Pulleniatina obliquiloculata and Globigerinoides ruber (e.g. Broecker and Clark, 2001a,
2001b, 2003). Broecker and Clark (2001a) report an average size
normalized foraminifer weight loss slope of 0.30 0.05 mg/shell
per 1 mmol/kg decrease in depth-normalized [CO2
3 ]. SNSW data for
WEP cores presented herein are from Broecker et al. (2001).
2.2.3. 230Th-normalized carbonate accumulation rate
We use 230Th-normalization to estimate carbonate accumulation rate (CAR) in a subset of our cores as a foraminifer-independent
calcite preservation proxy. This approach provides accurate estimates of the vertical flux of CaCO3 to the seafloor by correcting for
post- or syn-depositional redistribution of sediment by bottom
currents (see François et al., 2004 for a detailed explanation).
However, it cannot distinguish between changes in CaCO3 export
from surface waters and changes in CaCO3 preservation on the
seafloor, unless several cores are analyzed from the same area but
taken at different depths (e.g. Francois et al., 1990).
This approach is based on the approximation that the scavenged
flux of 230Th from the water column is equal to its known
production rate from the decay of 234U dissolved in seawater
(Bacon, 1984). The 230Th concentration in sediments in excess of the
lithogenic and authigenic fraction (ex230Tho in dpm/g) can be used
to quantify the vertical rain rate of sediment using Eqs. (5) and (6):
Bulk Sedimentation Rate ¼ ðb*water depthðkmÞÞ=ex230 Tho
(5)
where b ¼ constant production rate of 230Th in seawater from 234U
radioactive decay (2.63 dpm/cm3/kaper km of water depth)
Carbonate Accumulation RateðCARÞ
¼ Bulk Sedimentation Rate*Fraction Carbonate
(6)
Estimates of CAR compiled from literature and used herein are
from Bradtmiller et al. (2006), Loubere et al. (2004), François et al.
(1990) and Kienast et al. (2007) (see Table 2 for detailed listing).
Thorium-normalized CAR data for two cores on the Ontong Java
Plateau (OJP), ERDC 125 and ERDC 131, have not been previously
published. The CaCO3 fraction in dry bulk sediment for these two
cores was measured by coulometry and ex230Tho was determined
by alpha spectrometry as outlined in François et al. (1993).
2.3. Modeling percent calcite dissolved
Fig. 2. Comparison of the MFI-based CaCO3 preservation record in RC13-110 in the
tropical Pacific (red; data from Loubere et al., 2004) and the percent CaCO3 in core ODP
1089 in the Cape Basin (blue; data from Hodell et al., 2001). (For interpretation of the
references to colour in this figure legend, the reader is referred to the web version of
this article.)
We used the computational model Muds_constcal (Archer et al.,
2002) to calculate the effects of bottom water CO2
3 undersaturation,
sedimentary organic carbon flux and CaCO3 flux on the percent
CaCO3 dissolved at the seabed. Muds_constcal is a model of pore
water pH and redox chemistry and is driven by the sinking fluxes of
organic carbon and CaCO3 to the seabed. The model uses the
chemistry of the overlying water column as a boundary condition.
See Mekik et al. (2002, 2010) and Archer et al. (2002) for more
information about using Muds to estimate CaCO3 dissolution rates
and the percent CaCO3 dissolved (percent preserved ¼ 100 percent
dissolved).
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
3. Down core preservation records
3.1. Pacific Ocean
Among nine cores from the EEP, six show a steady drop in MFIbased CaCO3 preservation from the Last Glacial Maximum (LGM,
w22 ka) to the present without a deglacial preservation maximum
(Fig. 3a), whereas three cores exhibit somewhat improved preservation during the deglaciation (Fig. 3b). Core ME-24 has a relatively
high sedimentation rate (see Table 1) and its MFI record indicates
that the deglacial increase in CaCO3 preservation was interrupted
by a brief interval of enhanced dissolution (Fig. 3b).
