CPClimate of the PastCPClim. Past1814-9332Copernicus PublicationsGöttingen, Germany10.5194/cp-12-1919-2016Impact of meltwater on high-latitude early Last Interglacial climateStoneEmma J.emma.j.stone@bristol.ac.ukCapronEmilieLuntDaniel J.https://orcid.org/0000-0003-3585-6928PayneAntony J.https://orcid.org/0000-0001-8825-8425SingarayerJoy S.ValdesPaul J.https://orcid.org/0000-0002-1902-3283WolffEric W.https://orcid.org/0000-0002-5914-8531BRIDGE, School of Geographical Sciences, University of Bristol,
Bristol, UKBritish Antarctic Survey, Cambridge, UKCentre for ice and Climate, Niels Bohr Institute, University of
Copenhagen, Copenhagen, DenmarkDepartment of Meteorology, University of Reading, Reading, UKDepartment of Earth Sciences, University of Cambridge, Cambridge, UKEmma J. Stone (emma.j.stone@bristol.ac.uk)29September20161291919193221January201610February201628July201618August2016This 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/12/1919/2016/cp-12-1919-2016.htmlThe full text article is available as a PDF file from https://cp.copernicus.org/articles/12/1919/2016/cp-12-1919-2016.pdf
Recent data compilations of the early Last Interglacial
period have indicated a bipolar temperature response at 130 ka, with
colder-than-present temperatures in the North Atlantic and
warmer-than-present temperatures in the Southern Ocean and over Antarctica.
However, climate model simulations of this period have been unable to
reproduce this response, when only orbital and greenhouse gas forcings are
considered in a climate model framework. Using a full-complexity general
circulation model we perform climate model simulations representative of 130 ka conditions which include a magnitude of freshwater forcing derived from
data at this time. We show that this meltwater from the remnant Northern
Hemisphere ice sheets during the glacial–interglacial transition produces a
modelled climate response similar to the observed colder-than-present
temperatures in the North Atlantic at 130 ka and also results in
warmer-than-present temperatures in the Southern Ocean via the bipolar seesaw
mechanism. Further simulations in which the West Antarctic Ice Sheet is also
removed lead to warming in East Antarctica and the Southern Ocean but do not
appreciably improve the model–data comparison. This integrated model–data
approach provides evidence that Northern Hemisphere freshwater forcing is an
important player in the evolution of early Last Interglacial climate.
Introduction
Understanding the climate feedback processes that occur in the high-latitude
regions is essential because they are particularly sensitive to changes in
radiative forcing and act as amplifiers of climate change
(Vaughan et al., 2013). Peak high-latitude temperatures were several degrees warmer during the Last
Interglacial (LIG; approximately 129–116 kyr ago, based on
eustatic sea level variations; Masson-Delmotte et al., 2013) (Clark and
Huybers, 2009; Masson-Delmotte et al., 2011; Otto-Bliesner et al., 2006;
Sime et al., 2009; Turney and Jones, 2010), and maximum global sea level was
6 to 9 m higher than today (Dutton et al., 2015; Dutton and Lambeck,
2012; Kopp et al., 2009). Thus, the LIG represents an ideal case study to
understand and test the climate mechanisms that operate under warm climates.
The LIG, however, should not be considered a direct analogue for future
climate due to the difference in primary forcing mechanisms of seasonal
astronomical changes versus greenhouse gas (GHG) changes to explain the
observed warmth.
Until recently, climate model simulations of the LIG were typically compared
with a data synthesis for surface temperature consisting of one single
snapshot representing the warmest temperature anomalies for the whole LIG
(Lunt et al., 2013; McKay et al., 2011; Otto-Bliesner et al., 2013). In
particular, the annual surface temperature data synthesis from
Turney and Jones (2010) illustrates the large-scale spatial
pattern in peak LIG warmth but does not provide a global temporal climatic
evolution due to the difficulty in obtaining robust and coherent LIG
chronologies (Govin et al., 2015). Such
a compilation of LIG maximum warmth, as in the approach of Turney and Jones,
neglects any potential asynchronous temperature changes between regions, while previous studies (Bauch et al., 2011; CLIMAP Project Members, 1984;
Govin et al., 2012; Ruddiman et al., 1980; Van Nieuwenhove et al., 2011;
Winsor et al., 2012), though limited to only a few records, have provided
evidence of hemispheric surface temperature asynchrony during the early LIG.
