CPClimate of the PastCPClim. Past1814-9332Copernicus PublicationsGöttingen, Germany10.5194/cp-12-1721-2016Impact of ice sheet meltwater fluxes on the climate evolution at the onset
of the Last InterglacialGoelzerHeikoh.goelzer@uu.nlheiko.goelzer@vub.ac.behttps://orcid.org/0000-0002-5878-9599HuybrechtsPhilippehttps://orcid.org/0000-0003-1406-0525LoutreMarie-Francehttps://orcid.org/0000-0001-6944-4038FichefetThierryEarth System Sciences & Departement Geografie, Vrije Universiteit
Brussel, Brussels, BelgiumUniversité catholique de Louvain, Earth and Life Institute,
Georges Lemaître Centre for Earth and Climate Research (TECLIM),
Louvain-la-Neuve, Belgiumnow at: Institute for Marine and Atmospheric Research Utrecht, Utrecht
University, the NetherlandsHeiko Goelzer (h.goelzer@uu.nl, heiko.goelzer@vub.ac.be)25August20161281721173721August201517September201530June201612July2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://cp.copernicus.org/articles/12/1721/2016/cp-12-1721-2016.htmlThe full text article is available as a PDF file from https://cp.copernicus.org/articles/12/1721/2016/cp-12-1721-2016.pdf
Large climate perturbations occurred during the transition between the
penultimate glacial period and the Last Interglacial (Termination II), when
the ice sheets retreated from their glacial configuration. Here we
investigate the impact of ice sheet changes and associated freshwater fluxes
on the climate evolution at the onset of the Last Interglacial. The period
from 135 to 120 kyr BP is simulated with the Earth system model of
intermediate complexity LOVECLIM v.1.3 with prescribed evolution of the
Antarctic ice sheet, the Greenland ice sheet, and the other Northern
Hemisphere ice sheets. Variations in meltwater fluxes from the Northern
Hemisphere ice sheets lead to North Atlantic temperature changes and
modifications of the strength of the Atlantic meridional overturning
circulation. By means of the interhemispheric see-saw effect, variations in
the Atlantic meridional overturning circulation also give rise to
temperature changes in the Southern Hemisphere, which are additionally
modulated by the direct impact of Antarctic meltwater fluxes into the
Southern Ocean. Freshwater fluxes from the melting Antarctic ice sheet lead
to a millennial timescale oceanic cold event in the Southern Ocean with
expanded sea ice as evidenced in some ocean sediment cores, which may be
used to constrain the timing of ice sheet retreat.
Introduction
Understanding the climate and ice sheet evolution during past warm periods
in the history of the Earth may provide important insights for projections
of future climate and sea-level changes. The growing number of
palaeo-reconstructions for the Last Interglacial period (e.g. Govin et al.,
2012; Capron et al., 2014) in combination with improved model simulations of
this most recent warm period (e.g. Bakker et al., 2013; Lunt et al., 2013;
Langebroek and Nisancioglu, 2014; Loutre et al., 2014) makes it an
interesting target for studying the coupled climate–ice sheet system.
LOVECLIM model setup for the present study including prescribed
ice sheet boundary conditions from the Northern Hemisphere, Greenland, and
Antarctic ice sheets.
According to reconstructions, the Last Interglacial (LIG, from
∼ 130 to 115 kyr BP) was characterised by a global annual mean
surface temperature of up to 2 ∘C above the pre-industrial era (e.g.
Turney and Jones, 2010; Capron et al., 2014) and a sea-level high stand of
6–9 m above the present day (Kopp et al., 2009; Dutton and Lambeck, 2012).
As the penultimate glacial maximum was at least as severe as the Last
Glacial Maximum (LGM) in both hemispheres (EPICA community members, 2004;
Svendsen et al., 2004), this implies a large amplitude glacial–interglacial
transition in terms of temperature and ice sheet configuration. At the onset
of the LIG, a rapid warming of ∼ 10 ∘C from the
preceding cold state is recorded in deep Antarctic ice cores
(Masson-Delmotte et al., 2011) to have occurred between ∼ 135 and 130 kyr BP. Current ice core records from the Greenland ice sheet
(GrIS) do not extend long enough back in time to cover the entire
penultimate deglaciation and associated warming (NEEM community members,
2013), but a similar timing and magnitude of warming compared to the
Antarctic can be reconstructed for sea surface temperatures off the western
European margin (Sánchez Goñi et al., 2012). The warming is closely
related to an ice sheet retreat in both hemispheres. Despite large
uncertainties in reconstructions, the global sea-level stand at 135 kyr BP
of as low as -80 m (Grant et al., 2012) is indicative of the large amount of
freshwater that entered the ocean in the form of meltwater from the
retreating ice sheets during Termination II. Aside from determining the
amplitude of sea-level changes, which is the focus of many studies (e.g.
Robinson et al., 2011; Stone et al., 2013), the associated climate impacts
and possible feedbacks on the ice sheet evolution of this freshwater forcing
are an important element for a process understanding of the coupled
climate–ice sheet changes at that time.
