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Clim. Past, 9, 955–968, 2013
www.clim-past.net/9/955/2013/
doi:10.5194/cp-9-955-2013
© Author(s) 2013. CC Attribution 3.0 License.
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Model sensitivity to North Atlantic freshwater forcing at 8.2ka
C. Morrill1,2, A. N. LeGrande3, H. Renssen4, P. Bakker4, and B. L. Otto-Bliesner5
1Cooperative Institute for Research in Environmental Sciences, University of Colorado, Boulder, CO, USA
2National Oceanic and Atmospheric Administration’s National Climatic Data Center, Boulder, CO, USA
3NASA Goddard Institute for Space Studies and Center for Climate Systems Research, New York, NY, USA
4Department of Earth Sciences, VU University Amsterdam, the Netherlands
5Climate and Global Dynamics, National Center for Atmospheric Research, Boulder, CO, USA
Correspondence to: C. Morrill (carrie.morrill@colorado.edu)
Received: 31 July 2012 – Published in Clim. Past Discuss.: 21 August 2012
Revised: 12 March 2013 – Accepted: 13 March 2013 – Published: 10 April 2013
Abstract. We compared four simulations of the 8.2ka event
to assess climate model sensitivity and skill in responding to
North Atlantic freshwater perturbations. All of the simula-
tions used the same freshwater forcing, 2.5Sv for one year,
applied to either the Hudson Bay (northeastern Canada) or
Labrador Sea (between Canada’s Labrador coast and Green-
land). This freshwater pulse induced a decadal-mean slow-
down of 10–25 % in the Atlantic Meridional Overturning Cir-
culation (AMOC) of the models and caused a large-scale pat-
tern of climate anomalies that matched proxy evidence for
cooling in the Northern Hemisphere and a southward shift
of the Intertropical Convergence Zone. The multi-model en-
semble generated temperature anomalies that were just half
as large as those from quantitative proxy reconstructions,
however. Also, the duration of AMOC and climate anoma-
lies in three of the simulations was only several decades, sig-
nificantly shorter than the duration of ∼150yr in the paleo-
climate record. Possible reasons for these discrepancies in-
clude incorrect representation of the early Holocene climate
and ocean state in the North Atlantic and uncertainties in the
freshwater forcing estimates.
1 Introduction
The Atlantic Meridional Overturning Circulation (AMOC)
plays a key role in the climate system, particularly through
its control on heat transport and storage of carbon in the
deep ocean. Changes in the AMOC can have far-reaching
effects on the El Ni˜
no–Southern Oscillation (Timmermann
et al., 2005), Atlantic hurricane development (Zhang and
Delworth, 2006), tropical rainfall (Vellinga and Wood, 2002),
and marine ecosystems (Schmittner, 2005). Model simula-
tions of the 21st century with prescribed greenhouse gas
concentrations increasing according to the Intergovernmen-
tal Panel on Climate Change (IPCC) scenario SRESA1B
uniformly show a reduction in the strength of the AMOC
(Schmittner et al., 2005). This multi-model ensemble yields
a mean decrease of 25% by 2100, but there is a large range
in the individual model results that indicates substantial un-
certainties in the AMOC response to climate change.
Several previous model intercomparison projects were un-
dertaken to improve understanding of the large spread in
modeled AMOC. Schmittner et al. (2005) considered the
skill of nine coupled climate models in matching observa-
tions of modern hydrography. They found that the models
were more successful at reproducing temperature patterns
than either salinity patterns or pycnocline depth. Stouffer
et al. (2006) examined the response of both Earth system
models of intermediate complexity (EMICs) and coupled
atmosphere–ocean general circulation models (AOGCMs) to
North Atlantic freshwater forcings of 0.1 and 1.0Sv (Sver-
drup=106m3s−1) for 100yr. While there were some ro-
bust patterns among the models, important disagreements
existed in model sensitivity and in reversibility following
AMOC shutdown. Since these were idealized experiments,
no comparison to observations was possible. Otto-Bliesner
et al. (2007) compared AMOC in four Last Glacial Maxi-
mum simulations from the second phase of the Paleoclimate
Modelling Intercomparison Project (PMIP2). These models
gave very different glacial circulations and a comparison to
Published by Copernicus Publications on behalf of the European Geosciences Union.
