Quaternary Science Reviews 29 (2010) 193–205
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The role of Southern Ocean processes in orbital and millennial
CO2 variations – A synthesis
Hubertus Fischer a, b, c, *, Jochen Schmitt a, b, c, Dieter Lüthi a, b, Thomas F. Stocker a, b, Tobias Tschumi a, b,
Payal Parekh a, b, Fortunat Joos a, b, Peter Köhler c, Christoph Völker c, Rainer Gersonde c,
Carlo Barbante d, e, Martine Le Floch e, f, Dominique Raynaud e, f, Eric Wolff f, g
Climate and Environmental Physics, Physics Institute, University of Bern, Switzerland
Oeschger Centre for Climate Change Research, University of Bern, Switzerland
Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany
Department of Environmental Sciences, University of Venice, Ca’ Foscari, Italy
Institute for the Dynamics of Environmental Processes-CNR, Venice, Italy
Laboratoire de Glaciologie et Geophysique de l’Environnement (LGGE), CNRS, St Martin d’Hères, France
British Antarctic Survey, Cambridge, UK
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 19 January 2009
Received in revised form
2 June 2009
Accepted 4 June 2009
Recent progress in the reconstruction of atmospheric CO2 records from Antarctic ice cores has allowed
for the documentation of natural CO2 variations on orbital time scales over the last up to 800,000 years
and for the resolution of millennial CO2 variations during the last glacial cycle in unprecedented detail.
This has shown that atmospheric CO2 varied within natural bounds of approximately 170–300 ppmv but
never reached recent CO2 concentrations caused by anthropogenic CO2 emissions. In addition, the
natural atmospheric CO2 concentrations show an extraordinary correlation with Southern Ocean climate
changes, pointing to a significant (direct or indirect) influence of climatic and environmental changes in
the Southern Ocean region on atmospheric CO2 concentrations.
Here, we compile recent ice core and marine sediment records of atmospheric CO2, temperature and
environmental changes in the Southern Ocean region, as well as carbon cycle model experiments, in
order to quantify the effect of potential Southern Ocean processes on atmospheric CO2 related to these
orbital and millennial changes. This shows that physical and biological changes in the SO are able to
explain substantial parts of the glacial/interglacial CO2 change, but that none of the single processes is
able to explain this change by itself. In particular, changes in the Southern Ocean related to changes in
the surface buoyancy flux, which in return is controlled by the waxing and waning of sea ice may
favorably explain the high correlation of CO2 and Antarctic temperature on orbital and millennial time
scales. In contrast, the changes of the position and strength of the westerly wind field were most likely
too small to explain the observed changes in atmospheric CO2 or may even have increased atmospheric
CO2 in the glacial. Also iron fertilization of the marine biota in the Southern Ocean contributes to a glacial
drawdown of CO2 but turns out to be limited by other factors than the total dust input such as
bioavailability of iron or macronutrient supply.
Ó 2009 Elsevier Ltd. All rights reserved.
1. Introduction
Carbon dioxide represents one of the most important greenhouse gases in the Earth’s atmosphere, second only to the
atmospheric content of water vapor (IPCC, 2007). Accordingly,
significant variations in its atmospheric concentration lead to
a respective change in radiative forcing. Due to anthropogenic
* Corresponding author. Climate and Environmental Physics, Physics Institute,
University of Bern, Switzerland. Tel.: þ41 31 631 85 03; fax: þ41 31 631 87 42.
E-mail address:[email protected] (H. Fischer).
0277-3791/$ – see front matter Ó 2009 Elsevier Ltd. All rights reserved.
emissions of CO2 by fossil fuel burning and land use changes, CO2
has increased over the last 150 years from a preindustrial level of
approximately 280 ppmv to 385 ppmv in recent years, as documented in both atmospheric ( and ice
core records (Etheridge et al., 1996; Indermühle et al., 1999;
MacFarling Meure et al., 2006). From these data one can confidently state that the recent rate of increase (up to 2 ppmv/year)
has not been seen at decadal or longer timescales for at least the
last 20,000 years (Joos and Spahni, 2008) and past CO2 concentrations in ice cores over the last 800,000 years have never come
close to current CO2 levels (Petit et al., 1999; Siegenthaler et al.,
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
2005; MacFarling Meure et al., 2006; Lüthi et al., 2008). Resulting
from this anthropogenic CO2 increase is a change in radiative
forcing of about 1.7 W/m2 which is in large parts responsible for
the 0.7 C warming observed over the last century (IPCC, 2007).
Today, the global carbon cycle is being driven out of equilibrium
primarily by the burning of fossil fuels and deforestation leading to
a continuing increase in atmospheric CO2 since the start of the
industrialization. However, apart from the increasing anthropogenic emissions it is the carbon turnover and uptake capacity of the
ocean and terrestrial biosphere that determines the long-term fate
of this anthropogenic perturbation and its radiative forcing. In the
pre-anthropocene, the natural shift from glacial to warm climate
conditions initiated a significant increase in atmospheric CO2 and in
terrestrial biospheric carbon stocks while reducing the carbon
storage in the deep ocean. In return, higher natural CO2 levels lead
an additional warming of the atmosphere via their feedback on the
radiative balance (Fischer et al., 1999; Caillon et al., 2003). These
natural CO2 changes are well archived in Antarctic ice cores (Neftel
et al., 1982; Stauffer et al., 1998; Fischer et al., 1999; Indermühle
et al., 1999; Petit et al., 1999; Monnin et al., 2001; Siegenthaler et al.,
2005; Ahn and Brook, 2008; Lüthi et al., 2008) and show about
100 ppmv lower CO2 concentrations during glacial than interglacial
times as well as significant millennial CO2 variations during glacial
periods. Accordingly, the global carbon cycle is intimately coupled
to long-term climate changes. Besides the necessity of an increased
radiative forcing by CO2 and other greenhouse gases in the atmosphere to explain the glacial/interglacial temperature rise (Köhler
et al., 2010) the climate induced changes in the carbon cycle and
their consequence for atmospheric CO2 also represent an important
testbed for hypotheses invoking changes in ocean circulation or
marine and terrestrial biogenic productivity in the past. For
instance any theory invoking changes in ocean circulation would
affect the carbon storage in the ocean and, thus, would also have to
consider its effect on atmospheric CO2.
Despite the importance for climate changes today and in the
past, the glacial/interglacial CO2 changes could not be accounted for
in carbon cycle models until recently. In recent model experiments,
however, it has become possible to quantitatively explain the
100 ppmv glacial/interglacial shift in atmospheric CO2 concentrations by combining the effects of different processes acting on the
global carbon cycle, such as sea surface temperature (SST) and
salinity changes, gas exchange, ocean circulation, marine biological
export production, terrestrial carbon storage, and carbonate
compensation in the deep ocean (Archer et al., 2000a; Sigman and
Boyle, 2000; LeGrand and Alverson, 2001; Paillard and Parrenin,
2004; Köhler et al., 2005; Köhler and Fischer, 2006; Brovkin et al.,
2007). Although the total 100 ppmv change can now in principle be
accounted for, the contributions of each individual process to the
overall change still carry substantial uncertainties that do not allow
for a unique solution to the problem. However, all the models agree
that changes in the biological or physical carbon fluxes in the
Southern Ocean (SO) connected to export production of organic
material at the surface and SO circulation changes, together with
their carbonate compensation feedback (Broecker and Peng, 1987)
in the deep ocean, represent the most important factors influencing
atmospheric CO2 on orbital time scales. Former comparisons of
simulated atmospheric CO2 to changes in high latitude processes
suggested that more complex models are in some aspects less
sensitive on SO processes than simpler carbon cycle box models
(Archer et al., 2000b; Broecker et al., 1999). However, the comparison of completely different types of models is not straightforward
(Lane et al., 2006) and a given model maybe high latitude sensitive
to certain parameters but not to others. The challenge remains to
explain the low glacial CO2 in a self-consistent 3-dimensional
dynamical model setting.