A new LIG compilation (Capron et al., 2014) of surface
temperature changes has been produced for the high-latitude oceans
(latitudes northward of 40∘ N and southward of 40∘ S)
and polar ice sheets. In contrast to previous LIG datasets, this new data
synthesis benefits from a coherent temporal framework between marine and ice
core records. It thus provides the first spatio-temporal description of the
climate between 135 and 110 ka. In particular, surface temperature anomalies
have been calculated for four time windows: 114–116, 119–121, 124–126, and
129–131 ka, referred to as the data-based 115, 120, 125, and 130 ka time
slices. These four time slices are associated with quantitative estimates of
temperature errors, including the error in the reconstructed sea surface
temperature (SST) and the propagation of dating uncertainties: the 2σ uncertainty on SST anomalies is 2.6 ∘C on average and
1.5 ∘C for Antarctic surface temperatures (see
Capron et al., 2014, for methodological details and
2σ uncertainty estimates for individual records). Note that
Antarctic annual surface air temperature reconstructions are estimated based
on the water isotopic records after correction for sea water isotopic
composition and moisture source correction using deuterium excess data
(Masson-Delmotte et al., 2011). Capron et al. (2014) consider an error of
1.5 ∘C associated with these reconstructions. It accounts for the
uncertainty associated with this method and also partially accounts for the
uncertainty associated with possible impacts of changes in seasonality of
precipitation on the reconstructions, which remain difficult to quantify in
ice core data (Masson-Delmotte et al., 2011).
The data-based 130 ka time slice indicates robust new insights into the
early LIG climate with asynchronous maximum summer temperature changes
relative to the present day between the two hemispheres where the Southern Ocean
and Antarctic (annual) records show early onset of warming compared with the
North Atlantic records (Fig. 1c, d, e).
(a) 65∘ N (black) and 65∘ S
(grey) summer insolation (Laskar et al., 2004). (b)
EDC (EPICA Dome C) ice core CO2 concentration (Schneider et al., 2013)
(black). (c) North Atlantic core ODP-980 summer-SST reconstruction
(Oppo et al., 2006) (green) and associated 2σ
uncertainty envelope (Capron et al., 2014) (light
green). (d) North Atlantic core CH69-K09 δ13C record
(Govin et al., 2012) and 2σ uncertainty envelope (this study; see
details on the methodology in Capron et al., 2014). (e) Southern
Ocean core MD02-2488 summer-SST reconstruction (Govin et al., 2012)
(pink) and 2σ uncertainty envelope (Capron et
al., 2014) (light pink). (f) EDC surface temperature reconstruction
(dark blue) and associated 1.5 ∘C uncertainty envelope
(Masson-Delmotte et al., 2011) (light blue). Note that Govin et al. (2015) reports in Table 5 of their paper that the Antarctic
reconstructed surface temperature (based on EDC δD) starts increasing at
135.6 ± 2.5 ka based on the use of the RAMPFIT software. (g)
Red Sea relative sea level (RSL) data (probability maximum, red) with 95 %
confidence interval (Grant et al., 2012) (orange). (h) Red Sea
rate of RSL change (probability maximum for the first-order time derivative,
in Sv, black) with 95 % confidence interval (Grant et al., 2012)
(grey). Note that ice and marine records from (a) to (f)
are shown on the AICC2012 ice core chronology (Bazin et al.,
2013; Capron et al., 2014; Veres et al., 2013), while the Red Sea records
(g, h) are displayed on their original age scale, which is
independent of the AICC2012 ice core chronology (see Grant et al., 2012, for details) The grey vertical line marks 130 ka. The grey band highlights
the 129–131 ka time interval that has been considered for the construction
of the 130 ka data-based time slice for surface temperature (see Capron et
al., 2014, for details on the methodology).
Comparison with snapshot climate model simulations selected as part of an
“ensemble of opportunity” (Lunt et al., 2013) and presented in the most
recent IPCC report (Masson-Delmotte et
al., 2013) shows that the majority of models predict warmer-than-present
conditions earlier than documented in the North Atlantic records (Fig. 2),
while the magnitude of the reconstructed early Southern Ocean and Antarctic
warming is not captured (Fig. 2). An ensemble of LIG transient simulations
with climate models of intermediate complexity or general circulation models
(GCMs) with low-resolution/accelerated forcing also shows that only
including orbital and GHG forcing results in peak Northern Hemisphere (NH)
warming occurring earlier than that shown in the marine data records
(Bakker et al., 2013). These results highlight not only the importance of
producing defined time slices rather than a unique snapshot representative
of the whole LIG but also that important missing processes in the models are
likely required to account for this temporal mismatch between data and model
temperature anomalies (Capron et al., 2014). For
example, previous GCM simulations did not consider freshwater forcing from
melting of the NH ice sheets prior and during the onset of the transition
from glacial to interglacial conditions at 130 ka (Lunt et al., 2013).