956 C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka
Table 1. Participating models.
Model Atmospheric model Oceanic model Citations
CCSM3 CAM3: T42 (∼2.8×2.8◦), 26
levels POP: ∼1× ∼ 1◦;∼0.3× ∼ 0.3◦in
North Atlantic, 40 levels,
volume-conserving
Collins et al. (2006)
Otto-Bliesner et al. (2006)
Wagner et al. (2013)
GISS ModelE-R ModelE: M20 (4×5◦), 20 levels Russell: 4 ×5◦, 13 levels,
mass-conserving Schmidt et al. (2006)
Russell et al. (2000, 1995)
LeGrande et al. (2006)
LeGrande and Schmidt (2008)
LOVECLIM1.2 ECBilt2: T21 (5.625×5.625◦),
3 levels CLIO3: 3 ×3◦, 20 levels,
mass-conserving Goosse et al. (2010)
paleoclimate proxy evidence indicated serious mismatches
for several of the simulations.
For the third phase of PMIP, the 8.2 ka event has been
targeted for a new model intercomparison. Of past abrupt
changes in the AMOC, the 8.2ka event provides a par-
ticularly useful case study because its duration (∼150yr;
Thomas et al., 2007) and forcing are constrained by the
proxy record, making an achievable target for climate model
simulations (Schmidt and LeGrande, 2005). There are still
some uncertainties regarding the hypothesized forcing of
the event, including the volume of drainage from proglacial
Lake Agassiz-Ojibway (hereafter Lake Agassiz; Barber et
al., 1999) into the Hudson Bay (northeastern Canada) and the
possibility of multiple meltwater pulses from both the lake
and the collapsing Laurentide Ice Sheet (Teller et al., 2002;
Gregoire et al., 2012). Model sensitivity to some of these
uncertainties has been explored elsewhere (Renssen et al.,
2001; Wiersma et al., 2006; LeGrande and Schmidt, 2008;
Clarke et al., 2009; Wiersma and Jongma, 2010; Wagner et
al., 2013). The target of this intercomparison is to use a me-
dian value for the forcing of the 8.2ka event and compare
model sensitivity to North Atlantic surface buoyancy anoma-
lies that have precise dating and a duration short enough to
make simulations with state-of-the-art coupled climate mod-
els feasible (Schmidt and LeGrande, 2005; Thomas et al.,
2007; Kobashi et al., 2007).
2 Models and experiments
We compare 8.2ka experiments completed with three mod-
els: the Community Climate System Model version 3
(CCSM), the Goddard Institute for Space Studies (GISS)
ModelE-R and LOVECLIM version 1.2. CCSM and
ModelE-R are atmosphere–ocean general circulation models
(AOGCMs) coupled without flux adjustments. LOVECLIM
is an Earth system model of intermediate complexity with
its most significant simplifications applied to the atmosphere
component (Table 1). These simplifications include clouds
that are prescribed and vertical profiles of temperature and
specific humidity that are limited by three atmospheric lev-
els. LOVECLIM also employs a freshwater flux correction
between the atmosphere and ocean subcomponents that re-
moves excess precipitation from the Arctic and Atlantic and
adds it to the North Pacific (Goosse et al., 2010).
Of relevance to this study, the ocean models of ModelE-R
and LOVECLIM are mass-conserving, in which the addition
of freshwater causes a rise in the free surface of the ocean and
reduces salinity purely through dilution. The ocean model
component of CCSM uses the rigid-lid approximation, which
does not permit vertical motion at the top of the ocean and
parameterizes the addition of freshwater as a salt extrac-
tion while keeping the volume of the ocean constant. Yin
et al. (2009) discuss the differences between these two ap-
proaches and compare results from two versions of the GFDL
CM2.1 model using each formulation. For a large freshwater
forcing that is similar in magnitude to that used in 8.2ka ex-
periments, the rigid-lid version exaggerates the forcing and
there are significant regional biases in sea surface salinity
(SSS). Despite this, the AMOC behaves similarly in the two
versions and many fundamental aspects of the two simula-
tions are qualitatively similar.