Among four cores from the OJP in the WEP, two (MW91-9
51GGC and 56GGC) show broad deglacial increases in SNSW for
N. dutertrei and one (56GGC) exhibits a corresponding increase in
SNSW for P. obliquiloculata (Fig. 4). In none of the cores does the
MFI-based index suggest increased CaCO3 preservation during
deglaciation (Fig. 4).
As was the case for the MFI records, some profiles of 230Thnormalized CAR from the tropical Pacific exhibit deglacial maxima,
consistent with enhanced CaCO3 preservation, whereas others
do not. In the WEP, four of five CAR records have maxima
65
corresponding to late deglacial or early Holocene periods (Fig. 5a).
Core RC17-177, with a deglacial CAR minimum, is the clear exception to this pattern. Both of the cores from the central equatorial
Pacific examined here have deglacial CAR maxima (Fig. 5b). In the
EEP, five of ten cores examined have no detectable increase in CAR
through the deglaciation (Fig. 5c), whereas the other half of the EEP
cores show enhanced CaCO3 accumulation during deglaciation
(Fig. 5d).
3.2. Atlantic and Indian Oceans
Thorium-normalized CAR were measured on two sets of three
cores taken at different depths on Ceara Rise (western equatorial
Atlantic) and Sierra Leone Rise (eastern equatorial Atlantic).
Carbonate preservation can be assessed by comparing CAR in the
deeper and shallower cores of each set, showing a clear minimum
during deglaciation (Fig. 6a and c). G. menardii disappeared from
the Atlantic Ocean during the LGM, so our MFI record only reaches
down to 13 ka. The MFI record is consistent with the CAR results in
that it indicates greater CaCO3 dissolution during deglaciation,
followed by improving preservation through the early Holocene
(Fig. 6b).
Core WIND 28 KA from the western tropical Indian Ocean has
a record of bottom water [CO2
3 ] estimated from B/Ca ratios in
benthic foraminifera (Yu et al., 2010) to compare with our MFIbased record of CaCO3 preservation. The two records are consistent during the Holocene in showing a decrease in CaCO3 preservation associated with declining [CO2
3 ] (Fig. 7). However, whereas
the B/Ca proxy indicates lower [CO2
3 ] during the LGM than during
deglaciation, the MFI record indicates consistently high CaCO3
preservation from the LGM through deglaciation before declining
during the Holocene.
4. Why is the deglacial CaCO3 preservation maximum missing
in most tropical/subtropical cores?
Fig. 3. Down core MFI-based % calcite preserved data for nine cores in the eastern
equatorial Pacific. See Table 1 for core depths and age models. The time interval
between 10 and 17 ka has been shaded to show the deglaciation. A: Cores showing no
deglacial increase in CaCO3 preservation, B: Cores showing a deglacial increase in
CaCO3 preservation.
The expected peak during the deglaciation in the preservation
(or accumulation) of CaCO3 is missing in more than half of the 31
cores examined here. Foraminifers from deglacial sediments are
visibly more dissolved (fragmented and thinner shells) than their
glacial counterparts in these cores (FM unpublished observations),
supporting the absence of a deglacial enhancement in CaCO3
preservation inferred from the other proxies.
Furthermore, multiple proxies within the same cores provide
conflicting records of CaCO3 preservation. For example, RC13-114
shows no deglacial calcite preservation peak with either MFI
(Fig. 3a) or with 230Th-normalized CAR (Fig. 5C) whereas Marchitto
et al. (2005) showed a clear deglacial spike in [CO23 ] of bottom
waters inferred from Zn/Ca in benthic foraminifers. Similarly, SNSW
of foraminifers shows a clear increase during the deglaciation in
MW91-9 56GGC while MFI shows no change (Fig. 4a).
While uncertainties in the age models of the cores studied here
(Tables 1 and 2) could confound the exact timing of the deglacial
carbonate preservation peak, for cores where this peak is observed
(e.g. RC 13-110) its timing is similar to the timing of carbonate
preservation peaks detected by other proxies in other cores
(w11e18 ka, broadly). For cores where we do not observe increased
carbonate preservation at any time after the LGM, uncertainties in
the age models do not influence our interpretation. The expected
CaCO3 preservation maximum is simply not observed at any time
during the deglaciation.