A climatic mechanism that is thought to be directly related to changes in
the NH ice sheet freshwater fluxes (FWFs) is the interhemispheric see-saw
effect (Stocker, 1998) that links SH warming to a weakening of the Atlantic
meridional overturning circulation (AMOC). If the see-saw effect was active
during the onset of the LIG, NH ice sheet melting during Termination II
would have been the cause of a substantial AMOC weakening and NH cooling,
while reduced interhemispheric heat transport would have caused a gradual SH
warming (Stocker and Johnson, 2003). The see-saw mechanism was evoked to
explain part of the peak Antarctic warming during the LIG (e.g. Holden et
al., 2010; Marino et al., 2015), even though some Southern Ocean (SO)
warming was shown by Langebroek and Nisancioglu (2014) to be possible with
orbital forcing alone (without NH freshwater forcing). The see-saw mechanism
has been speculated to have caused increased Antarctic ice shelf melting and
West Antarctic ice sheet (WAIS) retreat (Duplessy et al., 2007). The retreat
of the WAIS, which is believed to have been grounded at the edge of the
continental shelf during the penultimate glaciation, generated a large
anomalous flux of freshwater into the SO. Such freshwater forcing could have
had a substantial influence on the SO configuration in terms of sea ice
extent and ocean circulation as shown in model experiments for the last
deglaciation (Menviel et al., 2011), for future global warming scenarios
(Swingedouw et al., 2008), and for the present day (Bintanja et al., 2013).
The impact of increased Antarctic FWFs is thought to consist of a surface
ocean freshening, stratification of the surface ocean, and cooling, in turn
promoting sea ice growth (e.g. Bintanja et al., 2013) and reduced Antarctic
Bottom Water (AABW) formation (Menviel et al., 2011). Recently, Golledge et
al. (2014) suggested that such a mechanism might also have provided a
feedback on Antarctic ice sheet (AIS) retreat for meltwater pulse 1A during
the last glacial–interglacial transition (Termination I), by promoting
warming of mid-depth ocean waters that provide additional heat for melting
ice shelves.
In the present work, we study the effect of evolving ice sheet boundary
conditions on the climate by simulating the climate evolution at the onset
and over the course of the LIG with an Earth system model of intermediate
complexity (EMIC). The model is forced with realistic ice sheet boundary
conditions from offline simulations of ice dynamic models of the AIS and
GrIS and reconstructions of the other NH ice sheets. With this study we
extend the work of Loutre et al. (2014) by additionally including dynamic
ice sheet changes of the GrIS and AIS and focusing on the effect of ice
sheet freshwater fluxes on the climate, particularly in the Southern
Hemisphere (SH). The model and experimental setup are described in Sects. 2
and 3, respectively, followed by results (Sects. 4, 5, and 6), their
discussion (Sect. 7), and conclusions (Sect. 8).
Evolution of reconstructed Northern Hemisphere ice sheets and
embedded modelled GrIS (top) and modelled AIS (bottom) used as boundary
conditions for the climate model.
Model description
We use the EMIC LOVECLIM version 1.3, which includes components representing
the atmosphere, the ocean and sea ice, the terrestrial biosphere, and the ice
sheets (see Fig. 1). The model has been utilised
in a large number of coupled climate–ice sheet studies (e.g. Driesschaert et
al., 2007; Swingedouw et al., 2008; Goelzer et al., 2011, 2012a; Loutre et
al., 2014) and is described in detail in Goosse et al. (2010).
In this study, the climate components are forced by time-evolving ice sheet
boundary conditions, which are calculated offline, i.e. uncoupled from the
climate evolution. Our modelling approach for the ice sheets consists of a
combination of reconstructed NH ice sheets (except the GrIS) based on
geomorphological data (Sect. 2.1) and of stand-alone
ice dynamic simulations of the GrIS and AIS (Sect. 2.2). In either case, the boundary conditions
provide time-evolving topography, ice sheet extent (albedo), and spatially
and temporally variable FWFs to the climate model. We conserve global
salinity and global volume in the ocean model to avoid numerical problems, a common practice for
simulations with large amounts of freshwater input.
Northern Hemisphere ice sheet forcing
We have little geomorphological evidence for Northern Hemisphere (NH) ice
sheet evolution during Termination II since it was mostly destroyed by the
re-advance leading to the LGM. Therefore, the reconstruction of NH ice sheet
evolution for the period of interest is made based on information from the
last deglaciation. The method has already been described in some detail in Loutre
et al. (2014). Nevertheless, we include a more thorough description here
(Appendix A). The resulting boundary conditions used to force the climate
model consist of a chronology of ice mask and surface elevation changes
(Fig. 2) and freshwater fluxes
(Fig. 3b) over the entire LIG period. Support for
the derived chronology of NH ice sheet evolution and their FWFs can be found
in records of ice-rafted detritus (IRD) from the subpolar North Atlantic
(Kandiano et al., 2004; Oppo et al., 2006). These records show variability of
similar signature during the deglaciation and in particular a last IRD peak
at ∼ 128 kyr BP preceding low IRD levels throughout the LIG.
Reconstructed ice volume (a, c) and freshwater forcing (b, d) from
the NH ice sheets (a, b) and from the GrIS and AIS (c, d). See Goelzer et al. (2012b) for definition of oceanic basins in (b).
Simulations of the Greenland and Antarctic ice sheets
For the present study, the climate components are partially forced by
results from stand-alone simulations of the GrIS and AIS, which have been
adapted from existing ice sheet model experiments (Huybrechts, 2002). The
configuration of both ice sheet models and the forcing interface follows the
description in Goosse et al. (2010) with the following exceptions. Forcing
for the ice sheet models is derived from scaling present-day observations of
temperature and precipitation with indices based on ice core records, as
often done for long-term palaeo-ice-sheet modelling (e.g. Huybrechts, 1990;
Letréguilly et al., 1991; Zweck and Huybrechts, 2005; Greve et al.,
2011). For the GrIS the forcing record was created following Fürst et al. (2015). We combine a synthesised Greenland δ18O record
derived from Antarctica Dome C using a bipolar seesaw model (Barker et al.,
2011) with the NEEM temperature reconstruction (NEEM community members,
2013) between 128.44 and 120 kyr BP. The Barker δ18O
record is converted to a spatially uniform temperature anomaly with a
constant temperature/isotope factor ΔT= 2.4 ∘C/‰ × (δ18O + 34.83) as in Huybrechts (2002).