An important role of the SO in terms of the carbon cycle is also
suggested by a very high correlation of atmospheric CO2 and
Antarctic temperatures over the last 800,000 years (Wolff et al.,
2005; Lüthi et al., 2008), the latter being reliably archived in the
stable water isotope record in Antarctic ice cores (Jouzel et al.,
2007). Despite this agreement on the role of the SO for atmospheric
CO2 levels, disagreement exists on the importance of ocean circulation vs. export production changes, and about the physical
processes that cause ocean circulation changes and affect carbon
fluxes in the SO (Sigman and Boyle, 2000; Matsumoto et al., 2002;
Köhler et al., 2005; Toggweiler et al., 2006; Watson and Garabato,
2006; Menviel et al., 2008; Parekh et al., 2008; Tschumi et al., 2008;
Martı́nez-Garcia et al., 2009).
The goal of this paper is to discuss the potential SO processes
that can lead to a glacial drawdown of atmospheric CO2 and to
confront these hypotheses with the various marine sediment, ice
core and modeling evidence. To this end, we will review the latest
ice core observations on glacial/interglacial and millennial CO2
changes, the theoretical background of SO circulation changes and
their role for carbon storage in the deeper ocean, as well as
modeling exercises to constrain those changes and their effect on
atmospheric CO2. We intentionally do not discuss carbon cycle
processes outside the SO region, which are either rather well
constrained (such as global SST and salinity) or beyond the scope of
this paper (such as the terrestrial biosphere or carbonate
compensation). However, it has to be kept in mind, that although
SO processes dominate the atmospheric CO2 changes in the past,
the full CO2 story can only be told when processes outside the SO
are also taken into account.
2. Ice core data of atmospheric CO2 changes
Bubble enclosures in polar ice cores represent the only direct
atmospheric archive that allows for reconstruction of the atmospheric composition over hundred thousands of years. In the case of
CO2, Antarctic ice cores represent the only unaltered archive of past
atmospheric CO2 concentrations, while in Greenland ice cores high
carbonate and low pH in combination with higher concentrations
of organic impurities in the ice leads to in situ production of CO2
(Anklin et al., 1995; Smith et al., 1997). In Antarctic ice cores, no
such artifacts have been observed so far as illustrated by the low
scatter of high-resolution samples in bubble ice (Indermühle et al.,
1999; Stauffer and Tschumi, 2000; Monnin et al., 2001; Lüthi et al.,
2008) that is below the analytical uncertainty, by the consistency of
Antarctic ice core records from different locations, and by the very
good agreement of ice core data with atmospheric measurements
in the overlapping interval between 1958 and 1975 (Etheridge et al.,
1996; MacFarling Meure et al., 2006).
However, due to the slow bubble enclosure process, the age of
the gas enclosed in air bubbles at a particular depth has a certain
distribution, leading to an inherent low-pass filtering of the
dynamics in atmospheric CO2 in the ice core archive (Schwander
and Stauffer, 1984). Because of this process, ice cores covering the
last glacial cycle and beyond (which are necessarily located in low
accumulation regions) only provide CO2 records with an effective
resolution of a few centuries (Spahni et al., 2003). The lowest
resolution is achieved during cold climate conditions, when snow
accumulation is reduced by about 50% compared to present
conditions (EPICA community members, 2004; EPICA community
members, 2006).
Fig. 1 shows a compilation of the CO2 concentrations from the
EPICA (European Project for Ice Coring in Antarctica) Dome C
(EDC) and Vostok ice cores, which as a whole cover the last
800,000 years (Fischer et al., 1999; Petit et al., 1999; Monnin et al.,
2001; Siegenthaler et al., 2005; Lüthi et al., 2008). Also shown are
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
Fe flux
[°/ ]
°° -410
EDC3/GT4 age [yr BP]
Fig. 1. CO2 concentrations measured on the EDC (red dots (Monnin et al., 2001; Siegenthaler et al., 2005; Lüthi et al., 2008)) and Vostok ice cores (filled orange dots (Petit et al.,
1999), open orange circles (Fischer et al., 1999)). Bold dashed blue lines represent a CO2 estimate derived from the linear regressions in Fig. 2. Also plotted are continuous stable
water isotope records (dD, black and orange lines) derived on the same cores (Petit et al., 1999; EPICA community members, 2004; Jouzel et al., 2007) in 500 yr resolution together
with spotwise iron fluxes (purple line) for EDC interpolated to 500 yr resolution (Martı́nez-Garcia et al., 2009). Ages are given on the EDC3 and GT4 age scale for the EDC and the
Vostok ice cores, respectively. Bold numbers indicate interglacial MIS, the yellow bar the time interval from MIS 16 to 18, when CO2 is lower than expected from a linear regression.
temperature proxy records (dD of water) in 500 yr resolution (i.e.
approximately the temporal resolution that can be achieved in
glacial CO2 records at these low accumulation sites) measured on
the same ice cores (Petit et al., 1999; EPICA community members,
2004; Jouzel et al., 2007). Note that the Vostok and EDC ice core
records are plotted on individual time scales, explaining the
significant differences in the timing of Marine Isotope Stages (MIS)
7 and 9 in Fig. 1.
The CO2 concentrations vary between 170–190 ppmv during
glacials and 250–300 ppmv during interglacials, with generally
lower interglacial concentrations prior to 450,000 years before
present (BP, where present is defined as 1950). In this time period,
interglacial temperatures were also significantly lower in
Antarctica than during the last four glacial cycles. Although the
amplitude of glacial/interglacial temperature changes in East
Antarctic is higher than for the SO sea surface temperatures (Gersonde et al., 2005), the temporal changes in Antarctic and SO
temperature (Martı́nez-Garcia et al., 2009) and sea ice coverage
(Wolff et al., 2006) are strongly linked. Accordingly, we assume that
the temporal evolution of the precise and high-resolution
temperature changes reconstructed from the EPICA ice cores are at
first order representative for the whole circum-Antarctic region
(Watanabe et al., 2003; EPICA community members, 2006; Martı́nez-Garcia et al., 2009).
A linear regression (Fig. 2) of CO2 and dD over the entire data
sets explains 71% and 84% of the variance in the CO2 data for the
Vostok and EDC ice cores, respectively. Note that this correlation is
only representative for the time scales of CO2 variations that can be
resolved in ice cores from the central East Antarctic ice sheet, i.e. on
the order of a few centuries. The somewhat larger scatter of the
Vostok data may be in part due to the fact that some of the
measured depth intervals come from the clathrate formation zone
between approximately 700 and 1300 m depth. In that depth
interval, the analytical uncertainty is higher because of different
extraction efficiencies of air from bubbles vs. clathrates, in combination with a fractionation of the CO2 concentration in clathrates
relative to that in air bubbles (Stauffer and Tschumi, 2000). A high
(0 - 420,000 yrs)
Dome C
(0-20,000 yr BP)
(400,000- 450,000 yr BP)
(450,000- 650,000 yr BP)
(650,000- 800,000 yr BP)
δD [°/ ]
Fig. 2. CO2 concentrations vs. dD as shown in Fig. 1. Respective dD values for each CO2
value were calculated by linear age interpolation of the dD values in 500 yr resolution
to ensure a temporal resolution comparable to the one for glacial CO2 data. Straight
lines represent two-sided linear regressions over the entire data sets. Note, that EDC
CO2 concentrations prior to 450,000 yr BP do not span the whole CO2 range but show
a similar linear relationship with temperature. Only CO2 data during MIS 16–18 are
systematically reduced by about 10 ppmv.
