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GEOPHYSICAL RESEARCH LETTERS, VOL. 35, L17705, doi:10.1029/2008GL034410, 2008
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Antarctic ice-sheet melting provides negative feedbacks
on future climate warming
D. Swingedouw,1 T. Fichefet,1 P. Huybrechts,2 H. Goosse,1 E. Driesschaert,1
and M.-F. Loutre1
Received 21 April 2008; revised 9 July 2008; accepted 16 July 2008; published 10 September 2008.
[1] We show by using a three-dimensional climate model,
which includes a comprehensive representation of polar ice
sheets, that on centennial to millennial time scales Antarctic
Ice Sheet (AIS) can melt and moderate warming in the
Southern Hemisphere, by up to 10°C regionally, in a 4 CO2 scenario. This behaviour stems from the formation of a
cold halocline in the Southern Ocean, which limits sea-ice
cover retreat under global warming and increases surface
albedo, reducing local surface warming. Furthermore, we
show that AIS melting, by decreasing Antarctic Bottom
Water formation, restrains the weakening of the Atlantic
meridional overturning circulation, which is a new
illustration of the effect of the bi-polar oceanic seesaw.
Consequently, it appears that AIS melting strongly interacts
with climate and ocean circulation globally. It is therefore
necessary to account for this coupling in future climate and
sea-level rise scenarios. Citation: Swingedouw, D., T. Fichefet,
P. Huybrechts, H. Goosse, E. Driesschaert, and M.-F. Loutre
(2008), Antarctic ice-sheet melting provides negative feedbacks
on future climate warming, Geophys. Res. Lett., 35, L17705,
doi:10.1029/2008GL034410.
1. Introduction
[2] Current anthropogenic greenhouse gas emissions are
likely to affect climate for millennia, notably due to the large
thermal inertia of the oceans and the long memory of the ice
sheets [Meehl et al., 2007; Hasselmann et al., 2003].
Archives of the past suggest noticeable Antarctic Ice-Sheet
(AIS) melting contributions to sea-level changes during the
last deglaciation [Clark et al., 2002; Philippon et al., 2006]
and glaciation [Kanfoush et al., 2000; Rohling et al., 2004],
illustrating the possibility of massive freshwater input into
the Southern Ocean, which could have influenced the
climate [Weaver et al., 2003]. Recent observations report
an accelerated melting of the West Antarctic Ice Sheet
[Rignot and Thomas, 2002; Cook et al., 2005; Velicogna
and Wahr, 2006; Shepherd and Wingham, 2007]. This ice
melting may partly explain the freshening of the Ross Sea
observed during the past four decades [Jacobs et al., 2002].
Freshening also appears in the Antarctic Bottom Water
(AABW) [Rintoul, 2007] and could limit this deep-water
formation in the future and affect climate. While none of the
coupled climate models participating to the IPCC Fourth
1
Institut d’Astronomie et de Géophysique Georges Lemaı̂tre, Université Catholique de Louvain, Louvain-la-Neuve, Belgium.
2
Department of Geography, Vrije Universiteit Brussel, Brussels,
Belgium.
Copyright 2008 by the American Geophysical Union.
0094-8276/08/2008GL034410$05.00
Assessment Report [Meehl et al., 2007] take into account the
ice sheets melting for projections going up to the year 2100,
it is necessary to evaluate the potential effect of this melting
for longer projections.
[3] Potential irreversible changes both in the ice sheets
and ocean could actually lead to dangerous effects for
the environment, society and economy [Rahmstorf and
Ganopolski, 1999; Oppenheimer and Alley, 2004]. It is
therefore urgent to account correctly for ice-sheet-climate
interactions in climate projections. Ice-sheet retreat can
regionally enhance climate warming through changes in
topography and albedo. Furthermore, ice-sheet melting
releases freshwater into the ocean that can modify the
ocean circulation and sea ice cover [Weaver et al., 2003;
Fichefet et al., 2003; Swingedouw et al., 2006], and thus the
climate. The Greenland and Antarctic ice sheets are rather
different from each other since the total melting of the
former would represent around 7 m of sea-level rise, while
the latter would correspond to about 61 m [Huybrechts,
2002]. Moreover, contrary to the Greenland Ice Sheet
(GIS), the AIS has massive ice shelves, bordering the Ross
and Weddell Seas, where the bulk of AABW is formed. The
impact of GIS melting on climate and ocean circulation has
been evaluated in several studies [Fichefet et al., 2003;
Ridley et al., 2005; Swingedouw et al., 2006; Driesschaert
et al., 2007], contrary to its southern counterpart, the AIS.