The lack of a consistent and unequivocal CaCO3 preservation
peak during the deglaciation in many proxy records is problematic
in light of the compelling evidence in support of the ventilation
hypothesis (see above). Furthermore, an alternative hypothesis that
66
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
Fig. 4. CaCO3 preservation record for 4 cores on the Ontong-Java Plateau. Blue points show MFI-based % CaCO3 preserved. Red symbols show N. dutertrei shell weight and green
symbols show P. obliquiloculata shell weight. The time interval between 10 and 17 ka has been shaded to show the deglaciation. See Tables 1 and 2 for sources of age models and
shell weight data. MW91-9 51GGC and MW91-9 56GGC were used by Broecker et al. (2001) but were cited as BC 51 and BC 56 in their work. (For interpretation of the references to
colour in this figure legend, the reader is referred to the web version of this article.)
does not require ventilation of CO2 from the deep ocean as a source
for the increase in atmospheric pCO2 is hard to construct. We do not
argue against the deglacial increase in deep ocean [CO2
3 ] because
the presence of the carbonate preservation peak in many of our 31
cores as well as carbonate preservation peaks observed in higher
latitude cores from the North Pacific (Jaccard et al., 2009, 2010)
clearly demonstrate that this was a global event. Instead, below we
explore possible causes for the disparate down core records of
CaCO3 preservation to learn more about the processes that may
influence the CaCO3 content of deep-sea sediments.
Fig. 5. Thorium-normalized CaCO3 accumulation rate (CAR) data for the tropical Pacific. A: Western equatorial Pacific, B: Central equatorial Pacific, C: Cores in the eastern equatorial
Pacific not showing a deglacial increase in CAR, D: Cores in the eastern equatorial Pacific showing a deglacial increase in CAR. The time interval between 10 and 17 ka has been
shaded to show the deglaciation.
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
67
Fig. 6. Thorium-normalized CaCO3 accumulation rate (CAR) and MFI-based percent CaCO3 preserved from the tropical Atlantic. A: CAR data for cores from the Ceara Rise. B: MFI
results for core GS7309-6PC from the central tropical Atlantic. C: CAR data for cores from the Sierra Leone Rise. The time interval between 10 and 17 ka has been shaded to show the
deglaciation.
4.1. Water depth issues
If CaCO3 preservation is mostly controlled by changes in DCO2
3
of bottom water, a preservation spike would only be apparent
within the depth range between the minimum depth of the CaCO3
saturation horizon (above which CaCO3 would be always 100%
preserved) and the maximum Carbonate Compensation Depth
(¼CCD; below which CaCO3 would always be 100% dissolved). In
addition, some indicators of CaCO3 dissolution are likely to be more
sensitive at greater depths, close to the CCD while others are likely
to be more sensitive at shallower depths, closer to the CaCO3
saturation horizon. For instance, when CaCO3 is the dominant
constituent of sediment, %CaCO3 is not sensitive to %CaCO3 dissolved near the CaCO3 saturation horizon (Broecker and Peng,
1987). On the other hand, close to the CCD, changes in DCO2
3 can
produce large relative changes in %CaCO3 (e.g. Jaccard et al., 2009,
2010). Changes in CAR are equally sensitive over the entire depth
range (230Th concentration is inversely proportional to %CaCO3
dissolved) but the relative changes in CAR are larger and more
easily discerned (and less easily muted by bioturbation) at greater
depths. This may partly explain why we find a clear preservation
spike in the relatively deep cores in the central equatorial Pacific
but not in the shallower cores from the WEP. In contrast, MFI
cannot detect % preserved less than 25% (Mekik et al., 2010) but
above that, MFI varies quasi-linearly with %CaCO3 preserved.
Therefore, MFI is particularly well suited to identify CaCO3 preservation spikes at shallower depths, closer to the CaCO3 saturation
horizon.