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
correlation between CO2 and dD (r2 ¼ 0.64) has previously been
observed in the Vostok ice core over the last glacial cycle (Cuffey
and Vimeux, 2001). Correcting the dD record for changes in the
isotopic composition and temperature of the water vapor source
using the deuterium excess, Cuffey and Vimeux (2001) were able to
show that this correction improves the correlation even further and
that such corrected Antarctic temperatures can account for more
than 84% of the CO2 variance. Note, that these very high correlation
values are representative for unlagged correlation, ignoring
a potential phase shift between CO2 and temperature, which may
increase the correlation even further. Thus, taking the Antarctic
isotope temperature records as a surrogate for the temporal
evolution of the entire SO ocean region more than 80% of the
multimillennial variability in CO2 concentrations could be attributed to circum-Antarctic temperature changes, in case some
physical link between the two can be established. Fig. 1 also shows
time intervals, where the deviations from this linear relationship
with Antarctic temperature are the largest, i.e. during interglacials
and the glacial inceptions, where processes outside the SO (such as
changes in the terrestrial carbon storage) may offset the balance
between the atmosphere and the SO. In addition, CO2 concentrations between MIS 16 and 18 are approximately 10 ppmv lower
than expected from the linear regression.
A strong connection between CO2 data and Antarctic temperature in centennial resolution also holds for millennial CO2 variations during the last ice age. Recently, Ahn and Brook (2008)
published a high-resolution CO2 record from the Byrd ice core over
the last glacial period as shown in Fig. 3, which is also in very good
agreement with CO2 data from Taylor Dome over the time interval
25,000–65,000 yr BP (Indermühle et al., 2000) as well as unpublished data from the EPICA ice core from Dronning Maud Land
(EDML) (Lüthi, unpublished). Since the accumulation rate at Byrd
is about 3–4 times higher than at EDC, this ice core allows for an
effective resolution of better than 200 years, even for glacial
conditions. The main feature of this CO2 record covering the last
glacial are concentrations elevated by about 20 ppmv in parallel to
the strongest Antarctic Isotope Maxima (AIM) (EPICA community
members, 2006). These AIM events are related to Dansgaard
Oeschger events 8, 12, 14, 17, 19, 20 and 21 in the northern
hemisphere (North Greenland Ice Core Project members, 2004;
Jouzel et al., 2007) via the bipolar seesaw (Stocker and Johnsen,
2003; EPICA community members, 2006). Despite the obvious
covariance in Fig. 3, the correlation between CO2 and d18O
measured on the Byrd ice core is not very strong (r2 ¼ 0.25 for the
entire data set from 20 to 90,000 yr BP and r2 ¼ 0.2 for the time
interval 30–90,000 yr BP) although significantly different from
zero (Fig. 4). A comparable CO2 record from the EDML ice core
(Lüthi, unpublished data) exhibits essentially the same CO2 variations as in the Byrd record but yields a much higher correlation
(Fig. 4) with EDML d18O values in 300 yr resolution (r2 ¼ 0.69 for
the entire data set from 30 to 110,000 yr BP and r2 ¼ 0.63 for the
time interval shared with the Byrd record from 30 to 90,000 yr BP).
This difference in the correlation coefficients could be explained in
two ways: i) the glacial Byrd d18O data show a higher variability
compared to the EPICA records, which may be in part related to
local temperature changes potentially connected to altitude
changes at Byrd, which is located on the West Antarctic Ice Sheet,
while the EPICA ice cores and Vostok are located on the more
stable East Antarctic Ice Sheet; ii) an insufficient age synchronization between the gas and ice record at Byrd may deteriorate the
correlation. Interestingly, a very high correlation (r2 ¼ 0.93) on
a millennial time scale is also clearly documented for the oldest
glaciation (750,000–778,000 yr BP) in the EDC record as shown in
Fig. 3.
In summary, we find that a large part of the variance is shared
between CO2 concentrations and Antarctic temperatures. This fact
calls for processes that connect atmospheric CO2 to climate changes
occurring in the SO region. In case such a link can be convincingly
established, the high correlation between CO2 and dD implies that
more than 80% of the glacial/interglacial and more than 60% of the
millennial CO2 variations during the last glacial could be ascribed to
such SO processes.
8 12 14 17
19 20 21
[°/ ]
GISP2/EDC3 age [yr BP]
EDC3 age [yr BP]
Fig. 3. Millennial CO2 variations over the last glacial (left) and the glaciation from MIS19 to MIS 18 (right) measured on the Byrd (blue dots) and the EDC (red dots) ice cores. Also
given are the isotope temperature records of both ice cores (Byrd d18O (left axis), pink line in 200 yr resolution (Ahn and Brook, 2008); EDC dD (right axis), black line in 500 yr
resolution (EPICA community members, 2004; Jouzel et al., 2007)) and the EDC non-sea-salt (nss) Ca2þ (purple line) as a mineral dust indicator. Ages are given on the Greenland Ice
Sheet Project (GISP) 2 age scale and the EDC3 age scale for Byrd and EDC, respectively. Bold numbers indicate Antarctic Isotope Maxima and the respective Dansgaard Oeschger
events in Greenland (North Greenland Ice Core Project members, 2004; EPICA community members, 2006; Jouzel et al., 2007).
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
(30,000 - 110,000 yr BP)
(20,000 - 90,000 yr BP)
δ18O [°/ ]
Fig. 4. CO2 concentrations vs. d18O isotope temperatures for the Byrd ice core (purple
dots) as shown in Fig. 3 as well as EDML CO2 concentrations (Lüthi, unpublished data)
vs. EDML d18O after correction for sea level and upstream altitude effects (EPICA
community members, 2006) (orange dots). Respective isotope values for each CO2
value were calculated by linear age interpolation based on the d18O values in 200 yr
resolution for Byrd and d18O values in 300 yr resolution for EDML to ensure a temporal
resolution comparable to the CO2 records. The straight lines represent two-sided linear
regressions over the entire data sets.
3. Southern Ocean carbon cycle processes
Current understanding essentially suggests that in addition to
the effect of the higher solubility of CO2 in SO surface waters during
colder climate periods, which accounts for about 15 ppmv lower
glacial atmospheric CO2 (Köhler and Fischer, 2006), two classes of
SO processes could account for such a correlation: (i) changes in
marine biological productivity in the SO, and (ii) changes in SO
circulation (advection or mixing), that vary the exchange between
carbon-enriched deep waters and the surface and change the
degree of nutrient utilization at the surface. In general, photosynthetic production of organic material in the surface water and the
subsequent export of that organic matter to the deep ocean
depletes atmospheric CO2 and at the same time enriches the dissolved inorganic carbon content of the deep ocean by respiration.
The ocean circulation has the opposite effect on atmospheric CO2,
bringing this phytoplankton-mediated carbon back to the surface.
In the following, we will briefly review the background of those two
classes of processes and discuss their influence on atmospheric CO2
based on observations and modeling evidence.