Accordingly, other work has invoked freshwater forcing from melting
ice sheets to account for a mismatch between model and data records in the
geological past (Smith and Gregory, 2009). Enhanced insolation
forcing in the NH during the penultimate deglaciation resulted in rapid
ice sheet retreat and an increase in freshwater input to the North Atlantic
and a suppression of the Atlantic Meridional Overturning Circulation (AMOC)
near the end of the deglaciation (Carlson, 2008). In addition,
marine sediment core evidence shows that North Atlantic Deep Water (NADW)
production was reduced compared with the present day but recovered to present-day values by 125 ka (Böhm et al., 2015; Lototskaya and Ganssen,
1999; Oppo et al., 1997).
Simulated summer (NH July, August, September; SH January,
February, March) (left and middle panels) SST and annual (right panel)
surface air temperature change relative to pre-industrial period, for GCM model
results previously published (Lunt et al., 2013) and their ensemble mean.
The simulations are compared with the new data-based 130 ka time slice
(Capron et al., 2014).
A 130 ka climate model simulation (Holden et al., 2010),
including freshwater forcing, shows warming over Antarctica with a
freshwater input of 1 Sv into the North Atlantic between 50 and
70∘ N but still underestimates the temperature anomaly
interpreted from East Antarctic ice cores. This mismatch between model and
data is reconciled if the West Antarctic Ice Sheet (WAIS) is removed in
the simulation of Holden et al. (2010). However, a freshwater flux of 1 Sv is unrealistic for this
time period when compared with rates of change in sea level (Grant et
al., 2012, Fig. 1h). Previous modelling studies (Loutre et al., 2014;
Ritz et al., 2011; Goelzer et al., 2016) using climate models of
intermediate complexity show a reduction in the strength of the AMOC as a
result of freshwater input into the North Atlantic. Although
Loutre et al. (2014) were able to model the delay in NH warmth
in the early LIG when freshwater forcing was included, there is still a
mismatch in timing and/or magnitude between their model temperature response
and the temperature reconstructions. Govin et al. (2012) considered the
melting of the Greenland ice sheet and its influence on surface temperatures
and NADW formation at 126 ka and showed a slowdown of the AMOC along with
reduced SSTs in the North Atlantic, but the timing of the cooling from the
new data synthesis of Capron et al. (2014) predates
conditions at 126 ka. Similar work by Bakker et al. (2012) and
Otto-Bliesner et al. (2006) showed that melting of the Greenland ice sheet
resulted in a reduction in the AMOC strength and cooling in the vicinity of
the Labrador Sea. Goelzer et al. (2016) showed that freshwater fluxes
from the decaying Laurentide ice sheet during Termination II resulted in a
decreased AMOC and associated increases in Southern Ocean temperatures,
whereas freshwater from the Antarctic ice sheet led to surface cooling in
the same region. It is worth noting that mechanisms other than freshwater
fluxes have been invoked which could cause millennial-scale variations in
climate through changes in AMOC behaviour. These include a salt oscillator in
the North Atlantic (Peltier and Vettoretti, 2014) and wind stresses over
the subpolar gyre caused by changes in the Laurentide ice-sheet geometry
(Zhang et al., 2014). Furthermore, high-latitude climates are influenced by
changes in the mode of atmospheric circulation (e.g. Kleppin et al., 2015).
However, our main focus here is on characterizing the role of freshwater
fluxes in contributing to the LIG model–data mismatches.
The North Atlantic region for freshwater input denoted by
the red box (50–70∘ N). Note that the freshwater amount is evenly
distributed within this region.
The recent studies (Capron et al., 2014; Govin et al., 2012; Marino et
al., 2015) based on proxy reconstructions of temperature and sea level
speculated that the input of freshwater into the North Atlantic could
explain the reconstructed NH versus Southern Hemisphere (SH) early LIG
temperature pattern, via a bipolar response. Although previous modelling
studies (e.g. Bakker et al., 2013; Holden et al., 2010; Loutre et al.,
2014; Sanchez-Goni et al., 2012) have looked at the impact of freshwater
forcing on early LIG climate, they did not link the response with the data
reconstructions in the high-latitude regions of the Northern and Southern
Hemispheres and did not attribute this to a bipolar seesaw mechanism. As
such, we perform the first rigorous model–data comparison approach to
examine the impact and sensitivity of freshwater forcing on the high-latitude climate of the early LIG to test whether the hypothesis of a
bipolar mechanism is feasible in the framework of a comprehensive fully
coupled climate model to explain the difference in peak warmth conditions
between hemispheres at 130 ka. We further perform an idealized simulation
with the WAIS removed to test whether this has any additional influence on
regional warming in our model framework, as recent work has indicated that
some of the warmth seen in Antarctic ice core records during the LIG could
partly be explained by a reduced West Antarctic Ice Sheet (Steig et al.,
2015).