Positive temperature anomalies of the NEEM record are scaled by a
factor 0.6 to fulfil constraints on maximal ice sheet retreat from Camp
Century and Dye3 ice core locations that are assumed to have been ice-covered during the LIG. This places the GrIS evolution in the range of
former model estimates during that period (e.g. Robinson et al., 2011; Born
and Nisancioglu, 2012; Stone et al., 2013). Such scaling is in line with
recent studies (e.g. van de Berg et al., 2013; Merz et al., 2014; Sjolte and Hoffmann, 2014; Steen-Larsen et al., 2014) that put in question the high
temperature of the central estimate reconstructed from the NEEM record.
Precipitation rates for ice sheet forcing vary percentage-wise as a function
of the δ18O record.
The AIS forcing is derived directly from the Antarctica Dome C record (EPICA
community members, 2004), following again procedures described by Huybrechts (2002). Here precipitation changes are assumed proportional to the saturated
water vapour pressure gradient relative to the temperature above the surface
inversion layer. Furthermore, both ice sheet models are forced by changes in
global sea-level stand based on the benthic deep-sea record of Lisiecki and
Raymo (2005) for the GrIS and on a more recent sea-level reconstruction
using Red Sea data (Grant et al., 2012) for the AIS, where the sea-level
changes are the dominant forcing. The chronology of the Red Sea record is
expected to be more accurate since new dating techniques are applied (Grant
et al., 2012). The impact of using another sea-level record for the GrIS
simulation over the LIG is small, because of the largely land-based
character of the ice sheet during that period. The AIS model is run at a
horizontal resolution of 20 km × 20 km instead of 10 km × 10 km (as in the
standard LOVECLIM configuration and for the GrIS model) due to computational
constraints for the relatively long duration of the LIG simulation.
To embed the dynamic GrIS simulation in the other NH boundary conditions,
the geometric evolution of the GrIS overrides prescribed changes where
Greenland ice is present. Therefore, the prescribed ice sheet evolution and
associated FWFs are not limited by the present-day configuration of the GrIS
as in Loutre et al. (2014). The ice sheet evolution is illustrated in
Fig. 2 for the modelled GrIS embedded in the NH
reconstruction (top) and for the modelled AIS (bottom). Ice volume evolution
for the NH ice sheets and the GrIS and AIS is given in
Fig. 3a and c,
respectively. The FWFs from the dynamic GrIS and AIS
(Fig. 3d) replace the background freshwater flux
from runoff over land calculated by the land model.
In our setup, the combined sea-level contributions from Antarctica and the
NH (including Greenland) fall within the 67 % confidence interval of
probabilistic sea-level reconstructions (Kopp et al., 2009) for the first
peak in sea-level contributions and the following period (∼ 124–120 kyr BP). For both hemispheres, the final 20 m rise in sea level at
the onset of the LIG is, however, steeper and occurs 1–2 kyr
earlier as compared to the reconstructions. When assuming a maximum
contribution from glaciers (0.42 ± 0.11) and an additional estimate for
thermal expansion of the ocean (0.4 ± 0.3) as given by Masson-Delmotte
et al. (2013), the assumed ice sheet evolution in our setup reproduces well
the average sea-level contribution between 125 and 120 kyr BP from the best
estimate of Kopp et al. (2009), but it does not represent the multi-peak
structure of global sea-level contribution during the LIG as suggested by
Kopp et al. (2009, 2013). More details about the ice sheet and sea-level
evolution can be found in a companion paper (Goelzer et al., 2016) that
specifically deals with the sea-level contribution of the ice sheets during
the LIG in a fully coupled model setup.
Initialisation
The goal of our initialisation technique is to prepare a climate model state
for the transient simulations starting at 135 kyr BP that exhibits a minimal
coupling drift. Both the GrIS and AIS models are integrated over the
preceding glacial cycles and the entire LIG in stand-alone mode. The climate
model is then initialised to a steady state with ice sheet boundary
conditions, greenhouse gas (GHG) forcing, and orbital parameters for the time
of coupling (135 kyr BP). In this way, when LOVECLIM is integrated forward
in time for transient experiments, the climate component is already relaxed
to the ice sheet boundary conditions and exhibits a minimal model drift in
unforced control experiments (not shown).
Matrix of all experiments and the respective ice sheet components
that evolve in time (yes) or are fixed (no). In the latter case, freshwater
fluxes (FWF, bold) are kept constant and topography and surface albedo are
fixed to the pre-industrial configuration.
Prescribed model forcings. (a) Average monthly insolation anomaly
relative to the pre-industrial era at 65∘ N in July (black) and
65∘ S in January (blue). (b) Combined radiative forcing
anomaly of prescribed greenhouse gas concentrations (CO2, CH4,
N2O) relative to the pre-industrial era. (c) Sea-level forcing for the ice
sheet components derived from either oceanic δ18O data
(Lisiecki and Raymo, 2005, red) scaled to a global sea-level contrast
between LGM and present day of 130 m or from a Red Sea relative
sea-level record (Grant et al., 2012, black).