3.1. Marine biological productivity changes
In contrast to most other ocean basins, the SO is characterized
by high macronutrient (nitrate and phosphate) concentrations in
surface waters and, yet, a limited biological productivity, as
reflected by a relatively low chlorophyll content. This implies that
the phytoplankton in the SO is not able to quantitatively use the
available nutrients, probably related to a limitation of micronutrients such as iron. This leads to a large fraction of the available nutrients being unutilized before they are exported again as
so-called ‘‘preformed’’ nutrients to deeper ocean layers (Marinov
et al., 2008). Accordingly, it has been recognized that a relaxation
or complete removal of the iron limitation by enhanced iron
deposition connected to an increase in aeolian dust could lead to
increased carbon fixation in the SO biosphere (if ocean circulation
remained unchanged). Connected to an increased export of this
organic matter to the deep ocean, this could lead to a draw down
of atmospheric CO2 (Martin, 1990). This iron fertilization
hypothesis was directly tested for recent conditions by ship-based
iron fertilization experiments in the SO (Boyd et al., 2000; de Baar
et al., 2005). These experiments, showed a clear increase in
primary productivity but very little evidence of increased carbon
export after artificial iron input into the surface waters, questioning the effectiveness of the iron fertilization in drawing down
atmospheric CO2.
Elevated dust fluxes are unambiguously recorded in the EDC ice
core over the last eight glacials (Fig. 1) and the flux of soluble iron in
the ice cores increased by a factor of 20 during glacial times (Gaspari et al., 2006; Wolff et al., 2006; Lambert et al., 2008; Martı́nezGarcia et al., 2009). From this it can be deduced that the aeolian iron
deposition into the SO surface waters was significantly higher
during glacial times as well. In fact, recently derived Th-normalized
iron fluxes in a marine sediment record from the Atlantic sector of
the SO show dust fluxes higher by a factor of 4–5 in the glacial
compared to the Holocene (Martı́nez-Garcia et al., 2009).
Besides these orbital variations in iron deposition, what observational evidence for a possibly enhanced CO2 uptake in the SO by
iron fertilization exists for the last glacial? In absence of a continuous Fe record over MIS 3, we use the ice core dust derived Ca2þ
records as a proxy for iron deposition. The ice core dust records
show clear dust variations in the last glacial (Röthlisberger et al.,
2004; Fischer et al., 2007) in parallel to the millennial CO2 variations (Fig. 3). Note, however, that the long-term trend in CO2
concentrations from AIM 17 to AIM 8 is not seen in the dust flux,
while it is clearly present in Antarctic temperatures. Again, the
steep decrease in CO2 around 70,000 yr BP has no equivalent in the
dust or the d18O record, suggesting that additional processes must
be responsible for this atmospheric CO2 change. During the last
glacial/interglacial transition the majority of the CO2 increase
occured only after iron fluxes had already reached interglacial
levels (Wolff et al., 2006) and – based on the timing of the dust
decline – a maximum of 20–30 ppmv of the CO2 rise during the last
transition can be attributed to the increasing iron limitation.
Similarly, dust fluxes decreased to nearly interglacial levels during
the warm events in MIS3, however, this decrease was accompanied
by an increase in atmospheric CO2 of only 20 ppmv (Röthlisberger
et al., 2004). Thus, additional factors are required to explain the full
glacial/interglacial increase of atmospheric CO2 of 100 ppmv (note
that the glacial/interglacial CO2 flux from the ocean was even
higher than required to explain a measured 100 ppmv increase in
the atmospheric carbon pool only. For instance the concurrent
carbon uptake by the terrestrial biosphere carbon storage during
the last glacial termination required an additional flux out of the
ocean equivalent to a 30 ppmv change in atmospheric CO2). The
EPICA ice core records also provide more direct evidence of changes
both in biological nitrogen and sulfur production, based on the NHþ
and SO2
fluxes derived from continuous chemical ice core
analyses. These records show glacial/interglacial variations of only
20–30%, i.e. within the uncertainty of the accumulation rate
estimate used to calculate the deposition fluxes (Kaufmann et al.,
2010; Wolff et al., 2006). In any case, they do not support a strong
increase in biological productivity in the SO during glacial times.
Marine geological evidence for an increased export production
during the glacial is controversial. Observations based on the joint
interpretation of various productivity indicators in the SO point to
a reduced biological export production south of the recent Polar
Front (PF), while they indicate somewhat elevated export rates
north of it (Kohfeld et al., 2005; Martı́nez-Garcia et al., 2009).
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
In contrast, the latest inferences on the productivity in the Atlantic
sector of the SO based on small diatom resting spores, which
contribute little to the total sedimentation rate but are less affected
by dissolution, also point to increased diatom productivity south of
the PF (Abelmann et al., 2006). Finally, an increase in d15N values in
SO sediments during the glacial points to an increased utilization of
nitrate in the surface waters at that time (Francois et al., 1997). This
could be due to enhanced export production caused by iron
fertilization (if ocean circulation remained unchanged) or to
reduced upwelling of nutrients to the surface and, thus, a higher
utilization of a reduced macronutrient concentration in the surface
waters as suggested by Francois et al. (1997). In fact, clear indications of changes in upwelling of deep waters have been recently
found in SO opal records (Anderson et al., 2009). Those show
a twofold increase in opal flux to the ocean floor during MIS 3 and 4
in parallel to the CO2 increases in the Byrd ice core (Ahn and Brook,
2008) and a most pronounced sixfold increase during the last
glacial termination, which has been attributed to increased
upwelling of Si enriched deep water to the SO surface stimulating
opal production (Anderson et al., 2009). Despite these intervals
being times of increased opal production at the surface they are
times of increasing atmospheric CO2. Obviously, the degassing of
CO2 from the old carbon-enriched upwelling water was able to
increase atmospheric CO2 concentrations, putting a strong limit on
the effectivity of increased biological (opal) production in the SO on
reducing atmospheric CO2.
Models of the marine carbon cycle may further elucidate how
strong an iron fertilization effect might be. The applicability of
observations from iron fertilization experiments in models for
glacial conditions, however, is hampered because of unsolved
issues regarding the percentage of mineral dust derived iron that is
in a form available for biological uptake. Carbon cycle models
assuming that available iron is proportional to the soluble iron
concentration (which again is proportional to the aeolian iron
deposition) predict a drawdown of glacial atmospheric CO2 of
10–40 ppmv relative to preindustrial levels, depending on model
configuration (Archer et al., 2000a; Watson et al., 2000; Bopp et al.,
2003; Ridgwell, 2003; Köhler et al., 2005; Brovkin et al., 2007).
However, the bioavailability of soluble iron is not only dependent
on the iron input but also on the concentration of iron-binding
organic ligands that ensure that the iron remains in its bioavailable
form (Parekh et al., 2008). When taking reasonable ligand
concentrations in the SO into account in an ocean carbon cycle
model, the iron fertilization effect of glacial dust is strongly reduced
and accounts for only a 10 ppmv change in atmospheric CO2
concentrations when the dust flux is increased by a factor of 100
and only by 5–8 ppmv, when realistic dust flux changes are
prescribed for AIM 8, 12, 14 and 17 (Parekh et al., 2008).