Locations of 130 and 125 ka data-based time slice data from Capron
et al. (2014). The colours denote the groups of data used in Method 2 (Eq. 2)
to calculate the RMSE for each region.
Greenhouse gas concentrations, orbital and freshwater
forcing, and state of the ice sheets (GrIS: Greenland ice sheet; M: modern-day
ice sheet; N: no ice sheet and orography flattened) for the GCM simulations
at 130 ka.
Greenhouse gas concentration Orbital parameters CO2CH4N2OObliquityEccentricityPerihelion(ppm)(ppb)(ppb)(∘)(day of year)25751223924.25 0.0401121.8Freshwater forcing (Sv)0.00.10.20.30.40.51.0WAIS; GrIS stateM; MM; M1. M; MM; MM; MM; MM; M2. N; M
Simulated 130 ka SSTs and near-surface air temperature
anomalies compared with data for the high-latitude regions. The top two rows
are SSTs (annual or summer as labelled) and the bottom row is annual mean
near-surface air temperature. Left panel, (a): LIG peak warmth data
synthesis of Turney and Jones (2010) (dots) compared with 130 ka
annual temperature anomalies (GHG and orbital forcing only). Middle panel,
(b): the 130 ka data-based time slice (dots) compared with simulated
summer-SST anomalies for the NH (July, August, and September) and SH
(January, February, and March) (GHG and orbital forcing only). Right panel,
(c): the 130 ka data-based time slice (dots) compared with
summer-SST anomalies for the NH and SH (GHG, orbital forcing, and a constant
freshwater input of 0.2 Sv into the North Atlantic). Note the non-linear
temperature scale. Anomalies calculated relative to the modern period for both the
model and the data.
Root mean squared error (RMSE) for NH and SH SSTs and East Antarctic Ice Sheet
(EAIS) near-surface air temperature regions. RMSE is calculated according to Eq. (1) (RMSE1) and Eq. (2) (RMSE2, values in brackets). The model output is
compared with the 130 and 125 ka time slices from
Capron et al. (2014). Note that the simulations
without freshwater forcing included were previously described in Lunt et
al. (2013) and references therein. JAS: July–August–September; JFM: January–February–March; ANN: annual mean.
In order to reconcile the high-latitude mismatch between the data and model
output at the beginning of the LIG for both hemispheres, we perform snapshot
climate model simulations, representative of 130 ka conditions. The LIG
starts at 129 ka when using a definition based on the eustatic sea level
(Masson-Delmotte et al., 2013); however, considering dating uncertainties
associated with paleoclimatic records during this time interval (see Govin et
al., 2015, for a review) and the fact that defining the boundaries of
interglacial periods is not trivial (see discussion in the Past Interglacials
Working Group of PAGES, 2016), we consider our 130 ka simulations as
representative of the “early LIG”. We use the UK Met Office fully coupled
GCM, HadCM3, with an atmospheric horizontal grid spacing of 2.5∘
(latitude) by 3.75∘ (longitude) and an ocean horizontal grid spacing
of 1.25∘ by 1.25∘ (Gordon et al., 2000), which includes the
MOSES 2.1 land surface scheme where water and energy fluxes are calculated.
For comparison with data we take advantage of the 130 ka data-based time
slice produced by Capron et al. (2014). Compared with the pre-industrial
period (see Table 1), the astronomical forcing resulted in greater
seasonality, leading to pronounced high northern-latitude summer insolation
during the early part of the LIG (Fig. 1a). GHG concentrations were similar
to pre-industrial values based on records obtained from ice cores (Loulergue
et al., 2008; Lüthi et al., 2008; Schilt et al., 2010) (Fig. 1b). In
addition to prescribing these forcings we further vary the amounts of
freshwater input between 0 and 1 Sv (Table 1, see also Fig. 9) injected
uniformly between 50 and 70∘ N in the North Atlantic Ocean (Fig. 3)
in order to determine the sensitivity of the model to freshwater forcing
under an early LIG climate regime. Given the uncertainty around the actual
location of the freshwater flux, we prescribe an idealized hosing region. The
climate simulations are run for 200 model years with fixed pre-industrial
vegetation and ice-sheet distributions. According to the highly resolved
millennial-scale global sea level reconstruction based on Red Sea records
(Grant et al., 2012), the rate of sea level rise was about 15.2 m kyr-1
during the glacial–interglacial transition (Fig. 1g, h). This is equivalent
to a flux of approximately 0.17 Sv, an estimate in agreement with the
0.19 Sv calculated by Carlson (2008) based on coral records. As a
consequence we choose an NH freshwater input (assuming no contribution from
the melting of the Antarctic ice sheet at this time) of 0.2 Sv
(HadCM3_BRIS_130ka_0.2Sv) as our best-estimate scenario with which to
compare our model temperature output and the high-latitude data synthesis at
130 ka. We also perform a 130 ka simulation with a freshwater
forcing of 0.2 Sv and the WAIS removed and its bedrock after removal defined
to be 200 m above sea level (HadCM3_BRIS_130ka_0.2Sv_NOWAIS) and
replaced with a bare soil surface, more akin to what is observed in the Dry
Valleys today. A land surface type was chosen instead of ocean due to
instabilities in the ocean numerics in HadCM3 close to the South Pole.