In this study, we quantify the interactions of future AIS
melting with climate, using the climate model LOVECLIM.
2. Experimental Design
[4] To capture the respective roles of the AIS and GIS
impact under global warming, we performed 5 different
experiments (Table 1) using LOVECLIM, a three-dimensional
Earth system model of intermediate complexity (EMIC)
that includes representations of the polar ice sheets (see
methods section in the auxiliary materials).1 The first
experiment is a control simulation (CTRL) under preindustrial conditions that satisfactorily reproduces the climate
mean state [Driesschaert et al., 2007]. In the other simulations, the atmospheric CO2 concentration is increased by 1%
per year (compounded) until it reaches four times its initial
value, where it remains unchanged for 3000 years. These are
idealized experiments (called scenarios hereafter) designed to
capture the relevant ice-sheet-climate interactions in a warming world at the millennial timescale. The first scenario
(iAiG) has fully interactive ice sheets over Antarctica and
Greenland, while in the second one (fAfG), climate compo1
Auxiliary materials are available in the HTML. doi:10.1029/
2008GL034410.
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Table 1. Description of the 3000-Year Numerical Experiments Performed With LOVECLIM
Name
Description
CTRL
Control simulation with a constant forcing corresponding
to pre-industrial conditions, notably with the CO2
concentration in the atmosphere set to 277.6 ppm.
Scenario simulation in which the CO2 concentration increases
from the pre-industrial level by 1% per year and is maintained
constant after 140 years of integration when it reaches a value
equal to four times the pre-industrial level (4 CO2 scenario).
The climate components experience constant Antarctic and Greenland
ice-sheet areas and elevations, fixed at their preindustrial estimate.
The potential melting of the ice sheets due to warming is however
calculated ‘‘off line’’, but the corresponding freshwater fluxes
are not released to the ocean.
Same as fAfG but with fully interactive Antarctic and
Greenland ice sheets. Freshwater fluxes associated
with melting are released to the ocean. Ice-sheet area
and elevation are free to evolve and to influence the climate.
Same as fAfG but with fully interactive Greenland ice sheet.
Same as fAfG but with fully interactive Antarctic ice sheet.
fAfG
iAiG
fAiG
iAfG
nents are forced with a fixed ice-sheet configuration. In this
experiment, we still force the ice sheets ‘‘off line’’ with the
simulated warming, but without the potential feedback of
melting on climate. The ice sheets in this experiment are
therefore only ‘‘one-way’’ coupled. Two complementary
experiments have been conducted to isolate the individual
role of the AIS and GIS. Experiment iAfG (fAiG) has
interactive (fixed) AIS and fixed (interactive) GIS.
Consequently, the weakening of the deep convection and
hence the reduction in vertical heat exchange in the ocean
enhance the sea-ice extent, which cools the climate
through the higher sea-ice albedo [Stouffer et al., 2007].