In summary, the occurrence of a dust-induced iron fertilization
of SO primary production seems very likely, however, its effect on
export production and, following, on atmospheric CO2 changes is
limited. Based on state-of-the-art model exercises, this effect can
likely explain about 10 ppmv (and maximum 40 ppmv) of the
glacial/interglacial change in atmospheric CO2. For the millennial
CO2 variations during the glacials, a higher relative portion of the
20 ppmv changes may be explained by iron fertilization but again
the effect is most likely limited to less than 10 ppmv. Accordingly,
other processes connecting atmospheric CO2 and SO temperature
changes are required to explain the strong correlation described
3.2. Southern Ocean circulation changes
Theoretical background
The carbon fluxes in the SO reflect a fine balance between
carbon-enriched deep waters and surface waters subject to air/sea
exchange of CO2. Part of the deep water in the SO has initially been
subducted in the North Atlantic and transported to the SO as North
Atlantic Deep Water (NADW) spreading southward. The other part
is derived from localized Antarctic Bottom Water (AABW) formation close to the Antarctic continent (Fig. 5), which represents
globally the second most important deep water source. In addition,
Antarctic Intermediate Water (AAIW) formation subducts water to
intermediate depths north of the Antarctic PF after air/sea exchange
of CO2. Finally, the upwelling of Circumpolar Deep Water (CDW)
and modified NADW at the Antarctic divergence (Rintoul et al.,
2001) closes the cycle, and brings carbon-enriched deep waters
back to the surface. In addition to these advective transport
Fig. 5. Salinity cross section through the Atlantic sector of the Southern Ocean together with a schematic view of the Southern Ocean water masses and the SOMOC. Abbreviations
indicate the location of the PF (Polar Front), AABW (Antarctic Bottom Water), LCDW (Lower Circumpolar Deep Water), NADW (North Atlantic Deep Water), UCDW (Upper
Circumpolar Deep Water), AAIW (Antarctic Intermediate Water). Figure adapted from Ocean Data View.
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
processes, convective mixing of deeper and surface water also
contributes to the ventilation of the deep SO. For instance, AABW
formation today mainly occurs in cold shallow shelf waters (most
importantly the Weddell and Ross Seas) where heat loss and brine
rejection during sea ice formation are strong enough to destabilize
the water column overlying the continental shelf and to cause
convection. Subsequently, AABW export at the shelf edge provides
deep water spreading into other ocean basins below NADW or
recirculating back to the SO surface.
Very important for the net upwelling of deep waters along the
Antarctic Circumpolar Current (ACC) is the Southern Ocean
Meridional Overturning Circulation (furtheron referred to as
SOMOC), established by the northward transport of water at the SO
surface that requires upward movement of water from below.
These deep waters are enriched in carbon due to the remineralization of organic carbon and provide CO2 to the atmosphere when
they come back to the surface (Watson and Garabato, 2006). The
net northward transport of water at the surface is partly a consequence of the meridional Ekman transport induced by the strong
westerly winds prevailing in the SO (Fig. 6a). Accordingly, a change
in zonal wind speed is expected to also change the SOMOC and,
thus, the carbon flux to the surface with higher zonal wind speeds
increasing atmospheric CO2 (Toggweiler et al., 2006). In addition,
Toggweiler et al. (2006) suggested that a glacial northward shift of
the westerly wind belt could reduce the SOMOC and, thus, CO2 in
the atmosphere.
However, Ekman transport is only part of the story. As demonstrated by Karsten and Marshall (2002) and later taken up with
respect to atmospheric CO2 by Watson and Garabato (2006), the net
meridional transport of water (referred to as ‘‘residual circulation’’
by Karsten and Marshall (2002)) is the sum of the mean Ekman
transport JEkman and opposing mean eddy transport Jeddy over the
ACC (see also Fig. 5). Using oceanographic data, Karsten and
Marshall (2002) were able to quantify these two components
(Fig. 7), showing that the southward eddy transport partly offsets
the northward Ekman transport south of the PF and even overcompensates the Ekman component north of it. Karsten and
Marshall (2002) conclude that the residual northward and upward
flow averaged over the ACC south of the PF is about 13 Sv, while the
40 oS
180o W
E-P [mm/yr]
60 oS
15 o
30 o
80 oS
12 o
60 o
90 o W
12 o
12 o
zonal wind speed [m/s]
30 o
60 oS
15 o
90 o E
90 o W
r0 cw
Qnet þ r0 b SðE PÞ
180o W
where a and b are the thermal and haline expansion/contraction
coefficients of sea water, r0 the reference density, cw the specific
heat capacity of sea water, and S the salinity. As becomes clear from
equation (2), the buoyancy flux at the surface is determined by the
net heat flux Qnet (Fig. 6b) and the freshwater flux E-P (Fig. 6c) at the
surface (where the latter may also be influenced by the lateral sea
ice transport in the SO). The vertical downward diffusion of buoyancy is generally small compared to the surface buoyancy flux in
the SO (Karsten and Marshall, 2002; Watson and Garabato, 2006).
Thus, taking into account that for modern conditions the net heat
flux averaged over the ACC dominates the buoyancy flux at the
surface, equation (1) shows that for a given mixed layer meridional
buoyancy gradient vbm/vy, the residual circulation is controlled
primarily by the heat gain of the waters when travelling northward.
Accordingly, a decrease in the surface buoyancy flux in the past may
have lead to a reduction in the SOMOC and atmospheric CO2,
independent of zonal wind speed changes (Watson and Garabato,
2006) if vbm/vy remained constant. Vice versa, an increase in the
Bs ¼
90 o E
90 o E
40 oS
15 o
where hm is the mixed layer depth, bm the mixed layer buoyancy,
and k the vertical eddy diffusivity that removes buoyancy from the
mixed layer downwards. Bs is the surface buoyancy flux
80 oS
60 oS
¼ Bs k
ðz ¼ hm Þ
80 oS
Jres ðz ¼ hm Þ
60 o
30 o
Ekman transport alone amounts to 21 Sv. On the northern side of
the polar front, Karsten and Marshall (2002) derive a southward
residual flux of about 10 Sv. In fact, it is this convergence of cold
waters from the south and warmer waters from the north that
forms the PF and leads to subduction of AAIW. Accordingly,
a change in the eddy transport over time could also lead to a change
in the SOMOC and, thus, atmospheric CO2 (Watson and Garabato,
Using the buoyancy balance at the surface, Karsten and Marshall
(2002) showed that the residual circulation Jres at the base of the
mixed layer is given by:
90 o W
60 o
40 oS
180o W
net heat flux [W/m2]
Fig. 6. a) Zonal wind speed, b) evaporation minus precipitation (E-P) with evaporation estimated from the latent heat flux, c) surface net heat flux calculated from the sum of net
shortwave and net longwave radiation, sensible heat flux and latent heat flux (net downward heat fluxes taken as negative) in the Southern Ocean region, derived from NCEP/NCAR
reanalysis data (Kalnay et al., 1996). Note that the reanalysis data are model dependent. Furthermore, b) represents only the vertical freshwater exchange with the atmosphere and
no salt/freshwater fluxes due to sea ice formation and melting are included. Accordingly, b) and c) are given only to provide an overall picture of the zonal and meridional structure
but cannot be directly compared to the buoyancy fluxes in Fig. 7. Dashed white lines represent the current summer and winter sea ice edge, while the solid white line indicates the
average position of the modern PF. In the glacial, the position of the winter sea ice edge is close to the position of the modern PF.
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
in atmospheric CO2; such a reduction in wind speed is clearly not
supported by any observational or modeling evidence. Similarly, in
the model used by Toggweiler et al. (2006) an increase in wind
speed by 50% is required to increase CO2 by 35 ppmv. Accordingly,
any significant change in the SOMOC due to a change in the
strength of the westerlies appears unlikely.