However, Holden et al. (2010) show with the GENIE climate model that
replacing the WAIS with ocean rather than land results in only a slight
increase in the surface air temperatures over Antarctica. Given the
uncertainty in the location and rate of freshwater forcing associated with
the WAIS removal, we do not prescribe additional freshwater fluxes from the
WAIS. Finally, we also perform a 130 ka simulation forced with a freshwater
forcing of 0.2 Sv and the WAIS removed, but with WAIS replaced with shrubs
instead of bare soil, to test the response to uncertainty in the land-cover
type which would replace the ice sheet. We perform analysis on the last
50 model years of each simulation. To test the robustness of the results to
the 200-year simulation length, we extended the 130 ka simulation with
0.2 Sv of freshwater forcing for a further 200 model years (400 years in
total). In the Southern Ocean the rate of change in summer SST with time is
very small, and the difference between the 50-year climate mean JFM (January–February–March) anomaly
after 200 years compared with the 50-year climate mean after 400 years is
very small (not shown); the difference ranges between -0.5 and
0.5 ∘C for the majority of the region, which is well within the
uncertainty of the data synthesis from Capron et al. (2014) of
2.6 ∘C on average.
For the model–data comparison, two methods have been used to calculate the
root mean square error (RMSE) to determine the influence of clustering of the
data points on the RMSE calculation. Method 1 is based on comparing each
observation (xi) at 130 ka with its coincident grid cell model value
(yi) according to Eq. (1):
RMSE1=∑i=1Nxi-yi2N,
where N is the total number of observations. Method 2 takes into account
the effect of clustering of the observations when compared with model values.
The RMSE is calculated according to Eq. (2):
RMSE2=∑i=1G∑j=1nixij-yij/ni2G,
where G is the total number of groups of clustered data points, ni is
the number of observations in each group, xi is the observation, and
yi is the model value. Each group (chosen based on geographical
proximity) is shown in Fig. 4 according to a different colour for the data
compilation at 130 and 125 ka for the three geographical regions considered.
The absolute error is calculated between each observation and its coincident
model value then averaged over the group.
Simulated summer (NH July, August, September; SH January,
February, March) SST and annual (bottom panel) surface air temperature
anomalies at 125 ka compared with the Capron et al. (2014) 125 ka data-based time slice. Left panel: HadCM3; right panel:
NorESM. Table 2 shows a similar agreement with data for NorESM and HadCM3
(which has no cooling in the North Atlantic) in the NH when compared with
the 125 ka data-based time slice from Capron et al. (2014). However, NorESM
shows poor agreement with the data synthesis over Antarctica. Anomalies
calculated relative to the modern period for both the model and the data.
Simulated summer (July, August, and September) near-surface air temperature anomaly compared with pre-industrial values at 130 ka over
southern Europe and the North Atlantic region. The simulation is forced with
0.2 Sv of freshwater flux as well as changes to the GHGs and orbital forcing
(HadCM3_BRIS_130ka_0.2Sv).
Results and discussion
Figure 5a shows results from the 130 ka climate simulation with no
additional freshwater input compared with the Turney and Jones (2010) time
slice, the latter assuming synchronous temperature changes across the globe
during the LIG. Figure 5b shows a comparison with the high-latitude 130 ka
time slice from the Capron et al. (2014) synthesis. Note that Turney and
Jones (2010) interpret the records as annual temperature means, while Capron
et al. (2014) interpret the marine records as summer temperature means, as
proposed by the authors of the original papers, and the ice core records as
annual means. In the North Atlantic, any similarity between the model and the
Turney and Jones data is misleading as the LIG temperature maximum recorded
by their study generally occurred later than 130 ka; a similar compilation
restricted to data from 130 ka would be much colder than the data shown in
Fig. 5a. This behaviour is seen in the Capron et al. (2014) 130 ka data
synthesis, now interpreted as seasonal, showing a cooling in the North
Atlantic. The model simulation with only orbital and GHG forcing (Fig. 5b)
matches the 130 ka compilation of Capron et al. (2014) poorly, with too
high temperature anomalies in the NH (RMSE1= 5.9 ∘C) and
too low temperature anomalies in the SH (RMSE1= 2.4 ∘C).