Boundary conditions specified for the control simulations
are listed in Table 2. Early Holocene orbital forcing increased
the seasonality of insolation in the Northern Hemisphere
and decreased seasonality in the Southern Hemisphere rel-
ative to the present (Berger, 1978). Greenhouse gas con-
centrations for the early Holocene were nearly identical to
those for the recent pre-industrial period (Fl¨
uckiger et al.,
2002; Monnin et al., 2004). Two of the control simulations,
CCSMall and LOVECLIM, incorporated the surface albedo
and elevation effects of the remnant of the Laurentide Ice
Sheet that was present near Hudson Bay at 8.5ka, as re-
constructed by Peltier (2004). These same control simula-
tions also included a small (∼0.05Sv) background flux of
Laurentide meltwater (Licciardi et al., 1999). In CCSMall,
this freshwater flux was added to the modeled St. Lawrence
River (Canadian/US Great Lakes Basin) at its outflow, and
was spread as a virtual salinity flux along the coast near
the river’s mouth. In LOVECLIM, the freshwater was added
Clim. Past, 9, 955–968, 2013 www.clim-past.net/9/955/2013/
C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka 957
Table 2. Boundary conditions for control simulations.
Simulation Orbital parameters Greenhouse gas concentrations Ice sheet Background meltwater flux
CCSMog 8.5ka CO2=260 ppm
CH4=660ppb
N2O=260ppb
none none
CCSMall 8.5ka CO2=260 ppm
CH4=660ppb
N2O=260ppb
ICE-5G 0.05 Sv added to St. Lawrence River
ModelE-R 1880 AD CO2=285ppm
CH4=791ppb
N2O=275ppb
none none
LOVECLIM 8.5ka CO2=260ppm
CH4=660ppb
N2O=260ppb
ICE-5G 0.05 Sv added to Hudson Strait
as a volume to the upper layer of the ocean at the Hud-
son Strait. Since the ocean model in LOVECLIM has a free
surface, this effectively means that the surface height was
raised. The temperature of the added freshwater in LOVE-
CLIM was assigned the same temperature as the water in the
ocean cell to which it was added. Both of these control simu-
lations with background meltwater flux were integrated until
reaching a quasi-equilibrium, in which SSS of the North At-
lantic had stabilized. Global mean ocean salinity decreases
slowly throughout these control simulations due to the back-
ground meltwater flux, a trend that parallels observed fresh-
ening since the Last Glacial Maximum (Adkins et al., 2002).
A second CCSM control simulation (CCSMog; OG=orbital
and greenhouse gas only) without a Laurentide Ice Sheet and
background meltwater flux is included in this study for a
more direct comparison to ModelE-R results.
For the 8.2ka event experiments, a meltwater pulse
(MWP) of 2.5Sv for 1yr was added to each of the control
simulations to represent the previously-mentioned drainage
of Lake Agassiz. This freshwater volume was the best esti-
mate for the drainage event based on flood hydrograph sim-
ulations (Clarke et al., 2004). Following the one-year per-
turbation, the MWP ceased and the climate was allowed to
recover. In the models with a free-surface ocean, the MWP
was added as a volume to a limited number of grid cells.
In ModelE-R, freshwater was added to the approximately
20 grid boxes in the Hudson Bay and was assigned a tem-
perature of 0◦C. In LOVECLIM, freshwater was added to
the upper layers of the ocean at the Hudson Strait and was
assigned the same temperature as the water in the ocean cell
to which it was added. The virtual salinity flux in CCSM re-
quired a larger area for the MWP (50–65◦N, 35–70◦W).