Furthermore, a northward shift of the westerly wind belt as
hypothesized by Toggweiler et al. (2006) is also unlikely to be
responsible for a significant CO2 drawdown. Again, the intercomparison of GCM runs for glacial conditions showed no clear shift in
the location of the westerly wind belt (Menviel et al., 2008; Rojas
et al., 2008) and only very small latitudinal movements of the wind
belt in opposite directions for different models. Artificially moving
the westerly wind belt in an ocean carbon cycle model northward
by up to 10 latitude changed the atmospheric CO2 concentration
by only about 10 ppmv. Moreover, it increased the atmospheric CO2
concentration (Tschumi et al., 2008) in contrast to the hypothesis
by Toggweiler et al. (2006) which stipulated a decrease. The reason
for this is primarily an enlargement of the upwelling region in the
model when the wind belt is moved further to the north.
Bs [10-9m3/s2]
salt flux
heat flux
transport [Sv]
mean latitude
Fig. 7. Bottom: Oceanographic reconstruction (Karsten and Marshall, 2002) of the
residual circulation (northward circulation positive) in the Southern Ocean (purple
line) together with its components derived from Ekman (pink line) and eddy (blue
line) transport. Top: Total surface buoyancy flux Bs (buoyancy fluxes increasing the
mixed layer buoyancy (directed downward into the ocean) are plotted positive)
derived from the residual circulation (purple line) and its components derived from
the net heat flux (blue line) and the salt/freshwater flux (pink line). Fluxes were
derived along lines of equal stream function of the ACC, which vary in latitude from
one ocean basin to the other. Accordingly the two graphs represent a zonally averaged
cross section through the ACC. Figures adapted from Karsten and Marshall (2002).
mixed layer meridional buoyancy gradient in equation (1) could
also decrease the SOMOC if Bs was fixed.
Changes in the Southern Ocean westerlies
We first discuss the possibility that only the Ekman transport
has changed in the glacial due to wind speed changes or the location of the westerly wind belt in the SO. Direct observational
evidence of wind speed (or wind stress on the ACC) changes is
sparse. Using the differential change in the flux of glacial dust
derived from Patagonian dust sources (Delmonte et al., 2004) at the
two EPICA ice core sites, Fischer et al. (2007) were able to show that
there was little change in the transport time of dust from Patagonia
to the East Antarctic Ice Sheet. This also suggests that there was
little change in zonal wind speed. If at all, zonal wind speed was
slightly enhanced during colder periods. This implies that the
Ekman transport and, thus, upwelling and CO2 fluxes out of the
surface ocean remained largely unchanged or slightly increased
during cold periods. Recent comparisons of coupled general
circulation models (GCM) for glacial conditions (Menviel et al.,
2008; Rojas et al., 2008) also show very little change in the strength
of the zonal wind stress on the SO (with minor changes in opposite
directions in different models) in line with the ice core evidence. In
addition, coupled ocean carbon cycle model experiments (Menviel
et al., 2008; Tschumi et al., 2008) fail to explain significant CO2
changes by reasonable variations in zonal wind speed. For example,
in the model experiment by Tschumi et al. (2008), zonal wind
speeds had to be reduced by 75% to explain a 50 ppmv drawdown
Changes in buoyancy forcing
Thus, the SOMOC was unlikely to have been substantially
changed in the past due to changes in the northward Ekman
transport, however, the combined effect of Ekman and eddy
transport, which is controlled by buoyancy forcing, could have
potentially changed the SOMOC. Northward expansion of the sea
ice extent represents an efficient way to change the surface buoyancy flux. For modern conditions, summer sea ice is very limited
and winter sea ice is still far away from the PF, where significant
heat uptake occurs. During the glacial winter, sea ice encroached as
far as the modern PF in all regions around Antarctica, as reflected in
diatom records from the SO (Gersonde et al., 2005). Summer sea ice
was also largely expanded in the Weddell Sea region as far as the
modern winter sea ice edge but little summer sea ice change
occurred between 90 and 120 W (Gersonde et al., 2005). For all
other regions, the current marine geological data base does not
allow a reliable reconstruction of the summer sea ice coverage.
However latest ice core (Wolff et al., 2006; Fischer et al., 2007;
Röthlisberger et al., 2008) and atmospheric aerosol studies from
central and coastal Antarctica (Wagenbach et al., 1998; Rankin et al.,
2000; Rankin and Wolff, 2003; Hara et al., 2004;Yang et al., 2008)
indicate that sea salt aerosol in Antarctica is at least partly derived
from the sea ice surface and that sea salt records in ice cores can
provide a continuous proxy for regional sea ice coverage in the SO.
The ice core studies show that for the major part of the climate
records the sea salt aerosol flux, hence sea ice coverage, is linearly
related to Antarctic temperatures (Wolff et al., 2006; Fischer et al.,
2007; Röthlisberger et al., 2008). Only for very cold glacial conditions the sea ice proxy becomes increasingly insensitive due to the
long transport time of sea salt aerosol from the northernmost sea
ice regions and its strong depletion during transport. Here we
assume that Antarctic temperatures are to first order also representative for the northward expansion of sea ice in the SO, showing
that Antarctic temperature and, hence, sea ice changed largely in
parallel to atmospheric CO2 (Figs. 1–4). Based on this evidence, the
waxing and waning of sea ice coverage could have directly modulated the annual net heat gain of the SO surface waters. In return,
the reduced buoyancy gain would have changed the strength of the
SOMOC controlling the upwelling of carbon-enriched deep water.
This physical link between SO temperature, sea ice extent and
the SOMOC may explain the high correlation between CO2 and
Antarctic temperature both on orbital as well as millennial time
scales. The natural northern border of the northward expansion of
sea ice is the PF, where SST increases sharply northward and does
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
not allow for the presence of sea ice. When this northern sea ice
border was reached during the Last Glacial Maximum (LGM), the
SOMOC is expected to have ceased completely and the maximum
drawdown effect on atmospheric CO2 levels would have been
reached. Using a multibox model approach to quantify the effect of
a reduction in the SOMOC during the LGM, a drawdown of CO2 by
about 35 ppmv could be explained without iron fertilization
(Watson and Garabato, 2006). This effect is subsequently amplified
by carbonate compensation (Broecker and Peng, 1987).
A fly in the ointment of this buoyancy forcing hypothesis is the
unknown location of the PF during glacial times. Marine geological
evidence based on diatom species (Gersonde et al., 2003) points to
a general northward movement of SST isotherms across the SO but
is so far not able to resolve the frontal structure in the SO for glacial
conditions. Thus, the marine sediment data allow for a northward
movement of the PF. However, the PF, as defined by a strong
temperature gradient, could also have remained at a similar location as today while the overall temperature north and south of it
was generally lower, allowing for a latitudinal shift in diatom
assemblages. Circumstantial evidence for an unchanged location
may be derived from the fact that the modern location of the PF has
remained the region of upwelling of Si enriched deep waters during
the last deglaciation and for intervals during MIS 3 as evidenced by
an increased opal production (Anderson et al., 2009). Physical
oceanographic reasoning requires that the location of the PF is
geographically pinned by the bottom form stress in regions with
strong bottom topography and the PF cannot migrate northward
during glacial times in these regions (Moore et al., 2000; Olbers
et al., 2004). This is the case in the Weddell Sea region as well as in
the region where the mid ocean ridges cross the ACC. In other
regions of the SO, however, a change in the PF may be possible,
which would reduce the sea ice effect on the SOMOC by an additional heat gain in regions were the PF had moved northward.