Experimental setup
All simulations are forced by time-dependent changes in GHG concentrations
and insolation running from 135 until 120 kyr BP
(Fig. 4). The radiative forcing associated with
the reconstructed GHG levels (Petit et al., 1999; Peìpin et al., 2001;
Raynaud et al., 2005; Loulergue et al., 2008; Spahni et al., 2005) is below
pre-industrial values for most of this period and barely exceeds it at
∼ 128 kyr BP. The changes in the distribution of insolation
received by the Earth are dynamically computed from the changes in the
orbital configuration (Berger, 1978) and represent the governing NH forcing
during peak LIG conditions aside from evolving ice sheet boundary
conditions. In the following, we will compare results of the reference
experiment with all ice sheet boundary conditions evolving in time
(Reference) to experiments in which the ice sheet boundary conditions are
partially fixed to the pre-industrial configuration
(Table 1). To disentangle the effects of the
individual ice sheets, the experiments noGfwf (suppressed GrIS freshwater
fluxes) and noAGfwf (suppressed FWFs from both AIS and GrIS) are complemented
by two predecessor experiments with fixed AIS and GrIS and evolving NH
boundary conditions (noAG), as well as a climate experiment forced by
insolation and GHG changes only with all ice sheet boundary conditions fixed
(noIS). The last two experiments correspond to the allLR and IGonly
experiments from Loutre et al. (2014).
Evolution of global mean (a), northern hemispheric (b) and
southern (c) hemispheric mean surface temperature for experiments with different ice
sheet forcing included. Curves are smoothed with a running mean of 200 years
for better comparison.
Effect of GrIS and AIS on the temperature evolution at the onset of the
LIG
Including the forcing from the NH ice sheets in terms of configuration and
FWFs has been shown by Loutre et al. (2014) to be crucial to simulate the
onset of the LIG temperature increase and its amplitude variations more in
line with proxy records. This helps to partially overcome problems of EMICs
(and general circulation models) in simulating the strong temperature
contrasts inferred from proxy reconstructions (Bakker et al., 2013; Lunt et
al., 2013). The increased amplitude of temperature changes in our
simulations is due to albedo and elevation changes in addition to the larger
effect of the implied freshwater forcing from the NH ice sheets (Loutre et
al., 2014). Here the Loutre et al. (2014) experiments are complemented with
runs that additionally include changes in ice sheet configuration and FWFs
from the GrIS and AIS. We first discuss the effect of including these
additional ice sheet boundary conditions. A specific focus on the FWFs
follows in Sect. 5.
Comparison of modelled East Antarctic temperature evolution with
reconstructed temperature changes at deep ice core sites. Modelled
temperature anomalies are averaged over a region 72–90∘ S and 0–150∘ E. Ice core temperature
reconstructions for the sites EPICA Dronning Maud Land (EDML; 75∘00′ S, 00∘04′ E), Dome Fuji
(DF; 77∘19′ S, 39∘40′ E), Vostok (VK; 78∘28′ S, 106∘48′ E), and EPICA Dome C
(EDC;
75∘06′ S, 123∘21′ E) are from
Masson-Delmotte et al. (2011).
Freshwater forcing and oceanic response characteristics. NH
(a) and Antarctic ice sheet freshwater fluxes (f), strength of the AMOC (b), NH
sea ice area (c), SH sea ice area (d), and strength of AABW formation (e) for
the different experiments with and without freshwater forcing from
Greenland, Antarctic, and NH ice sheet melting. Curves are smoothed with a
running mean of 200 years for better comparison.
The temperature evolution (Fig. 5) before 127 kyr BP is in both hemispheres strongly influenced by the ice sheet boundary
conditions and in particular by the freshwater forcing from the ice sheets.
The experiments including FWFs from the NH ice sheets (Reference and noAG)
clearly show temperature variations on the multi-millennial timescale in
both hemispheres following variations in ice sheet freshwater input (cf.
Fig. 3). Differences in the temperature evolution
between noAG and the reference experiment are small in the NH, where the
additional freshwater flux from Greenland is small compared to the other
sources. In the SH, by contrast, a large perturbation arises around 130 kyr BP, when FWFs from the AIS peak. Global mean and hemispheric mean
temperatures are similar in all runs after ∼ 127 kyr BP, when
the ice sheets have largely reached their interglacial configuration and FWFs
are similar between the different experiments. An exception is the GrIS,
which is retreating until ∼ 120 kyr BP but accounts for only a
small FWF contribution. The similarity of the results in the runs after
∼ 127 kyr BP implies that the temporal memory of the response
to ice sheet changes in the system is limited to the multi-centennial timescale, at least for the surface climate. The location of largest
freshwater-induced temperature variations in the NH is the North Atlantic between
40 and 80∘ N. Here changes in the AMOC cause
a perturbation of the northward oceanic heat transport and temperature
changes, which are further amplified by sea ice–albedo and insulation
feedbacks. Greenland experiences maximum warming in the reference experiment
around 125 kyr BP of up to 2.7 ∘C in the annual mean compared to
the pre-industrial era over remaining ice-covered central Greenland. Here the
temperature evolution is largely similar to the experiment with GrIS changes
not accounted for (noAG), which exhibits a maximum warming of 2.4 ∘C (Loutre et al., 2014). However, the summer warming reaches up to
10 ∘C at the northern margin and even up to 14 ∘C over
southern margins over a then ice-free tundra (not shown). The strong warming
in the ice sheet periphery is due to a combination of elevation changes and
local albedo changes, confined to the immediate region of ice sheet lowering
and retreat. In the SH, the largest temperature perturbations linked to both
NH and SH freshwater fluxes occur in the SO. The largest warming over the
ice sheet itself is simulated over the WAIS (not shown) and is mainly a
consequence of the local elevation changes as the ice sheet retreats.
However, mainly due to the marine-based character of the WAIS, albedo
changes are much more limited compared to Greenland as the retreating ice
sheet surface is mostly replaced by sea ice. Modelled temperature changes
over the East Antarctic ice sheet (EAIS) have been compared to temperature
reconstructions for four ice core locations (Fig. 6). The reference experiment shows a more pronounced warming between 135
and 129.5 kyr BP compared to the experiments excluding Antarctic ice sheet
changes (noAG and noIS). While the modelled warming still appears to be
underestimated and delayed compared to the reconstructions, the reference
simulation clearly improves the representation of the EAIS temperature
evolution compared to experiments with fixed Antarctic boundary conditions.