However, CaCO3 preservation is also affected by respiratory
CO2 released into sediment pore waters, which dissolves CaCO3
above the CaCO3 saturation horizon as well as below it. As a result,
CaCO3 preservation may also be sensitive to changes in DCO2
3
above the CaCO3 saturation horizon, since lower DCO2
3 in bottom
water would increase CaCO3 dissolution for a given release of
metabolic CO2.
4.2. MFI Detection Limits
Fig. 7. MFI-based CaCO3 preservation record in WIND 28KA from 4157 m in the
western Indian Ocean. Bottom water [CO2
3 ] was estimated by Yu et al. (2010) based on
B/Ca ratios in benthic foraminifera. See Table 1 for more core details. The time interval
between 10 and 17 ka has been shaded to show the deglaciation.
We address this issue in two steps. First, what is the expected
amplitude of the deglacial increase in [CO2
3 ] of deep water? This is
not easy to predict as it depends on location, on the initial and final
conditions of ocean stratification, and on the rate of the deglacial
transition (Marchitto et al., 2005). Marchitto et al. (2005)
summarized the results of a suite of models that predicted an
increase in [CO2
3 ] of deep Indo-Pacific water ranging between 15
and 30 mmol/kg.
Results from three empirically calibrated proxies are consistent
with a value near the upper end of this range. Broecker et al. (2001)
interpreted changes in the weight of planktonic foraminifera shells
recovered from the OJP to indicate a deglacial increase in [CO2
3 ] as
large as 30 mmol/kg at 4.0 km water depth. Marchitto et al. (2005)
interpreted Zn/Ca ratios of benthic foraminifera at 3.4 km water
depth in the EEP to indicate a deglacial increase in [CO2
3 ] of
w25 mmol/kg Yu et al. (2010) drew similar inferences from their B/
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F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
Ca data from a core in the western tropical Indian Ocean. Based on
this evidence, a range of 25e30 mmol/kg is a reasonable
expectation.
Second, what is the expected impact on CaCO3 dissolution of
a 25e30 mmol/kg increase in [CO2
3 ]? The answer will depend on
several factors, including the initial state of CaCO3 saturation in
bottom water and the organic carbon rain to the sea bed. For
example, at sites well above the CaCO3 saturation horizon with
relatively low organic carbon rain rates, an increase in bottom
water [CO2
3 ] will have negligible impact on CaCO3 dissolution.
However, at sites close to the CCD, an increase in [CO2
3 ] of
25e30 mmol/kg is expected to have a substantial impact on CaCO3
preservation.
Results from a study of equatorial sediments at 140 W provide
a basis for estimating the expected change in CaCO3 dissolution.
Berelson et al. (1997) used a sediment diagenesis model together
with measured rain rates of CaCO3 and of organic carbon to estimate that CaCO3 dissolution in the central equatorial Pacific Ocean
had increased during the late Holocene by 1.1e1.8 g/cm2/ka in
response to a decrease in bottom water [CO2
3 ] of 10e15 mmol/kg.
This increase in CaCO3 dissolution corresponds to approximately
half the measured CaCO3 rain rate, a change that would be detected
easily with MFI. This example illustrates the potential for large
changes in CaCO3 preservation associated with a deglacial increase
in [CO2
3 ], even if it is substantially smaller than the estimates cited
above (25e30 mmol/kg). If the deglacial increase in [CO2
3 ] had been
close to these estimates, then a preservation peak should be evident
at the sites below the calcite saturation horizon studied with MFI.
4.3. Increased deglacial organic carbon to calcite rain ratios in the
eastern equatorial Pacific
We performed sensitivity tests with modeling to address
whether or not a change in organic carbon rain rate that is
consistent with the observed sediment record can cause enough of
an increase in CaCO3 dissolution to offset the increase in CaCO3
preservation that would be expected from a 20e30 mmol/kg
deglacial increase in DCO2
3 of bottom water. The biogeochemical
model, Muds (Archer et al., 2002), was used to test the sensitivity of
CaCO3 preservation to changes in organic carbon rain rate, organic
carbon to CaCO3 rain ratio, and bottom water DCO2
3 (for modeling
details see Mekik et al., 2002, 2010). Specifically, we used Muds_constcal where the input parameters for each sample location are
water depth, organic carbon flux, %CaCO3 in sediments and DCO2
3 .