The control simulation for ModelE-R displayed a number
of transient, quasi-stable states with either strong or weak
AMOC (LeGrande et al., 2006; LeGrande and Schmidt,
2008). For this study, we use an experiment begun from a
period of weak AMOC. The weak case was chosen because
it exhibited the longest response to the 2.5Sv ×1yr forcing,
and because it may more closely emulate the early Holocene
than periods with strong AMOC since it lacks deep convec-
tion in the Labrador Sea between Canada’s Labrador coast
and Greenland (see LeGrande and Schmidt, 2008 for fur-
ther detail). Since the weak case exhibits some high ampli-
tude decadal variability, we examined “decadal” results for
this model (i.e., the 10-yr mean of the MWP experiment less
the 30-yr mean of the closest control years) in order to more
clearly show the transient response to the MWP.
We calculated the Student’s ttest for the differences be-
tween control and MWP experiments for the two models with
annual output (i.e., LOVECLIM and CCSM). Since annual
output is no longer available from the ModelE-R experiments
and resources do not exist to re-run these simulations, we are
unable to make statements about the statistical significance
of the model’s response to the MWP, about the skill of the
model on the annual time-scale and about the relative am-
plitude of decadal and annual variability. The model simula-
tions in the present study are an “ensemble of opportunity”,
meaning that most were completed before this intercompari-
son was planned. One limitation of not coordinating experi-
ments is that perfect comparisons are sometimes impossible
to make. However, the main conclusions of this paper would
not change were annual data available from all simulations
and ensembles of opportunity are important for informing
future coordinated intercomparisons.
3 Response to freshwater forcing
3.1 AMOC
AMOC intensity is defined here as the maximum of the
Atlantic overturning streamfunction excluding the surface
(<500m) wind-driven overturning circulation. Mean values
for the control simulations range from 16 to 20Sv (Fig. 1),
and interannual variability is small in the three simulations
www.clim-past.net/9/955/2013/ Clim. Past, 9, 955–968, 2013
958 C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka
Fig. 1. The Atlantic meridional overturning streamfunctions of the control simulations (see Table 2), in Sv (1 Sv =106m3s−1). Plotted
values are 200yr means except for CCSMog, which is a 150 yr mean. Values in parentheses following the model names are long-term means
for the maximum of the streamfunction below 500m water depth.
with available annual output (standard deviations: LOVE-
CLIM=0.7, CCSMog =1.1, CCSMall =0.9Sv). AMOC
intensity is lower by severalSv in the simulations with a
background meltwater flux. AMOC has a similar structure
in all the control simulations. The northward flow of warm,
salty water occurs in the upper 1000m, while the southward
return flow of North Atlantic Deep Water occurs between
1000–3000m. The anticlockwise cell in the deep ocean,
associated with Antarctic Bottom Water formation, has a
strength of about 4Sv in all control simulations.
The values of AMOC intensity in the control simula-
tions are generally similar to the strength of the modern-day
AMOC (Meehl et al., 2007). Proxy evidence suggests that
the strength of the AMOC during the early Holocene was
probably not that different from today (Bianchi and McCave,
1999; Hall et al., 2004; Oppo et al., 2003; McManus et
al., 2004; Praetorius et al., 2008). There is some proxy ev-
idence for lack of convection and deep water formation in
the Labrador Sea during the early Holocene, however (e.g.,
Hillaire-Marcel et al., 2001; Solignac et al., 2004; Fagel
et al., 2004). To reconcile a vigorous AMOC with lack of
Labrador Sea convection, some other convection area, per-
haps in the Irminger Basin, might have been stronger in the
early Holocene to offset the weaker Labrador Sea convection
(Hall et al., 2010).
The location and strength of convection areas in the North
Atlantic varies significantly among the control simulations
(Fig. 2). Convection occurs primarily in the Nordic Seas in
one of the models (LOVECLIM), primarily in the Irminger
Sea in another (ModelE-R), and in both the Nordic Seas and
just east of the Labrador Sea in the third model (CCSM).
Notably, the background meltwater flux of 0.05Sv does not
shut down convection just east of the Labrador Sea in the
CCSMall control simulation (Fig. 2), as that flux is routed to
the south of the Labrador Sea by ocean surface currents.