Role of ice sheet meltwater fluxes
To study the role of the different freshwater contributions from the ice
sheets in more detail and evaluate their importance for the climate
evolution, we compare additional simulations where FWFs from the GrIS and AIS
are suppressed relative to the reference experiment
(Fig. 7). The ice sheet configuration (topography
and albedo) remains unchanged in these experiments. The effect of AIS FWFs
can therefore be evaluated as the difference between noGfwf and noAGfwf,
whereas the effect of GrIS FWFs becomes apparent from comparing the reference
simulation with noGfwf. The AIS FWFs (Fig. 7f)
lead to considerable changes in the SH but have very little impact on the
NH temperature evolution (cf. Fig. 5b).
Conversely, variations in the NH (Fig. 7a) and
GrIS freshwater forcing on millennial timescales imply temperature changes
in the SH on a background of general LIG warming.
Differences between the experiments in the AMOC evolution
(Fig. 7b) are largely explained by whether FWFs
from the NH ice sheets and the GrIS are included or not. Here AMOC strength
is calculated as the maximum value of the meridional overturning stream
function below the Ekman layer in the Atlantic Ocean between 45
and 65∘ N. The effect of the FWFs from the GrIS (cf. Reference and
noGfwf in Fig. 7b) is limited compared to the
large impact of the general NH ice sheet forcing and consists of an
additional weakening of the AMOC. It is most pronounced during periods of
AMOC recovery and after 130 kyr BP, when melting of the GrIS beyond its
present-day configuration sets in. Note that the simulated evolution of AMOC
strength in the reference experiment is in good agreement with palaeo-evidence based on δ13C data (Bauch et al., 2012) and in
particular with a recent reconstruction based on chemical water tracers
(Böhm et al., 2015). The timing of Heinrich Stadial 11 (∼ 132 kyr BP) and the variations in AMOC strength after that are well captured
by our reference simulation, which gives independent credibility to our NH
ice sheet reconstructions.
The evolution of NH sea ice area (Fig. 7c)
generally shows maxima at times of AMOC minima and vice versa and is closely
linked to NH surface temperature variations (cf.
Fig. 5b) by modifying the heat exchange between
ocean and atmosphere. The largest sea ice area between 135 and 130 kyr BP is
simulated in the reference experiment, which also exhibits the lowest AMOC
strength of all experiments.
The situation in the SH is more complex as surface temperature and sea ice
evolution are influenced by freshwater forcing from the AIS and also by the
FWFs in the NH. The AMOC variability gives rise to changes in the SH through
the so-called interhemispheric see-saw effect (Stocker, 1998). The SH begins
to warm as the NH cools due to modified oceanic heat transport across the
equator. Minima in SH temperature (cf. Fig. 5c)
and maxima in SH sea ice area (Fig. 7d) are
therefore associated with maxima in AMOC strength. The additional effect of
including GrIS freshwater forcing is consequently also felt in a warmer SH
with less sea ice formation. However, the influence of GrIS freshwater
fluxes and consequential AMOC variations on the SH temperature appears to be
mostly limited to the beginning of the experiment between ∼ 135 and 131 kyr BP. It could be speculated that this is related to the
larger extent of the SH sea ice in a colder climate, making the system more
sensitive due to an increased potential for sea ice–albedo and insulation
feedbacks. We also note that modelled periods of increased NH freshwater
fluxes, reduced AMOC strength, and higher SH temperatures are roughly in
phase with periods of steeper increase in GHG concentrations (cf.
Fig. 4b), in line with evidence from marine
sediment proxies that indicate that CO2 concentration rose most rapidly
when North Atlantic Deep Water shoaled (Ahn and Brook, 2008). Since GHGs and
NH freshwater fluxes are (independently) prescribed in our experiments, the
described in-phase relationship lends further credibility to our NH ice
sheet reconstruction.
The FWF from AIS melting (Fig. 7f) increases the
SO sea ice area (Fig. 7d) by freshening and
stratifying the upper ocean waters, which in turn leads to lower surface
temperatures. In our experiments, the increased freshwater flux from the
retreating AIS (cf. noGfwf versus noAGfwf) between 131 and 129 kyr BP is in
phase with a period of transient AMOC strengthening
(Fig. 7b), which leads to a combined effect of
surface cooling and sea ice expansion in the SO.
The formation of AABW is strongly controlled by salinity and sea ice area
(and therefore temperature) of the polar surface waters and hence directly by
Antarctic freshwater fluxes and indirectly by NH freshwater fluxes. Here the strength of AABW
formation is calculated as the minimum value of the global meridional
overturning stream function below the Ekman layer south of 60∘ S.
The AABW formation (Fig. 7e) is stronger for
saltier and colder surface conditions and therefore strongest in the case of
noAGfwf, where FWFs are suppressed from the AIS (saltier) and the GrIS
(colder). For a similar Antarctic freshwater forcing, the AABW formation is
stronger for a larger SH sea ice area. Including Antarctic FWFs leads to a
generally weaker AABW formation as surface waters become fresher (cf. noGfwf
versus noAGfwf). These relationships also imply that a stronger decrease in
AABW formation, associated with decreased CO2 uptake by the ocean can
be found for periods of steeper increase in prescribed radiative forcing.
Again, this appears to support consistency in timing between prescribed
radiative and NH ice sheet forcing in our modelling.