The mismatches in temperature between data and model are much too large to be
resolved, even taking into account the uncertainties in the marine temperature
reconstructions. Regarding temperatures over Antarctica, near-surface annual
air temperature anomalies are several degrees cooler in the model compared
with the Capron et al. (2014) synthesis, even considering the uncertainty in
the temperature reconstructions (RMSE1= 1.7 ∘C).
Furthermore, Table 2 shows that the RMSE result for each region is similar
for both RMSE1 and RMSE2. Note that the Capron et al. (2014)
dataset cites uncertainties in the data of 2.6 ∘C on average for the
data, which we do not propagate into the RMSE values.
(a) Simulated 130 ka SST and near-surface air temperature
anomalies, with the WAIS removed and replaced with bare soil and a North
Atlantic freshwater input forcing of 0.2 Sv, compared with the
Capron et al. (2014) 130 ka time slice. (b)
Difference in SST and near-surface air temperature between (a) and the 130 ka simulation with 0.2 Sv freshwater forcing only (Fig. 4c). (c)
Difference in SST and near-surface air temperature when the WAIS is replaced
with shrub rather than bare soil (Fig. 8a). Left: summer-SST anomalies for
the NH (July, August, and September); middle: summer-SST anomalies in the SH
(January, February, and March); right: annual near-surface air temperature
anomalies over Antarctica. It has been shown that this warming over East
Antarctica is attributed to climatic effects rather than isostatic effects
arising from a reduction in the WAIS (Bradley et al., 2012).
Inclusion of a constant freshwater forcing of 0.2 Sv in the North Atlantic
in the model results in a decrease in the strength of the AMOC of more than
10 Sv and an associated change from warming in the North Atlantic to a
cooling compared with the present day (Fig. 5c). This leads to a considerable
improvement in RMSE1 from 5.9 ∘C to 3.3 ∘C
for the North Atlantic compared with Capron et al. (2014). A warming compared
with the present day is observed in the climate model during the summer months for
the Southern Ocean, similar to when no freshwater forcing is included, but is
more extensive in the vicinity of the WAIS, with SSTs up to 2 ∘C
warmer than present. However, there is a lack of temperature records from
ocean sediment cores to further validate the model simulation in this region.
The addition of freshwater into the North Atlantic results in a bipolar
seesaw response (Stocker, 1998), with a redistribution of heat between the
hemispheres resulting from decreased northward heat transport through the
Atlantic (Crowley, 1992), although the response in the NH is stronger
compared with that simulated in the Southern Ocean. Here we use a snapshot
approach and, therefore, do not consider the timing of phasing between the
hemispheres in relation to the bipolar seesaw.
Other mechanisms have been suggested to explain the colder-than-present North
Atlantic at 130 ka. A study using the NorESM climate model (Langebroek and
Nisancioglu, 2014) (see Fig. 2 and Table 2) shows cooling in the North
Atlantic without the need to invoke freshwater input. Langebroek and
Nisancioglu (2014) attribute this to an expansion of the southeastern part of the subpolar gyre and an eastward
shift in the North Atlantic Current combined with a stronger AMOC. However,
marine sediment core evidence suggests that the AMOC was temporarily weaker
at this time (e.g. Böhm et al., 2015). Furthermore, this cooling persists
at 125 ka when the data show an overall warming compared with the present day
(see Fig. 6 and Table 2 for details).
Response of the North Atlantic summer SSTs and strength
of the overturning circulation to varying amounts of freshwater input
injected between 50 and 70∘ N (0, 0.1, 0.2, 0.3, 0.4, 0.5, and 1 Sv). Left axis: average 50-year model temperature anomaly
relative to the present day. The temperature is averaged over all model
grid boxes where data points are located. Right axis: 50-year average
maximum Atlantic Meridional Overturning Circulation strength at
30∘ N. The dashed horizontal line corresponds to the average
summer (July, August, September) temperature anomaly for all marine data
located at > 40∘ N for the 130 ka time slice (Capron
et al., 2014).