Another open question concerns the unknown change of the
meridional gradient of the mixed layer buoyancy vbm/vy for glacial
conditions. Hypothetically, even with a strongly reduced surface
buoyancy flux, residual circulation could be maintained if the
meridional mixed layer buoyancy gradient across the ACC (vbm/vy in
equation (1)) decreased in parallel to the surface buoyancy flux.
Watson and Garabato (2006) regard such a change in vbm/vy as rather
unlikely. In fact, the glacial SSTs point to a significant positive
temperature and, thus, density gradient over the ACC north of the sea
ice edge (Gersonde et al., 2005). In addition, the increase in sea ice
formation during glacial times is expected to have increased salinity
in the mixed layer in the sea ice formation zone, due to brine rejection
and to have reduced salinity further north due to melting of exported,
salt-depleted sea ice. Accordingly, this process would tend to
strengthen the meridional buoyancy gradient, which would enforce
the effect of the reduction in surface buoyancy flux on the residual
circulation in equation (1). Experiments using eddy-resolving
models (Hallberg and Gnanadesikan, 2006) for changed surface
buoyancy fluxes in the past could elucidate the response of the
meridional mixed layer buoyancy gradient and, hence, the SOMOC.
Southern Ocean stratification changes
An alternative way to decrease the ventilation of the deep SO
and, thus, to decrease atmospheric CO2 is to decrease the deep
vertical convective mixing. This requires no net change in the
SOMOC. Köhler et al. (2005) were able to reduce atmospheric CO2
by about 35 ppmv by shutting off vertical mixing in the SO during
the glacial in a multibox model. Again, this effect is subsequently
amplified by carbonate compensation (Broecker and Peng, 1987).
Such a reduction of vertical mixing has to be linked to a change in
a physical process in the SO. By simulating climate and biogeochemical changes during MIS3 in response to changes in NADW
formation, a clear bipolar seesaw response in Antarctic temperatures on the freshwater forcing in the North Atlantic and an
accompanying change in CO2 could be found in an ocean GCM
coupled to an atmospheric energy balance model (Schmittner and
Galbraith, 2008). A similar CO2 response on a NADW reduction
during the Younger Dryas was previously discussed by Marchal et al.
(1999). Both the time scale and the amplitude (25 ppmv) of the CO2
changes during MIS 3 are in line with the ice core observations
(Schmittner and Galbraith, 2008). They ascribe the CO2 change to
a reduced stratification of the SO, as a consequence of the decrease
in the NADW salt export into the SO, and an accompanying
expansion of AABW in the deep ocean. Both effects favor the
increased export of preformed nutrients to the deep ocean, thus
diminishing the biological pump and increasing atmospheric CO2.
Note that no iron fertilization was prescribed in their model
experiment and that the circulation changes in the SO could entirely
account for the millennial CO2 changes during MIS3 observed in the
ice cores. This could provide an alternative explanation for CO2
variations following Heinrich events and their strong connection
with SO temperature in MIS 3 by changes in the Atlantic Meridional
Overturning Circulation via the bipolar seesaw. However it cannot
explain the overall reduced CO2 level during the glacial, where the
strength of the Atlantic Meridional Overturning Circulation was
probably only slightly reduced (McManus et al., 2004; LynchStieglitz et al., 2007). Note also, that in the study by Schmittner and
Galbraith (2008), the CO2 background level was about 255 ppmv,
which is significantly higher than the CO2 concentrations found in
the ice core record during MIS 3 and indicates that their model
setup did not reflect full glacial boundary conditions.
Alternatively, the enhanced export of sea ice during cold climate
periods provides freshwater to the surface north of the perennial
sea ice, which stabilizes the water column and prevents communication of surface and deeper layers in the SO. This process is
mainly restricted to the upper ocean layers and does not directly
influence deep convection. Enhanced brine rejection in the
perennial sea ice zone may increase the mixing there, partly
compensating for the higher stratification further north. Surface
stratification in the sea ice export region, however, would also
decrease the nutrient supply to the surface water in that region and,
thus, increase the nutrient utilization and decrease the export of
preformed nutrients to the deeper ocean, reducing atmospheric
CO2 (Francois et al., 1997; Marinov et al., 2008).
A decrease in atmospheric CO2 due to SO processes may be
additionally strengthened by a reduction of the air/sea gas
exchange in sea ice covered areas. Such a hypothesis, linking
atmospheric CO2 to the Antarctic sea ice extent via variation in the
air/sea gas exchange, was put forward by Stephens and Keeling
(2000). This process alone cannot account for the majority of the
glacial/interglacial CO2 change since it requires an unrealistically
impermeable lid over the SO all the way to the PF to draw down CO2
by 70 ppmv (Stephens and Keeling, 2000). In view of the strong
seasonality of sea ice coverage over the SO during the glacial
(Gersonde et al., 2005), when summer sea ice did not reach the
upwelling regions south of the PF, and in view of the divergent
wind field in the SO that continuously opens leads within the sea
ice, such a lid seems impossible to maintain. However, a reduction
in air-sea gas exchange during winter may have strengthened other
processes, leading to a further reduction of CO2 in the atmosphere.
4. Other evidence for SO hydrography changes
Both the iron fertilization as well as the westerly wind belt
hypothesis have now been tested using climate and carbon cycle
models, showing limited effect of a dust-induced increase in
marine biological productivity and no change in zonal wind and the
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
connected SOMOC. The surface buoyancy forcing hypothesis and
the SO stratification hypothesis still lack a stringent model test, at
least partly because of the limitation of global ocean circulation
models in resolving eddy processes in the ocean. In the absence of
such a test, the role of the buoyancy forcing on SOMOC changes or
changes in deep mixing remains an open question. However, based
on other marine evidence, we can at least substantiate the consequence of such a forcing, i.e. changes in the strength of the deep and
surface water exchange and a connected higher carbon storage in
the deep ocean.
Probably the most direct evidence of reduced ventilation of the
deep SO is based on the very depleted carbon isotopes in benthic
foraminifera in the Atlantic sector of the SO (Hodell et al., 2003) and
the extremely high glacial salinities in this ocean basin, derived
from pore water reconstruction (Adkins et al., 2002). The former is
a consequence of continuing mineralization of organic material in
a poorly ventilated deep water mass in the SO. The d13C decline
observed in benthic foraminifera reflects the isotopic depletion in
dissolved inorganic carbon in the deep SO during the LGM (Hodell
et al., 2003). The size of this d13C depletion (1.5&) is quantitatively in line with carbon cycle modeling results for a shutdown of
deep mixing in the SO (Köhler et al., 2005). The salinity increase is
a consequence of the increased brine rejection in the SO, due to sea
ice formation in the glacial, in combination with a reduced recirculation of AABW to the surface. Interestingly, higher salinities of
bottom water formed by brine rejection in the absence of melting
caverns below ice shelves are also predicted by an ice shelf model
(Hellmer, 2004); a situation encountered during the LGM when the
ice sheet progressed all the way to the continental shelf break.