Evolution of annual mean sea surface temperature (a) and mid-depth
(485–700 m) ocean temperature (b) anomalies relative to the
pre-industrial era in close proximity to the AIS (south of 63∘ S).
(c) Meltwater-related changes in annual mean sea ice area at 129.5 kyr BP from differences
between experiments Reference and noAGfwf in per cent. The blue contour
outlines the observed ice-shelf edge and grounded ice margin of the
present-day AIS for illustration. All curves (a, b) are smoothed with a
running mean of 200 years for better comparison.
Time of maximum surface air temperature (MWT) in kyr BP for
experiments Reference (a), noGfwf (b), and noAGfwf (c) and difference in MWT
between experiments noGfwf and noAGfwf (d) in kyr, showing the shift of the
MWT when Antarctic freshwater fluxes are included.
Temperature evolution in the Southern Hemisphere
Millennial-scale sea-surface temperature variations in the SH induced by NH
freshwater fluxes are the strongest in the SO, where anomalies can be
amplified by sea ice–albedo and insulation feedbacks. This is also the
region that experiences the largest temperature change due to FWFs from the
AIS itself (not shown).
In order to study the effect of Antarctic FWFs in more detail, we also
analysed the oceanic temperature evolution south of 63∘ S
(Fig. 8). The effect of the AIS freshwater flux
in the reference experiment (compare noAGfwf with reference) becomes visible
in the sea surface temperature after 132 kyr BP
(Fig. 8a) as a cooling due to stratification and
sea ice expansion (Fig. 8c). At the same time,
the subsurface ocean warms (Fig. 8b) as heat is
trapped under the stratified surface waters and expanding sea ice area. When
the FWFs decline towards the end of the AIS retreat around 128 kyr BP, sea
ice retreats again and the heat is released to the atmosphere, where it
generates an overshoot in sea surface temperature compared to the experiment with constant
Antarctic freshwater fluxes (noAGfwf). The largest effect of this heat
buffering is found in winter in regions of strongest warming in the
Bellingshausen Sea and off the Gunnerus Ridge adjacent to Dronning Maud
Land. The maximum sea ice extent in the SH (Fig. 8c) occurs at the time of largest surface cooling at 129.5 kyr BP. This
freshwater-induced surface cooling at the onset of the LIG appears to be
superficial and relatively short-lived and of clearly different signature
compared to, for example, the Antarctic cold reversal during the last deglaciation.
The cooling event is indeed not recorded in our modelled temperature
evolution over central East Antarctica, in line with a lack of its signature
in Antarctic ice core records for that time period (Petit et al., 1999;
EPICA community members, 2004). A sea ice expansion during Termination II
together with an oceanic cold reversal around 129.5 kyr BP
(Fig. 8c) is, however, recorded in some deep-sea
sediment cores, where the composition of planktonic diatoms suggests
meltwater as the primary cause (Bianchi and Gersonde, 2002; Cortese and
Abelmann, 2002).
As a further consequence, the timing of maximum annual mean surface air
temperature (defined as MWT for maximum warmth timing; Bakker et al., 2013)
in the SO differs by several thousand years between experiments
(Fig. 9). Including Antarctic FWFs leads to an
earlier MWT (by up to 2 kyr) in large parts of the SO south of 60∘ S and in the central and eastern parts of the Atlantic sector of the SO up
to 40∘ S (Fig. 9d). Conversely, a later
MWT (by up to 3 kyr) is found in the Indian and Pacific sectors of the SO
north of 60∘ S when Antarctic FWFs are accounted for
(Fig. 9d). In the reference experiment
(Fig. 9a) and noGfwf
(Fig. 9b), the MWT lies relatively homogeneously
between -129 and -128 kyr for the entire SO south of 45∘ S and
coincides with the overshoot in sea surface temperature after the peak input of Antarctic FWFs.
The observed changes in the MWT in the SO due to the additional Antarctic
freshwater input can therefore in either way be understood as a shift
towards the time when heat from the mid-depth ocean buffer is released to
the surface.
Discussion
Despite remaining uncertainties in the timing of ice sheet retreat during
Termination II, we find several lines of evidence in support of our ice
sheet reconstructions and the associated climatic signatures. The NH ice
sheet reconstruction shows some similarity with the IRD signal recorded in
North Atlantic sediment cores (Kandiano et al., 2004; Oppo et al., 2006),
while the simulated evolution of the AMOC strength
(Fig. 7a) is in good agreement with a recent
reconstruction based on chemical water tracers (Böhm et al., 2015). The
combination of NH- and SH-sourced freshwater forcing variations produces a
stronger decrease in AABW formation, associated with decreased CO2
uptake by the ocean for periods of steeper increase in prescribed radiative
forcing, in line with evidence from marine sediment proxies that indicate
that CO2 concentration rose most rapidly when North Atlantic Deep Water
shoaled (Ahn and Brook, 2008). Reconstructing the NH ice sheet evolution
during Termination II with the same method but using the Grant et al. (2012)
sea-level record for comparison with Termination I has been shown to worsen
agreement of the modelled climate with proxy reconstructions (Loutre et al.,
2014).
Our modelling results furthermore suggest that the major AIS retreat from
its glacial configuration could be constrained by an oceanic cold event
recorded in several SO sediment cores around Antarctica (Bianchi and
Gersonde, 2002; Cortese and Abelmann, 2002). As a schematic sensitivity test
to uncertainties in the overall glacial AIS volume and retreat rate, we have
performed one more experiment identical to the reference experiment except
for Antarctic FWFs scaled to 50 % of their reference value. The resulting
magnitude of the SO cold event and overshoot is lower but exhibits the same
timing and spatial expression as in the reference case. The described
mechanisms and effects can therefore be considered robust to differences in
the magnitude of the freshwater flux, resulting from uncertainties in
glacial ice volume or AIS retreat rate. Notably, the improved representation
of the central East Antarctic temperature evolution in the model when
including Antarctic ice sheet changes (Fig. 6) is
largely independent of the chosen freshwater forcing. This implies that
changes in the geometry of the ice sheet and modified atmospheric
circulation patterns are the cause of the stronger simulated temperature
contrast.