The Southern Ocean warming is coherent with the warmer-than-present
conditions suggested in ice core records from East Antarctica. There is a
small improvement in the RMSE over East Antarctica (RMSE1= 1.5 ∘C) when freshwater forcing is included compared to without
(RMSE1= 1.7 ∘C), although the model is still too cold by up
to 2 ∘C, similar to Holden et al. (2010). Recent work has suggested
that the Southern Hemisphere cooling arising from changes in the northward
heat transport in the Atlantic, such as we have here, can be communicated to
Antarctica by feedbacks associated with sea ice expansion; in particular, the
expanded sea ice reduces the winter warming effect of the Southern Ocean
(Pedro et al., 2016).
Although the new LIG data synthesis of Capron et al. (2014) does not extend
to continental records and to latitudes lower than 45∘ N, note that
forcing the model at 130 ka with a 0.2 Sv freshwater flux leads to
simulated surface air temperatures over Europe that are consistent with
existing datasets (e.g. Sanchez-Goni et al., 2012; see Fig. 7).
The contribution of the Greenland ice sheet to global LIG sea level rise has
recently been quantified (Born and Nisancioglu, 2012; Colville et al., 2011;
Helsen et al., 2013; NEEM community members, 2013; Quiquet et al., 2013;
Stone et al., 2013), with the IPCC Fifth Assessment Report stating a range
very likely between 1.4 and 4.3 m of equivalent sea level height
(Masson-Delmotte et al., 2013). Taking contributions from thermal expansion
and mountain glaciers into account and that global sea level was at least
6 m higher than today (Dutton et al., 2015), this implies that a contribution
is likely also required from the WAIS (noted specifically by Colville et al.,
2011) and/or other parts of the Antarctic ice sheet. Although studies have
suggested the possibility of an East Antarctic contribution (Bradley et al.,
2012; Fogwill et al., 2014; Pingree et al., 2011), this has yet to be
quantitatively supported by observational or modelling evidence. Future
research using ice-sheet models could investigate whether the warming of the
Southern Ocean via the bipolar seesaw mechanism leads to enhanced basal
melting of the WAIS and retreat of the grounding line (Joughin et al., 2012;
Timmermann and Hellmer, 2013; Goelzer et al., 2016; DeConto and Pollard,
2016) at the beginning of the LIG. Indeed, a recent study has suggested that
the water isotopic data from the Mount Moulton ice core drilled in West
Antarctica compared with water isotopic profiles from East Antarctic ice
cores are consistent with a collapse of the WAIS during the LIG (Steig et
al., 2015). This potential melting of the WAIS during the early LIG could
explain or partially explain the mismatch between the model simulations and
Southern Ocean/East Antarctic data time slices at 130 ka.
However, in the additional simulation (Fig. 8a) where we remove the WAIS and
include the freshwater forcing input of 0.2 Sv
(HadCM3_BRIS_130ka_0.2Sv_NOWAIS), the model–data match is not improved
(see Table 2) over East Antarctica and still underestimates the temperature
response by at least 1 ∘C (Fig. 8a), although there is an increase
in overall warming compared with the case when only freshwater forcing is
considered (Fig. 8b). This result supports, to some extent, the findings of
Holden et al. (2010), where the WAIS was removed and 1 Sv of freshwater was
added in the North Atlantic, leading to enhanced warming over East Antarctica, but, in our case, a more realistic amount of freshwater forcing based on data
is implemented.
There is some uncertainty as to the extent or type of vegetation which may
or may not have grown on an unglaciated West Antarctica during the LIG, and
the vegetation type replacing a previously glaciated surface can have a
significant effect on the magnitude of warming (Stone and Lunt, 2013).
Figure 8c shows that warming over Antarctica is sensitive to the land
surface type chosen to replace the WAIS with an increase in annual
temperature by up to 2 ∘C over Antarctica when covered with a
shrub surface type instead of bare soil. Another study using the CCSM3 model
(Otto-Bliesner et al., 2013), but without additional NH freshwater
forcing, found very limited improvement in the model response when the WAIS
was removed and replaced with ocean. It is possible that our simulations
with WAIS replaced by a land type are overestimating the warming.