3. Results
[5] The AIS begins to loose mass after a few centuries
in iAfG and iAiG. This is in contrast with previous studies
[Meehl et al., 2007; Mikolajewicz et al., 2007] and is
related to a large warming over the AIS in this model,
which leads to a larger increase in ablation than accumulation for the grounded AIS (see Figure S1 and Text S1 in
the auxiliary material). The melting of the AIS reduces the
increase in surface air temperature by 10% (0.3°C) on a
global average after 500 years and beyond in iAfG and
iAiG compared to fAfG and fAiG (Figure 1a). The relative
cooling between iAiG and fAfG occurs mostly in the
southern high latitudes (Figure 1b) and reaches 10°C in
the Weddell Sea sector (Figure 1c). This is associated with a
smaller decrease in sea-ice cover in the Southern Ocean in
iAiG compared to fAfG (Figure 1d). A slightly larger
warming appears north of 60°N in iAiG compared to fAfG,
mostly after 2000 years. At that time, 70% of the GIS has
melted (Figure S2), which explains this larger warming north
of 60°N when GIS is interactive, and is due to a reduction in
elevation and albedo over Greenland [Driesschaert et al.,
2007]. In the Northern Hemisphere, the annual mean sea-ice
extent decreases approximately at the same rate in the
different scenarios and evolves from 15 1012 km2 to 6 1012 km2 after 3000 years. The annual mean sea-ice extent
in the Southern Hemisphere decreases from 10 1012 km2
to 3 1012 km2 in iAiG and to 0.9 1012 km2 in fAfG after
3000 years. Contrary to the melting of the GIS, the climatic
impact of AIS melting is therefore mainly due to interactions
with the ocean and sea ice. After 3000 years, there is an
additional freshwater input into the Southern Ocean of up to
0.14 Sv in iAiG as compared to fAfG. This freshwater
decreases the surface density of the Ross and Weddell
Seas leading to the formation of a shallow halocline.
Figure 1. Time series of the annual mean surface air
temperature (SAT in °C): (a) globally averaged from CTRL
(black), iAiG (red), fAfG (green), iAfG (blue) and fAiG
(purple dotted line) and (b) zonally averaged: difference
between iAiG and fAfG. A 10-year running mean has been
applied to all time series. (c) SAT difference between iAiG
and fAfG averaged over years 2900 to 3000 expressed in °C
and (d) same difference but for sea-ice concentration for
each grid (ratio between 0 and 1), which is an index of seaice cover.
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Figure 2. Time series of the annual mean value of (a) the
minimum of the oceanic global meridional overturning
streamfunction at 30°S (in Sv, 1 Sv = 106 m3/s), representing
the export of Antarctic and Circumpolar Deep Water
(AABW and CDW) at 30°S, and (b) the maximum of the
Atlantic meridional overturning streamfunction at 30°S,
representing the export of North Atlantic Deep Water
(NADW) at 30°S. CTRL is in black, iAiG in red, fAfG in
green, iAfG in blue and fAiG in purple dotted line. A 21-year
running mean has been applied to all time series.
[6] Furthermore, the freshwater input associated with AIS
melting influences the ocean circulation in the scenarios.
Without AIS melting, the annual mean AABW export at
30°S (which is an index of the strength of the AABW cell)
weakens during the first 300 years and then recovers (in
agreement with studies from Bi et al. [2001] and Bates et al.
[2005]), and is even enhanced compared to CTRL after 1000
years (Figure 2a). This is caused by changes in the sea-ice
freshwater forcing related to the retreat of the sea-ice cover
(Figure S3). Indeed, the net annual mean sea-ice melting in
the Weddell and Ross Seas is lower in fAfG compared to
CTRL. This increases the surface salinity and density, and
counteracts the density loss stemming from the temperature
increase, leading to an increase in AABW formation in these
seas in fAfG compared to CTRL after 3000 years.
[7] The AABW export is 35% smaller in iAiG than in
fAfG, due to a decrease in surface density around Antarctica
and a reduction in AABW formation, associated with AIS
melting. Interestingly, the AIS melting also affects the North
Atlantic Deep Water (NADW) export (which is an index of
the strength of the NADW cell). At 30°S, this export
diminishes in all the scenarios (Figure 2b), but recovers after
1000 years in iAfG contrary to fAfG, illustrating the stabilizing effect of AIS melting on the NADW cell weakening.
When GIS melting is accounted for, the NADW cell further
weakens. This melting notably leads to a peak difference of
3.3 Sv (23% of NADW export at 30°S in CTRL) in fAiG
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compared to fAfG after 2000 years. The AIS melting once
more reduces the NADW cell weakening by 1.2 Sv in iAiG
compared to fAiG. This stabilization effect of the AIS
melting on the NADW cell can be explained by the so-called
bi-polar ocean seesaw [Stocker et al., 1992; Seidov et al.,
2001; Brix and Gerdes, 2003], which emphasizes that a
reduction in AABW density allows the NADW to penetrate
deeper and further south in the Atlantic, enhancing the
associated cell (see Text S1).