The GrIS is generally assumed to have remained largely intact during the LIG
(e.g. Robinson et al., 2011; Colville et al., 2011; Stone et al., 2013; NEEM
community members, 2013) and indirect evidence of its freshwater
contribution may be difficult to find due to the low amplitude compared to
the other NH ice sheets. However, recent ice core reconstructions of the
temperature evolution at the NEEM ice core site (NEEM community members,
2013) point to a late retreat with a peak sea-level contribution close to
120 kyr BP. The GrIS can be assumed to lose mass approximately as long as
the temperature anomaly above the ice sheet remains above zero. Based on the
NEEM record, which has been used as forcing time series in our stand-alone
GrIS experiment, FWF from the GrIS peaks at ∼ 125 kyr BP but
remains elevated until around 120 kyr BP above the steady-state background
flux of an ice sheet in equilibrium with the climate. The additional FWF
from melting of the GrIS results in relatively low temperatures over
southeastern Greenland in response to a weakening of the AMOC (not shown). The
interaction between GrIS meltwater fluxes and oceanic circulation hence gives
rise to a negative feedback on ice sheet retreat. This aspect could play an
important role for the stability of the southern dome of the ice sheet and
should be examined further with fully coupled climate–ice sheet simulations.
In general, the NH freshwater forcing leads to variations in the strength of
the AMOC and North Atlantic cooling and additionally, through the bipolar
see-saw effect, to temperature changes in the SH. The only moment mid-depth
ocean temperatures close to AIS grounding lines are above pre-industrial
values in our experiments is during the oceanic cold reversal around 129.5 kyr BP, induced by anomalous FWFs from the retreating AIS. During this
period, SO mid-depth temperature anomalies relative to the pre-industrial era
reach up to 0.3 K, which could provide a positive but rather limited
feedback on ice sheet retreat, similar to what has been suggested by
Golledge et al. (2014) for meltwater pulse 1A during Termination I. However,
the oceanic warming recorded in our model is not strong and the duration of
the perturbation does not appear to be long enough for a sustained impact on
the retreat of the ice sheet. Furthermore, the peak in freshwater flux
appears when the ice sheet has already retreated considerably and WAIS
grounding lines are located mostly on the continental shelves, more
protected from the warm water build-up in the mid-depth ocean. A large-scale
marine ice sheet retreat of the likely less vulnerable EAIS sectors (Mengel
and Levermann, 2014) appears particularly unlikely, given the atmospheric
and oceanic forcing at the time apparent in our modelling results. However,
in-depth studies of these interactions require detailed coupled simulations
of the entire ocean–ice sheet system.
Despite aforementioned lines of evidence in support of the reconstructed NH
ice sheet evolution, a limitation to our modelling approach is the rescaling
of ice sheet retreat during Termination I, an attempt to address the
sparseness of geomorphological field evidence for Termination II. An
alternative approach would be to physically model all ice sheets together in
one framework (e.g. de Boer et al., 2013), although spatial and temporal
resolution of the models is a limiting factor in that specific case. A
rigorous modelling approach like the latter could also help to prevent
possible inconsistencies when combining ice sheet reconstructions from
different approaches. Nevertheless, any modelling approach will ultimately
be confronted with the same problem of scarce data for model validation
during that period. The exclusion of climate feedbacks on ice sheet
evolution of our present one-way coupled modelling approach is a general
limitation, which we have addressed in a separate study with a fully coupled
model setup (Goelzer et al., 2016).
Conclusions
We have presented a transient simulation of Termination II and the Last
Interglacial period with realistic ice sheet boundary conditions from
reconstructed NH ice sheets and detailed stand-alone simulations of the
Greenland and Antarctic ice sheets. Our results show that the temperature
evolution at the onset of the Last Interglacial was in both hemispheres
considerably influenced by meltwater fluxes from the retreating ice sheets.
While Antarctic freshwater fluxes lead to strong perturbations of the
Southern Ocean, NH freshwater fluxes have an influence on both NH
and SH temperature evolution through the oceanic see-saw effect. The
importance of additional freshwater input from the GrIS during Termination
II is small compared to the much larger fluxes from the other NH ice sheets
and becomes more important only later during the Last Interglacial, when it is the
only remaining ice sheet contributing freshwater fluxes to the North
Atlantic. In the SH, anomalous freshwater input from the AIS leads to an
episode of surface freshening, increased stratification, and sea ice cover
and consequently reduced ocean heat loss to the atmosphere, with temporary
heat build-up in the mid-depth ocean. We argue that the surface ocean
cooling associated with this event may be used to constrain an early
Antarctic retreat when matched with similar signatures evident in some
deep-sea sediment cores from the Southern Ocean.
Our transient simulations confirm results from earlier studies that stress
the importance of ice sheet boundary conditions for the climate evolution at
the onset of the LIG. However, most of the freshwater-induced changes remain
visible for at most 1–2 kyr after cessation of the perturbations, indicative
of a relative short memory of the (surface) climate system. Additional
effects may arise from climate–ice sheet feedbacks not considered in the
present model configuration, which should be investigated in fully coupled
experiments.
Data availability
The LOVECLIM version 1.3 model code can be downloaded from http://www.elic.ucl.ac.be/modx/elic/index.php?id=289.