The output parameter is CaCO3 dissolution rate.
We modeled the sensitivity of CaCO3 dissolution to increasing
organic carbon rain rate for a constant bottom water DCO2
3 value
and constant CaCO3 rain rate of 30 mmol/cm2/yr, approximately the
global average (Milliman, 1993) and about 50% greater than
measured in the central equatorial Pacific (Berelson et al., 1997). For
modern conditions we chose a bottom water DCO2
3 value of 10 m
mol/kg based on average bottom water [CO2
]
in the EEP from
3
GLODAP bottle data (Key et al., 2004; Sabine et al., 2005),
and þ20 mmol/kg for the deglaciation (Marchitto et al., 2005). Deep
sediment traps in the EEP collected 7.5e16 mmol/cm2/yr of organic
carbon flux (Dymond and Lyle, 1993). Berelson et al. (1997) reported
organic carbon fluxes of 7e20 mmol/cm2/yr for the central equatorial Pacific, also from deep sediment traps. Consequently, we chose
15 mmol/cm2/yr as the baseline organic carbon rain rate for modern
EEP conditions. Although the model is overly simple, in that surfacesediment CaCO3 concentration is constant despite varying organic
carbon rain, our modeling results (Fig. 8) show that the modern %
CaCO3 dissolved can be achieved under elevated deglacial bottom
water DCO2
by raising the organic carbon rain rate between
3
a factor of two and three (follow the dashed arrow in Fig. 8).
Fig. 8. Modeling results showing relationship between organic carbon flux and %
CaCO3 dissolved in EEP sediments for conditions approximating those of the deglaciation and the late Holocene (see text for details). Modern and deglacial bottom water
DCO2
3 are set to 10 and þ20 mmol/kg, respectively. Calcite rain is a constant 30 mmol/
cm2/yr in all cases. The modern organic carbon rain rate is estimated at 15 mmol/cm2/yr
(see text). The “X” marks modern conditions of CaCO3 dissolution and the dashed line
from X to Y illustrates the increase in organic carbon rain rate needed to offset
a 30 mmol/kg increase in bottom water DCO2
3 to maintain a constant percent CaCO3
dissolved.
The preserved flux of organic carbon in EEP sediments was
much greater during deglaciation than during the LGM or during
the Holocene (Fig. 9; Kienast et al., 2007). Although the rain rate of
organic carbon need not scale linearly with its burial in sediments,
due to variable preservation, the results in Fig. 9 indicate that
a deglacial increase in organic carbon rain rate of between a factor
of two and three is not inconsistent with the sediment record. Of
course, if the CaCO3 rain scaled with the organic carbon rain rate, so
that there was little change in the rain ratio, then the model results
presented here will not be appropriate. Fortunately, recent
biomarker evidence for EEP sediments indicates that most of the
deglacial increase in organic carbon rain was associated with diatoms and not with coccolithophorids (Calvo et al., 2011), indicating
Fig. 9. Thorium-normalized organic carbon accumulation rate data for five cores in the
eastern equatorial Pacific (data from Kienast et al., 2007). The time interval between 10
and 17 ka has been shaded to show the deglaciation.
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
that the rain ratio was larger during deglaciation and that the
simple model used here can therefore be used to investigate the
impact of increasing organic carbon rain on CaCO3 preservation.
Keeping in mind the limitations of this modeling approach, we
conclude that it is plausible, and consistent with the sediment
record, for the expected deglacial CaCO3 preservation peak to have
been obscured in EEP sediments by increased CaCO3 dissolution
driven by enhanced production of respiratory CO2 (i.e., by the
deglacial increase in organic carbon to CaCO3 rain ratio).