[8] Another important impact of ice-sheet melting concerns the sea-level rise. Here, we evaluate how interactions
between climate and ice-sheet melting can feed back on this
melting and influence sea-level rise in the various scenarios
(Table S1). According to its relative warming effect, the GIS
melting yields a positive feedback: in line with earlier finding
using LOVECLIM [Driesschaert et al., 2007], the whole ice
sheet has melted in fAiG and iAiG after 3000 years, while
60% remains in fAfG and iAfG. This positive feedback is due
to the reduction in albedo and altitude of the ice sheet, which
accelerates the melting. On the contrary, according to its
relative cooling effect, the AIS melting produces a negative
feedback, quantified by the comparison of the Antarctic
contributions to global sea-level rise in iAiG (3.2 m) and in
fAfG (10.0 m, calculated but not released to the ocean) after
3000 years. Moreover, the AIS melting tends to increase the
oceanic heat content (Figure 3) and leads to a larger thermal
expansion in iAiG compared to fAfG. This effect increases
the sea-level rise by 1.4 m in iAiG compared to fAfG and
corresponds to a warming at depth, while the surface,
particularly in the Southern Ocean, experiences cooling. This
is due to the capping of the ocean surface by freshwater
coming from the AIS melting, which inhibits the vertical
mixing of heat in high latitudes and warms the ocean interior.
On the whole, after 3000 years, the sea-level rise is 13.8 m in
iAiG, or 0.8 m less than the 14.6 m calculated in fAfG,
illustrating the compensation, in terms of sea-level rise,
between the GIS positive feedback and the AIS negative
feedback.
4. Conclusions
[9] A number of factors should however be borne in
mind when interpreting our results. The model used is an
Figure 3. Latitude-depth distribution of the annually
averaged temperature difference (in °C), years 2900 to
3000, of iAiG minus fAfG in the global ocean. Blue (red)
shading indicates values where the water is colder (warmer)
in iAiG than in fAfG. The contour interval is 0.2°C.
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EMIC and has therefore a rather coarse resolution. This
could affect deep water formation and the interaction
between the ocean and the ice-shelves [Nicholls, 1997]
but this is presently unavoidable to simulate the long-term
evolution of climate. Nonetheless, LOVECLIM has reached
sufficient realism concerning ice-sheet-climate interactions
to correctly capture the underlying mechanisms we have
illustrated here. The present study should not be seen as a
forecast but gives insight on the potential feedbacks between
climate and ice sheets melting for a given warming scenario.
Regarding the ice-sheet model, some of the potentially fast
processes (basal lubrication from penetrating surface melt
water, ice-flow acceleration induced by ice-shelf disintegration) by which warming may contribute to the ice-sheet mass
loss are not fully represented [Alley et al., 2005] so that a
faster decay could potentially happen. Note that ice sheet
melting might also be more rapid if processes responsible for
the widespread glacier acceleration currently observed in
Antarctica [e.g., Rignot et al., 2008] were taken into account
in the model. We therefore argue that ongoing efforts in icesheet modelling should continue and that AIS models should
be incorporated interactively in current ocean-atmosphere
general circulation models for centennial and millennial
projections of the climate system.
[10] Acknowledgments. We thank Chris König-Beatty, Gilles
Ramstein and Susan Solomon for comments on an earlier version of the
manuscript. We gratefully acknowledge the constructive comments from
two anonymous reviewers. This work was supported by the Marie Curie
Research Training Network NICE from the EU FP6 programme and by the
ASTER project of the Belgian Federal Science Policy Office Programme on
Science for a Sustainable Development. The authors wish to acknowledge
use of the Ferret program for analysis and graphics in this paper and the help
of Patrick Brockmann for the use of this program.