The continued production of AABW together with a reduced
recirculation in the SO and a shallower glacial NADW mass necessarily requires that AABW spread further into other ocean basins
during the glacial. This is clearly observed in synchronous d18O
records from planktonic and benthic foraminifers in the North
Atlantic all the way to the latitude of the Iberian peninsula during
the LGM (Shackleton et al., 2000). These show fast Dansgaard/
Oeschger changes at the surface but a slower Antarctic bipolar
seesaw temperature signal at the bottom. It is also reflected by the
distribution of paleonutrient proxies in the North Atlantic such as
benthic foraminiferal d13C (Duplessy et al., 1988) or benthic foraminiferal Cd/Ca (Marchitto and Broecker, 2006), which clearly
illustrate the northward expansion of AABW during the glacial,
filling the North Atlantic below a depth of about 3000 m. A similar
horizontal hydrographic front with poorly ventilated water below
2000 m could also be evidenced in the Indian Ocean (Kallel et al.,
1988). Note that another direct consequence of a reduced SOMOC
during the glacial (but not of a reduced deep mixing in the SO) is
a lower AAIW formation, which could be observed in marine
sediment records. In fact, an increase of AAIW formation has been
traced in the Southwest Pacific in parallel to the large AIM events
during glacial times (Pahnke and Zahn, 2005) and could be attributed to intermittent increases in the SOMOC.
Finally, the slow build up of a poorly ventilated deep ocean
water mass during the glaciation would lead to an accumulation of
carbon in that water mass. During the transition from a glacial
maximum to an interglacial, when the deep water mass is returning
to the surface, this carbon must be released to the atmosphere.
Several lines of evidence support this. Both marine radiocarbon as
well as marine d13C records indicate the transport of an old water
mass (depleted in 13C and 14C) to upper water bodies (Spero and
Lea, 2002; Marchitto et al., 2007). The decrease in d13C and D14C in
marine records over the last termination started at around
18,000 yr BP in line with the increase in atmospheric CO2 recorded
in the EDC ice core (Monnin et al., 2001). However, the maximum
190& decline in atmospheric D14C in the middle of the last glacial/
interglacial transition would have required a very large 14C depleted
deep ocean reservoir, which so far could not be quantitatively
accounted for (Broecker and Barker, 2007). The opal production in
the SO (Abelmann et al., 2006; Anderson et al., 2009) also shows
a pronounced maximum during the transition (not in the glacial
when dust-derived iron input was the largest), pointing to the
transport of deep water with high nutrient concentrations back to
the surface enhancing biological export production. The latter
indicates that the vertical transport of macro- and micronutrients to
the SO surface at that time had a much stronger influence on biological productivity in the SO than the input of iron from aeolian dust,
as is also suggested for modern conditions (Wagener et al., 2008).
5. Summary & conclusions
Already more than 20 years ago, the important role of high latitude oceans in controlling atmospheric CO2 was recognized (Sarmiento and Toggweiler, 1984; Siegenthaler and Wenk, 1984; Knox
and McElroy, 1985). However, only recent ice core data from the
EPICA and other ice cores, as well as improved model approaches,
have allowed us to appreciate the dominant role of the SO to its full
extent. Although none of the processes is able to explain the
100 ppmv glacial/interglacial CO2 change by itself, models and
paleoclimatic data point to a reduction in SO ventilation (either by
a reduced SOMOC or by decreased mixing) being a prime factor in
the control of carbon storage in the abyss and, therefore, atmospheric CO2. Model experiments show that it is most likely not
a change in the westerly wind belt and the accompanying northward
Ekman transport that alter the SOMOC. Theoretical considerations
point to a strong increase in the southward eddy return flow forced
by changes in the surface buoyancy flux being responsible for
a potential change in SOMOC. Such a buoyancy forcing of the SOMOC
would be in line with the idea that waxing and waning of sea ice (due
to temperature changes) controls the heat exchange at the SO
surface. Alternatively, reduced convective mixing during the glacial
may also lead to reduced transport of dissolved inorganic carbon and
macronutrients to the surface. However, a physical link with
temperature or sea ice extent to explain the high correlation of CO2
and Antarctic temperature seems not to be as straightforward as in
the case of SOMOC changes.
In addition, a change in biological productivity as a result of
dust-induced iron fertilization appears to be very likely, however,
its effect is probably not as large as initial estimates predicted. In
fact, some of the indications of increased productivity in the SO
may reflect the higher utilization of a reduced supply of nutrients to
the SO surface from below (Francois et al., 1997) and not a higher
productivity. We estimate the increase in export production to
explain 10–20 ppmv of the total glacial/interglacial CO2 increase
and 5–10 ppmv of the millennial CO2 changes during MIS 3. The
question of how much and where export production and nutrient
utilization changed in the SO has not been settled and requires
further marine geological evidence to resolve apparent contradictions of various productivity proxies as well as lab experiments on
the bioavailability of iron supplied by glacial dust.
A limited influence of iron fertilization on SO carbon storage also
has far reaching implications for potential geoengineering plans to
draw down anthropogenic CO2 by artificial iron fertilization of the
SO. In addition to the unknown ecological consequences of such
plans, a similar experiment was already performed by nature in the
last glacial with very limited effect, despite a 5–20 fold increase in
iron supply. Moreover, the much stronger ventilation of the modern
SO compared to the glacial (as suggested by ample marine evidence)
would be able to bring any additional carbon that was exported by
an enhanced biologically pump quickly back to the surface, damping
any potential iron fertilization effect in the longer run.
H. Fischer et al. / Quaternary Science Reviews 29 (2010) 193–205
Clearly, other factors – such as changes in the terrestrial
biosphere, global ocean temperature changes and, especially,
carbonate compensation – are responsible for part of the change in
atmospheric CO2 between glacials and interglacials but these
processes cannot explain the extraordinary correlation between
atmospheric CO2 and Antarctic temperatures, hence, climate
conditions in the SO. The hypothesis of a reduced SOMOC joins
together the available puzzle pieces of marine geological as well as
ice core evidence on changes in the carbon cycle both on an orbital
as well as a millennial time scale. Alternatively, changes in SO
mixing could also have taken place. A best guess estimate of these
SO ventilation changes on glacial CO2 is a reduction by about
40 ppmv (with a range between 20 and 50 ppmv). Together with
the solubility effect for colder surface waters, the iron fertilization
effect and the carbonate compensation feedback, this can explain
a glacial/interglacial CO2 change of about 80 ppmv. However,
a crucial piece of this puzzle is still missing, i.e. a stringent model
representation of the surface buoyancy forcing and the eddyinduced circulation in the SO for past climate conditions. On the
observational side, the position of the glacial PF represents a very
important open question with a strong impact on any hypothesis
that invoke changes in the SO hydrography as cause of a glacial
drawdown of CO2. Moreover, a changed hydrography of the SO has
also an impact on the strength in the zonal circulation of the SO.
Hence, detailed reconstruction of the SO temperature distribution
in the glacial SO, of the strength and structure of the ACC, together
with respective ACC model experiments for past climate conditions
will further advance our understanding on paleoceanographic and
carbon storage changes in the SO.
This work is a contribution to the European Project for Ice
Coring in Antarctica (EPICA), a joint European Science Foundation/
European Commission scientific programme, funded by the EU
(EPICA-MIS) and by national contributions from Belgium,
Denmark, France, Germany, Italy, the Netherlands, Norway,
Sweden, Switzerland and the United Kingdom. The main logistic
support was provided by IPEV and PNRA (at Dome C) and AWI (at
Dronning Maud Land). TFS acknowledges support by the Swiss
National Science Foundation and the Prince Albert II of Monaco
Foundation. This is EPICA publication no. 225.
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