4.4. Changes in circulation and bottom water chemistry
The potential for overprint by changes in circulation and water
mass distribution is particularly important for the Atlantic Ocean,
where deep waters of northern or southern origin with very
different corrosiveness for CaCO3 are vying for dominance. While
some of our Pacific cores seem to show a deglacial CaCO3 preservation maximum, MFI and 230Th-normalized CAR show greater
percent CaCO3 dissolved in the Atlantic Ocean during the deglaciation (Fig. 6). This could be explained in part by the high fluxes of
organic carbon reaching the seabed during the deglaciation, which
was a time of high surface ocean productivity (Loubere et al., 2003;
Bradtmiller et al., 2007). More important, perhaps, is the corrosive
effect of Antarctic Bottom Water (AABW), which spread well into
the North Atlantic Ocean during Heinrich Stadial 1 and Younger
Dryas (McManus et al., 2004; Robinson et al., 2005). Increased
CaCO3 dissolution below w4000 m depth during the period of
southern source water incursion was confirmed by CAR records
from two depth transects in the tropical Atlantic Ocean (Francois
et al., 1990). Deglacial carbonate accumulation minima in Atlantic
cores can be attributed to dissolution of calcite (Fig. 6a and c) by the
difference in 230Th e normalized carbonate fluxes between shallow
and deep cores (François et al., 1990). Consequently, an anticipated
global deglacial CaCO3 preservation maximum may have been
obscured at the deepest tropical Atlantic core sites by the enhanced
dissolution forced by an increased presence of AABW. On the other
hand, the deglacial pteropod preservation spike in the Atlantic reported by Berger (1977) indicates that ventilation of deep waters
may have raised [CO2
3 ] above LGM and Holocene levels at depths
shallower than 3500 m.
4.5. Proxy Issues
It is challenging to reconcile the preservation peaks observed in
SNSW (Broecker et al., 2001; Broecker and Clark, 2003) in down
core work on the OJP with MFI’s down core record in the same cores
69
(Fig. 4). The discrepancies are most likely due to the influence of
ecological and geochemical variations in ambient waters during life
on initial foraminifer shell weights (such as surface ocean [CO2
3 ];
Barker and Elderfield, 2002; Bijma et al., 2002). These variations
could bias the initial shell thickness of foraminifers and obliterate
their dissolution response in the sediment. Core top work in the
EEP by Mekik and Raterink (2008) supports this hypothesis by
showing that initial SNSW related to [CO2
3 ] of habitat waters is
clearly discernable even in foraminifer tests from sediments that
have high MFI values. That is, the initial shell weight signal is not
erased even under conditions of substantial post-depositional
CaCO3 dissolution.
Core RC13-114 (Fig. 10) is another example where multiple
proxies tell different stories in the same core. While neither MFIbased %calcite dissolved nor 230Th-normalized CAR (Bradtmiller
et al., 2006) seem to have a carbonate preservation maximum
between 11 and 17 ka, reconstruction of bottom water DCO2
3 using
Zn/Ca ratios shows two large deglacial peaks, one during deglaciation and one during the early Holocene (Marchitto et al., 2005).
The second, younger peak is consistent with CAR data but the older
deglacial peak is not corroborated by either of the other two
proxies. However, artifacts can influence Zn/Ca ratios in benthic
foraminifera (Marchitto et al., 2005). For example, at ODP Site 849
in the EEP, MnCO3 overgrowths dominated the Zn/Ca signal
throughout most of the core. Furthermore, in core RC13-114, which
has a well-defined Zn/Ca maximum during Termination I, there is
no corresponding maximum during Termination II. Instead, a Zn/Ca
maximum of reduced amplitude is observed after the termination,
during Marine Isotope Stage (MIS) 5. In addition, bottom water
DCO2
inferred from Zn/Ca of core top samples in RC13-114 is
3
greater than DCO2
3 inferred for MIS 3, which is inconsistent with
other results suggesting that CaCO3 preservation was at nearmaximum values during MIS 3 (Hodell et al., 2001; Anderson
et al., 2008).
While the MFI proxy tracks the evolution of the bulk carbonate
system much more closely than other dissolution proxies, a consequence of its calibration against a model of sedimentary CaCO3
diagenesis (Mekik et al., 2002; Mekik et al., 2006; Mekik and
Raterink, 2008; Mekik et al., 2010), it is not without its shortcomings. Based on its current calibration, it has an accuracy of 10e15%
calcite dissolved (Mekik et al., 2002, 2010). By contrast, the precision of MFI measurements, which determines the ability to detect
downcore changes in CaCO3 preservation, is 0.04 MFI units
(Mekik et al., 2010), which corresponds to w2e4% calcite dissolved.