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1. Methods: Description of the climate model LOVECLIM
LOVECLIM consists of five components representing the atmosphere (ECBilt), the
ocean and sea ice (CLIO), the terrestrial biosphere (VECODE), the oceanic carbon cycle (LOCH) and the Greenland and Antarctic ice sheets (AGISM). ECBilt is a quasigeostrophic atmospheric model with 3 levels and a T21 horizontal resolution (Opsteegh et
al. 1998), which contains a full hydrological cycle and explicitly computes synoptic variability associated with weather patterns. Cloud cover is prescribed according to presentday climatology, which is a limitation of the present study. CLIO is a primitive-equation,
free-surface ocean general circulation model coupled to a thermodynamic-dynamic seaice model (Goosse and Fichefet 1999). Its horizontal resolution is 3◦ × 3◦ , and there are
20 levels in the ocean. VECODE is a reduced-form model of vegetation dynamics and
of the terrestrial carbon cycle (Brovkin et al. 2002). It simulates the dynamics of two
plant functional types (trees and grassland) at the same resolution as that of ECBilt.
ECBilt-CLIO-VECODE has been utilized in a large number of climate studies (please
refer to
http://www.knmi.nl/onderzk/CKO/ecbilt-papers.html
for a full list of references). LOCH
is a comprehensive model of the oceanic carbon cycle (Mouchet and Francois 1996). It
takes into account both the solubility and biological pumps, and runs on the same grid
as the one of CLIO. This model was not activated in the present study, and we pre-
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scribe the evolution of the atmospheric CO2 concentration. Finally, AGISM is composed
of two three-dimensional thermomechanical models of the ice-sheet flow, coupled to a
visco-elastic bedrock model and a model of the mass balance at the ice-atmosphere and
ice-ocean interfaces (Huybrechts 2002). For both ice sheets, calculations are made on a 10
km × 10 km resolution grid with 31 sigma levels. Given the long time-scales investigated
here, the model is among the most complex climate models that can be applied to study
this type of questions at present. Note that the mask of the ice shelves is fixed under
present-day configuration, and the land-sea mask is not modified during the integration
for the ocean, but can change for the ice-sheet models.
The atmospheric variables needed as input for AGISM are surface temperature and precipitation. Because the details of the Greenland and Antarctica surface climates are not
well captured on the ECBilt coarse grid, these boundary conditions consist of present-day
observations as represented on the much finer AGISM grid onto which climate change
anomalies from ECBilt are superimposed (Driesschaert et al. 2007). Monthly temperature differences and annual precipitation ratios, computed against a reference climate
corresponding to the period 1970-2000 AD, are interpolated from the ECBilt grid onto
the AGISM grid and added to and multiplied by the observed surface temperatures and
precipitation rates, respectively. The oceanic heat flux at the base of Antarctic ice shelves
is also calculated in perturbation mode using the parameterization proposed by Beckmann
and Goosse (2003). After performing mass balance and ice dynamics computations, AGISM transmits the calculated changes in land fraction covered by ice and in orography to
ECBilt and VECODE. In addition, AGISM provides CLIO with the geographical distri-
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bution of the annual mean surface freshwater flux resulting from ice sheet runoff, iceberg
calving, runoff from ice-free land and basal ice melting from both below the grounded ice
sheet and its surrounding ice shelves. All of these sources of fresh water are added to the
surface layer of coastal oceanic grid boxes. The Greenland (Antarctic) ice-sheet module
was first integrated over the last two (four) glacial cycles up to 1500 AD with forcing
from ice core data to derive initial conditions for coupling with the other components of
LOVECLIM. The control experiment (CTRL) of 3000-year duration was then conducted
with LOVECLIM under forcing conditions corresponding to 1500 AD. The same initial
conditions are used for all the scenario simulations performed in this study.
The model version used here is LOVECLIM1.1. Three main improvements have been
incorporated in this version compared to LOVECLIM1.0 (Goosse et al. 2007). First, the
land-surface scheme has been modified (see http://www.astr.ucl.ac.be/ASTER/doc/E AR SDCS01A v2.pdf)
in order to take into account the impact of the changes in vegetation on the evaporation
(transpiration) and on the bucket depth (i.e. the maximum water that can be hold in
the soil). Second, the emissivity, which was the same for all the surface types in LOVECLIM1.0, is now different for land, ocean and sea ice. Third, in order to reduce the
artificial vertical diffusion in the ocean caused by numerical noise, the Coriolis term is
now treated in a fully implicit way in the equation of motion for the ocean, while a
semi-implicit scheme was used in LOVECLIM1.0.