Reconstruction of NH ice sheet forcing
A direct reconstruction of NH ice sheet evolution during Termination II
based on geomorphological data is not possible, due to the scarcity of field
evidence that was mostly destroyed by the re-advancing ice sheets during the
last glacial period. Therefore, a reconstruction of Termination II is made
by remapping the much better constrained ice sheet retreat of Termination I.
Ice extent during Termination I
The evolution of the NH ice extent since the LGM was estimated based on
published sources (Table A1) dating back to the time of the NH ice sheet
studies of Zweck and Huybrechts (2003, 2005). For the large Laurentide and
Eurasian ice sheets, inferred ice extents are relatively well determined from
geomorphological data and the reconstruction remains in good agreement with
most recent sources (e.g. Hughes et al., 2016). For smaller ice sheets such
as the European Alps, previous modelled ice extent was used (Zweck and
Huybrechts, 2005).
For ice sheets with multiple sources of data the isochrones were merged
using the most recent source when conflicts occurred (e.g. Dyke et al., 2002, instead of Dyke and Prest, 1987, for the Innuitian ice sheet,
Svendsen et al., 1999, instead of Andersen, 1981, for the LGM maximum of the
Eurasian ice sheet). The most recent source was then used as a mask of
maximum ice extent for most recent isochrones of all sources. The only
region which experienced an advance in ice extent using this technique was
the southern Cordilleran ice sheet according to the reconstruction of Clague
and James (2002).
The INTCAL98 timescale of Stuiver et al. (1998) was used to convert
radiocarbon dates to calendar years for the sources in Table A1. The retreat
of the ice sheets between the LGM and present day was prescribed at 200-year
resolution. Even for well-determined geomorphological observations,
uncertainties in dating and from the conversion of radiocarbon to calendar
years well exceed the 200-year temporal resolution used here. Fig. A1 in
Appendix A1
shows the deglaciation chronology reconstructed in this manner.
Sources of geomorphological data or modelling results used to
prescribe changes in Northern Hemisphere ice sheet extent for the retreat
during Termination I.
Ice sheetSourceIsochrone time period(kyr BP)LaurentideDyke and Prest (1987)18–present dayInnuitianDyke et al. (2002)18CordilleranClague and James (2002)20–present day (south)Dyke et al. (2002)Mayewski et al. (1981)18 (north)21–7 (interior)IcelandAndersen (1981)20–present dayEurasianAndersen (1981)20–present dayLandvik et al. (1998)15–12 (Barents Sea)Mangerud et al. (2002)18 (southern Barentsand Kara seas)Svendsen et al. (1999)18European AlpsZweck and Huybrechts (2003)21–present day(modelled ice extent)
Interpolated ice sheet extent during the last deglaciation for
the Northern Hemisphere ice sheets as a function of time (kyr BP). Hatched
regions indicate present-day ice.
Ice sheet elevations during Termination I
The NH ice sheets introduced significant changes to the surface topography
of the region. As LOVECLIM1.3 has only three atmospheric height levels, details
regarding topography are not strongly sensed. To include changes in surface
topography in the model, parabolic profile ice sheets are constructed using
the extents shown in Fig. A1, neglecting isostatic adjustment (i.e.
present-day surface elevation of the Earth's surface). The basal shear
stress for the parabolic profile reconstruction is chosen so that the
difference in ice volume between LGM and present day corresponds to 86 m of eustatic
sea-level change. With isostasy accounted for, a similar elevation would
result in an additional contribution of 24 m to a total equivalent eustatic
sea-level change of 110 m (cf. Zweck and Huybrechts, 2005). Using this
procedure the maximum elevation of the Laurentide ice sheet is 3000 m near
present-day Churchill in Hudson Bay, and the maximum elevation of the
Eurasian ice sheet is 2600 m 100 km west of present-day Helsinki.
Remapping Termination I to Termination II
Remapping of the retreat during Termination I to Termination II is done
using a benthic δ18O record (Lisiecki and Raymo, 2005), assumed
as an indicator of the global ice volume. In practice, the NH ice sheet
configuration for a given time (and δ18O value) during
Termination II is taken from a time during Termination I when the δ18O value had the same value. The LGM sea-level contribution of the NH
ice sheets relative to the present day of -110 m translates into a similar
magnitude for the penultimate glacial maximum (Lisiecki and Raymo, 2005).
The resulting NH ice volume evolution for Termination II is shown in
Fig. 3a. However, the method does not guarantee
that the sea-level contribution of the reconstructed NH ice sheets closely
follows the global ice volume curve. This is generally due to the mismatch
between global ice volume and NH ice sheet reconstruction during Termination
I, and in part related to the unconstrained contribution of other components
(AIS, thermal expansion). Due to the assumed analogy, different
configurations of the NH ice sheets (e.g. Obrochta et al., 2014) and
different relative timing of NH and SH deglaciation between last and
penultimate glaciation are not represented in these reconstructions. NH
freshwater fluxes were estimated from the same method by using derived
volume changes as input to a continental runoff-routing model (Goelzer et
al., 2012b) to identify the magnitude and location of meltwater fluxes to
the ocean.
Acknowledgements
We acknowledge support through the Belgian Federal Science Policy Office
within its Research Programme on Science for a Sustainable Development under
contract SD/CS/06A (iCLIPS). Computational resources were provided by
the supercomputing facilities of the Université catholique de Louvain
(CISM/UCL) and the Consortium des Equipements de Calcul Intensif en
Fédération Wallonie Bruxelles (CECI), funded by the Fond de la
Recherche Scientifique de Belgique (FRS-FNRS). We thank the two anonymous
reviewers and the editor for constructive comments and their follow-up of
the manuscript.
Edited by: G. Lohmann
Reviewed by: two anonymous referees
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