So, a change in % dissolved of greater than 4% among down core
samples should be detectable by MFI. Sample breakage resulting
Fig. 10. Comparison of proxy records related to CaCO3 preservation in RC13-114 from the eastern equatorial Pacific. CaCO3 accumulation rate (CAR), MFI-based fraction calcite
preserved and DCO2
3 reconstruction based on Zn/Ca ratios in benthic foraminifera. The time interval between 10 and 17 ka has been shaded to show the deglaciation. See Tables 1
and 2 for data sources.
70
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
from bioturbation or handling could potentially contribute to
fragmentation of G. menardii tests and potentially confound MFI
measurements. However, fragmentation due to physical breakage
is often marked by angular fragments, whereas dissolution-related
fragmentation tends to produce more rounded fragments and holes
within fragments. Visual assessment of samples was made to
ensure that G. menardii fragments were generally not the result of
physical breakage. Furthermore, based on the work of Barker and
Elderfield (2002) it can be argued that glacial G. menardii may
have had thicker shells due to higher [CO2
3 ] in calcification waters
during that time. This could negatively affect the reliability of MFI
in glacial age sediments. This idea is certainly plausible, but Mekik
and Raterink (2008) have shown from core top sediments in the
tropical Pacific that [CO2
3 ] of foraminiferal growth waters and MFI
show no relationship. For example, at the same surface-water
[CO2
3 ] (w180 mmol/kg), MFI can range from 0.4 to 1; and at the
same MFI value (w1), surface [CO2
3 ] can range from 90 to
200 mmol/kg. Mekik and Russo have new unpublished data
showing that the SNSW of G. menardii tests are minimally affected
by the [CO2
3 ] of their habitat waters and mostly trace CaCO3
dissolution in deep-sea sediments.
4.6. Western equatorial Pacific and Indian Ocean cores
While deglacial changes in water mass distributions and in rain
ratios are plausible explanations for the obscured deglacial CaCO3
preservation maximum in the Atlantic and EEP regions, respectively, the lack of a MFI-based CaCO3 preservation peak in WEP
cores (Fig. 4) and in one core in the western Indian Ocean (Fig. 8) is
difficult to reconcile with the ventilation hypothesis. Very low
sedimentation rates in cores from the OJP (Fig. 4; Table 1) may
account for the seeming lack of a preservation maximum. That is,
bioturbation may have homogenized the sediments to the point
that the MFI cannot pick up the deglacial signal (Fig. 4a), even
though a maximum is evident in the CAR records (Fig. 5a).
Certainly, more work in deeper cores with higher sedimentation
rate is necessary to resolve this issue.
Our results from the Indian Ocean (WIND 28K; Fig. 7) do not
exhibit a clear deglacial calcite preservation peak, even though the
MFI proxy faithfully records the Holocene trend of increasing CaCO3
dissolution observed throughout the deep Indian and Pacific
Oceans (Berger, 1977; Broecker et al., 2001; Hodell et al., 2001;
Broecker and Clark, 2001b; Marchitto et al., 2005; Anderson et al.,
2008; Yu et al., 2010; Lalicata and Lea, 2011). It is possible that
a change in rain ratio obscured the deglacial CaCO3 preservation
maximum there, as in the EEP, but we do not have sufficient
information about this site to explore this possibility.
The data in Loubere et al. (2004) show focusing factors that
generally decreased from the LGM to the present in several cores in
the EEP. This would lead to better preservation of CaCO3 in these
cores during the LGM when compared to subsequent times simply
as a result of buffering the sediment pore waters with CaCO3 that
was redistributed and focused laterally. Consequently, reduced
sediment focusing in these EEP cores following the LGM may have
complemented the higher rain ratio (Section 4.3) in obscuring the
anticipated deglacial CaCO3 preservation maximum.