2. Supplementary discussion 1: Mass balance of the Antarctic ice sheet
Under global warming conditions, the mass balance of the grounded AIS depends on the
rate of change in accumulation over the ice sheet and ice loss around its perimeter, through
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surface runoff and ice discharge across the grounding line into the surrounding ice shelves.
Few studies have analysed the long-term mass balance of the AIS in millennial projections.
Huybrechts and de Wolde (1999) find a negative sea-level rise contribution from the AIS
of -0.3 m after 1000 years in a 4×CO2 experiment similar to the one performed in the
present study, but with the forcing derived from a two-dimensional climate model. An
annual mean temperature rise over Antarctica of 5.5◦ C is simulated in the Huybrechts
and de Wolde (1999) study after 1000 years. In an 8×CO2 experiment, Huybrechts and
de Wolde (1999) simulate an 8.5◦ C warming over Antarctica and a positive sea-level
rise contribution of 0.8 m after 1000 years. More recently, Mikolajewicz et al. (2007),
using another ice-sheet model coupled to a state-of-the-art climate model, find negative
contribution in terms of sea-level rise for the AIS in different projections using emission
scenarios going from B1 up to A2. In their model the increase in accumulation over the
grounded AIS is always larger than the increase in ablation for the grounded AIS.
In the present study, the simulated warming over Antarctica after 3000 years is rather
large in the projections as compared to CTRL. The annual mean temperature rise over
Antarctica reaches values of 9◦ C in iAiG and 12◦ C in fAfG for a spatial average over
Antarctica. This is due to a large polar amplification in this model that leads to an
important warming over the AIS in fAfG (Supplementary Figure 1.b). This polar amplification put the model used here in the higher range of polar amplification as simulated
in climate models (Meehl et al. 2007, Masson-Delmotte et al. 2006). Nonetheless, since
this model is on the lower range for climate sensitivity (Meehl et al. 2007), the simulated
warming over the AIS is not unrealistic, given the several centuries necessary to reach
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such a warming, due to the thermal inertia of the Southern Ocean. This large warming
over the AIS leads to a rather rapid decay of the West AIS in fAfG and also a substantial retreat in some coastal parts of the East AIS, notably in the Ninnis-Mertz glacier
basins in Victoria Land (Supplementary Figure 4). The processes causing this decay are
both the development of a peripheral surface ablation zone, as in Greenland today, as
the demise of the surrounding ice shelves from both large increases in surface and bottom
melting. Consequently, the grounded AIS already looses a substantial fraction of its mass
after 1000 years in iAiG and fAfG (Supplementary Figure 2), due to a larger increase of
ablation (sum of surface runoff from grounded ice, basal melting below grounded ice, flux
across grounding line) over accumulation (Supplementary Figure 1.a). The AIS melting
corresponds here, after 1000 years, to a sea-level rise of 0.5 m in iAiG and 1.5 m in fAfG.
This result differs from Huybrechts and de Wolde (1999) for a 4×CO2 experiment, but is
coherent with the AIS response to a larger warming as found in the 8×CO2 experiment,
in which the warming over Antarctica in Huybrechts and de Wolde (1999) is similar to our
4×CO2 experiment. The differences in mass balance response of the AIS compared to the
Huybrechts and de Wolde (1999) and Mikolajewicz et al. (2007) are therefore due to the
different climate model used here that exhibits a large polar amplification and warming
over Antarctica.
3. Supplementary discussion 2: Issues concerning the bi-polar ocean seesaw
The effect of the bi-polar ocean seesaw has been illustrated in numerical simulations
(Stocker et al. 1992, Seidov et al. 2001) and it has been shown that changing surface
buoyancy forcing in key deep-water formation areas can disturb the balance between
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NADW and AABW cells on millennial time-scales. Thus a decrease in AABW formation
reduces the AABW cell and enhances the NADW cell. The exact oceanic mechanism that
yields this interaction however remains unclear. Furthermore, recent simulations (Stouffer
et al. 2007, Seidov et al. 2005) show that, on centennial time-scales, an additional 1 Sv
input of freshwater into the ocean, south of 60◦ S, has nearly no impact on the NADW cell.