Threshold behaviour of modelled AMOC strength in response to varying amounts of
freshwater forcing has been previously investigated, with models showing a
range from 0.1 to 0.5 Sv at which NADW formation can no longer be sustained
(Rahmstorf et al., 2005). As a result of this range in response of AMOC
collapse to freshwater input, we perform an analysis of the response of the
high-latitude regions to varying amounts of freshwater forcing in the North
Atlantic to test the sensitivity of the model under 130 ka forcing
conditions. This is similar to the study of Bakker et al. (2012), which looked
at the sensitivity of the AMOC to Greenland ice-sheet melt during the LIG
using a climate model of intermediate complexity. Figure 9 shows the model
summer North Atlantic temperature response (averaged over the locations for
which Capron et al., 2014, provide temperature records) for freshwater input, varying from 0 to 1 Sv compared with the average NH temperature anomaly from
the Capron et al. (2014) dataset (horizontal dashed line). In addition the
strength of the AMOC at 30∘ N for the varying amounts of freshwater
forcing is included. It is clear that HadCM3 shows a distinct threshold at
around 0.2 Sv under LIG boundary conditions, where freshwater input amounts
greater or equal to this lead to sufficient freshening in areas of NADW
formation and a reduction in the mixed layer depth in these regions. This
freshening results in reducing the overturning strength of the AMOC
considerably by more than 10 Sv. As a result the average temperature
response observed in the North Atlantic becomes cooler than in the present due to a
reduction in northward ocean heat transport. The weakening of the overturning
circulation occurs within 50 model years.
The implication of these simulations and the NH forcings depicted in Fig. 1
is that the freshwater forcing from the melting of the remnant ice sheets
provides a mechanism to warm the Antarctic and Southern Ocean during the
early LIG for a limited amount of time. From about 128 ka onwards, NH
surface temperature records and modelling studies (Capron et al., 2014) show
that surface warming relative to today also occurred in the NH. At 125 ka, when
the meltwater flux (Fig. 1) had likely returned to a low baseline, the match
between HadCM3 (orbital and GHG forcing only) and a similar compilation of
data (targeted at 125 ka, Fig. 6) is reasonable (Capron et al., 2014),
strengthening the case that a bipolar seesaw signal is required to reconcile
the evolution of temperature between 130 and 125 ka. This inter-hemispheric
bipolar seesaw pattern in temperature response during the penultimate
deglaciation first suggested by CLIMAP Project Members (1984) was also
highlighted in recent studies by Masson-Delmotte et al. (2010) and Marino et
al. (2015), while such a pattern has also been shown during Termination 1
(Shakun et al., 2012). Thus, this hemispheric asynchrony represents an
important feature of at least the last two glacial terminations. However, in
order to fully explore the temporal variations in temperature through the
LIG, fully transient simulations with time-evolving forcings would be
required.
Conclusions
Using new 130 ka snapshot GCM simulations and benefiting from the advent of a
new time-varying data-based representation of the climate evolution across
the LIG, we provide valuable modelling insights to explain the
inter-hemispheric asynchrony in temperature response during the early part of
the LIG. We show that the inclusion of freshwater input (determined from data
records) in the North Atlantic due to the melting of the remnant NH
ice sheets from the penultimate glaciation can explain the cold summer
temperature anomalies observed in the NH paleorecords and an extensive early
warming of the SH at 130 ka. Conversely, removing the WAIS in the simulations
does not improve the model–data comparison in East Antarctica or the Southern
Ocean. However, the lack of data coverage does not allow us to draw
conclusions regarding the configuration of the LIG WAIS. Our new results
highlight the need for additional paleoclimatic records (e.g. marine sediment
records in the vicinity of the WAIS) in order to better characterize both the
spatial and temporal high-latitude climatic patterns during the LIG. Possible
future work should include analysing ice-sheet model simulations of the WAIS
to test whether the ocean warming in these simulations is substantial enough
to increase basal melting of the ice sheet and grounding line retreat and to
account for the warming observed from ice core records in East Antarctica at
130 ka. This study shows the importance of studying the LIG not in isolation
but also in the context of the preceding glaciation. It further emphasizes
the importance of considering other forcings in addition to changes in
orbital and GHG forcings (which can lead to abrupt changes in the climate) in
future model simulations to improve the evaluation of their impact on climate
change, particularly in the high-latitude regions.
Data availability
The model data for this paper are available on request from the following
website: http://www.bridge.bris.ac.uk/resources/simulations. To access the data, click on Access Simulations. The data
synthesis can be accessed from the online Supplementary Material in Capron et
al. (2014).
Acknowledgements
This work was carried out with
funding from the UK-NERC consortium iGlass (NE/I009906/1) and is also a
contribution to the European Union's Seventh Framework programme
(FP7/2007–2013) under grant agreement 243908, “Past4Future. Climate change
– Learning from the past climate”. This is Past4Future contribution no. 85.
The climate model simulations were carried out using the computational
facilities of the Advanced Computing Research Centre, University of Bristol
– http://www.bris.ac.uk/acrc/. We thank the reviewers, whose
constructive comments improved the paper.
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