5. Conclusions
There is abundant evidence in support of the deglacial ventilation hypothesis: the widespread distribution of 13C-depleted
carbon that invaded the upper ocean and atmosphere during
deglaciation (Smith et al., 1999; Spero and Lea, 2002; Köhler et al.,
2005), the drop during deglaciation in 14C activity of dissolved
inorganic carbon in North Pacific intermediate waters (Marchitto
et al., 2007), and clear records of improved deglacial CaCO3 preservation in high latitude cores (Jaccard et al., 2009, 2010). We also
see a clear deglacial carbonate preservation peak with MFI in RC13110 (EEP) and in CAR of RC11-238 (EEP). B/Ca ratios in benthic
foraminifera indicate a deglacial increase in deep-water [CO2
3 ] at
several locations (Yu et al., 2010). This gives us confidence that deep
water [CO2
3 ] was higher during the deglaciation compared to the
LGM and to the Holocene. However, despite using a multi-proxy
approach on a large number of cores over a large geographic area,
we do not find unequivocal and globally traceable evidence of
a deglacial CaCO3 preservation maximum in deep-sea sediments.
Therefore, it is likely that the expected deglacial CaCO3 preservation
maximum is obscured by other factors.
One of these factors may be the lack of an ideal CaCO3 preservation proxy, which leads to disagreement between records of
CaCO3 preservation among different proxies. Other factors that
likely affected CaCO3 preservation in certain oceanic regions
include changes in the organic carbon to CaCO3 rain ratio reaching
the seabed (EEP), changes in ocean circulation patterns (tropical
Atlantic) and changes in sediment focusing (EEP).
While the deglacial ventilation hypothesis remains the best
explanation for the source of atmospheric CO2 after 18 ka, additional CaCO3 preservation records from subtropical and midlatitude regions will more clearly define the geographic extent of
the expected deep-sea CaCO3 preservation maximum. Our best
chances of finding additional evidence for a deglacial CaCO3 preservation spike from %CaCO3 or CAR would be in sediment cores
with relatively high sediment accumulation rates collected at
depths close to the CCD. For shallower cores, developing MFI
records in locations not affected by large changes in rain ratio or by
changes in deep-water circulation is also a promising approach.
4.7. Sediment focusing
Acknowledgments
The accumulation rate data used herein have all been corrected
for sediment redistribution by 230Th e normalization. However,
whereas the modeled dissolution rates are independent of the
focusing factor, the %CaCO3 dissolved values are not. Percent CaCO3
dissolved is determined by dividing the CaCO3 dissolution rate by
the total CaCO3 supply, which includes net lateral transport by
focusing as well as the vertical rain from above. For example, let us
suppose that the CaCO3 vertical rain rate were 1 g/cm2/ka and the
dissolution rate were 0.5 g/cm2/ka. If we assume a focusing factor of
1, then the % CaCO3 preserved is 50%. But if the focusing factor is 2,
then the total CaCO3 supply to the site (1 g/cm2/ka vertical rain and
1 g/cm2/ka by lateral transport) becomes 2 g/cm2/ka, while the
dissolution rate remains 0.5 g/cm2/ka. This would raise the % CaCO3
preserved to 75%. This point was illustrated by Berger (1992).
We extend many thanks to core curators at Oregon State
University (June Padman and Bobbi Conard), Lamont Doherty
Earth Observatory (Rusty Lotti-Bond) and the Ocean Drilling
Program for providing us with samples. Special thanks to Dan
McCorkle and Ellen Roosen at Woods Hole Oceanographic Institution for speedily providing us with samples for cores from the
Ontong-Java Plateau. Also many thanks to Andy Ridgwell for very
stimulating discussions and Tom Marchitto for providing us with
data. Lastly, we would like to thank three anonymous reviewers
whose thoughtful comments improved our manuscript. This work
was supported by grants to Mekik from NSF (OCE0326686 and
OCE0825280) and from the NASA Michigan Space Grant Consortium, Seed Grant, 2001.
F.A. Mekik et al. / Quaternary Science Reviews 39 (2012) 60e72
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