This result questions the validity of the bi-polar ocean seesaw since AABW formation is
strongly reduced in those experiments. Two explanations arise to account for this issue:
(i) the Seidov et al. (2005) and Stouffer et al. (2007) experiments use transient simulations
and the bi-polar ocean seesaw effect applies on longer time-scales, due to adjustment in
the ocean interior that necessitates thousands of years; (ii) the experimental design of
the numerical simulations from Stocker et al. (1992) and Seidov et al. (2001) on the one
side, and Seidov et al. (2005) and Stouffer et al. (2007) on the other side, are different
since the first-named impose surface buoyancy forcing anomalies in some key regions of
the Southern Ocean, using an ocean-only model, while the last-named, using an oceanatmosphere coupled model, put freshwater anomalies in the ocean south of 60◦ S, which
can spread through the intense currents of the Southern Ocean. Furthermore, observations
over the last decades suggest that variability in the NADW circulation is hardly influenced
by AABW (Koltermann et al. 1999).
In the present study, we have shown that even with an experimental design where freshwater is released into the ocean, using a coupled climate model, the bi-polar ocean seesaw
applies in LOVECLIM. We propose to analyze in further depth the exact mechanisms of
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the bi-polar ocean seesaw in a future study, since it is not the focus of the present one, in
order to try piecing some of the puzzle together (see Swingedouw et al. 2008).
References
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for climate models, Ocean Modell., 5, 157–170.
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Experiments with the CLIMBER-2 model. Global Biogeochem. Cycles, 16, doi:10.1029/
2001GB001662.
Goosse, H., and T. Fichefet (1999), Importance of ice-ocean interactions for the global
ocean circulation: A model study, J. Geophys. Res., 104, 337–355.
Goosse, H., E. Driesschaert, T. Fichefet, and M.-F. Loutre (2007), Information on the
early Holocene climate constrains the summer sea ice projections for the 21st century.
In press in Climate of the Past Discussion.
Huybrechts, P., and J. de Wolde (1999), The dynamic response of the Greenland and
Antarctic ice sheets to multiple-century climatic warming. J. Climate, 12, 2169–2188.
Huybrechts, P. (2002), Sea-level changes at the LGM from ice-dynamic reconstructions of
the Greenland and Antarctic ice sheets during the glacial cycles. Quat. Sci. Rev., 21,
203–231.
Koltermann, K. P.,et al. (1999), Decadal changes in the thermohaline circulation of the
North Atlantic, Deep-Sea Res. II, 46, 109–138.
Masson-Delmotte, V., et al. (2006), Past and future polar amplification of climate change:
climate model intercomparisons and ice-core constraints, Clim. Dyn., 27, 437–440.
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Mouchet, A., and L. M. Francois (1996), Sensitivity of a global oceanic carbon cycle model
to the circulation and to the fate of organic matter: Preliminary results. Phys. Chem.
Earth, 21, 511–516.
Opsteegh, J. D., et al. (1998), ECBILT: A dynamic alternative to mixed boundary conditions in ocean models. Tellus A, 50, 348–367.
Seidov, D., R. J. Stouffer, and B. J. Haupt (2005), Is there a simple bi-polar ocean seesaw?
Global Planetary Change, 49, 19–27.
Swingedouw, D., T. Fichefet, H. Goosse, M.-F. Loutre (2008), Impact of transient freshwater releases in the Southern Ocean on the AMOC and climate. Clim. Dyn., submitted.
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!! Please write \lefthead{<AUTHOR NAME(s)>} in file !!: !! Please write \righthead{<(Shortened) Article Title>} in fileX!! -
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Table 1. Sea-Level Rise for the Different Experiments in Comparison With CTRL After 3000
Yearsa
a
iAiG
fAfG
fAiG
iAfG
(in m). Sea-level
Antarctica Greenland Thermal
Total
Expansion
3.2
8.0
2.6
13.8
10.0
3.4
1.2
14.6
9.8
7.9
1.5
19.2
3.2
3.6
2.3
9.1
rise is decomposed into the contribution from Antarctic and Greenland
ice sheets melting and thermal expansion. The figures in italic stand for the fact that they have
been calculated, but the associated melting has not been released to the ocean and has therefore
no impact on ocean circulation.
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