Following the 2.5Sv MWP for one year, AMOC intensity
decreases in all simulations (Fig. 3). The maximum decadal-
mean decline in LOVECLIM and CCSM is about 10%,
Clim. Past, 9, 955–968, 2013 www.clim-past.net/9/955/2013/
C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka 959
Fig. 2. Control values of February mixed layer depth, in meters. Plotted values are 100yr means.
Fig. 3. Time series of AMOC intensity anomalies following the
MWP, expressed as a fraction of the long-term control mean. The
MWP of 2.5 Sv for one year was added at Model year 1. AMOC in-
tensity is defined as the maximum value of the overturning stream-
function below 500m water depth (excludes shallow wind-driven
overturning). Heavy lines are decadal averages. Vertical lines on the
right show the 2-sigma range of interannual variability in the con-
trol simulations, and are not shown for ModelE-R since only 30yr
control averages are available.
while for ModelE-R it is about 25%. The decline in AMOC
intensity in LOVECLIM and CCSM is relatively short-lived,
on the order of several decades, and generally within the
range of natural variability of AMOC in their control sim-
ulations. Similarly, mixed-layer depths shoal significantly
following the MWP, but this weakening of convection also
lasts several decades or less (not shown). The response in
ModelE-R is more pronounced and longer-lived, extending
on the order of 100–120yr. Proxy records do not provide
a quantitative estimate of AMOC weakening at 8.2ka, but
do suggest a duration of 100–200yr (Ellison et al., 2006;
Kleiven et al., 2008).
3.2 Ocean salinity and temperature
Significant freshening of the North Atlantic occurs follow-
ing the MWP in all simulations (Fig. 4). The largest anoma-
lies are generally along the coast of Labrador and are up to
1psu when averaged over the first fifty years following the
MWP. Areas of positive SSS anomalies at the mouth of the
St. Lawrence River in CCSMall are caused by cessation of the
0.05Sv background meltwater flux once Lake Agassiz has
drained. Globally, negative anomalies greater than 0.2 psu are
confined to the North Atlantic and Arctic oceans (not shown).
Patterns of SSS anomalies suggest that freshwater trav-
els eastward from the Labrador Sea into the North Atlantic
in all simulations. For most of the simulations, SSS de-
creases in both the Greenland-Iceland-Norwegian Seas and
in the subtropical gyre. This pathway is different from that
inferred by Keigwin et al. (2005), who used δ18O of plank-
tic foraminifera to suggest salinity was decreased near Cape
Hatteras around 8.2 ka. Also, it has been argued that freshwa-
ter released from Hudson Strait would not reach the Nordic
Seas, instead being trapped along the North American coast
(e.g., Wunsch, 2010) or circulating in the subtropical gyre
www.clim-past.net/9/955/2013/ Clim. Past, 9, 955–968, 2013
960 C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka
Fig. 4. Anomalies of annual-mean sea surface salinity in the first fifty years following the MWP relative to the control simulation, in practical
salinity units. Stippling shows statistical significance at the 95% level according to a Student’s ttest. Statistical tests were not performed for
ModelE-R since only decadal averages were available.
Fig. 5. Anomalies of annual-mean sea surface temperature in the first fifty years following the MWP relative to the control simulation, in
degrees Celsius. Stippling shows statistical significance at the 95% level according to a Student’s ttest. Statistical tests were not performed
for ModelE-R since only decadal averages were available.
Clim. Past, 9, 955–968, 2013 www.clim-past.net/9/955/2013/
C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka 961
(Condron and Winsor, 2011). However, several proxy records
from the Irminger and Labrador Seas combine δ18O and
Mg/Ca of planktic foraminera to infer decreases in δ18O of
seawater at 8.2 ka of up to 1 ‰ (Came et al., 2007; Thornalley
et al., 2009; Ellison et al., 2006; Winsor et al., 2012; Hoffman
et al., 2012), which would be equivalent to a freshening of
∼0.7psu assuming the Laurentide Ice Sheet meltwater was
about −25‰ (Hillaire-Marcel et al., 2007). Also, the loca-
tion of detrital carbonate layers deposited around 8.2ka in-
dicate greater freshwater transport in the outer branch of the
Labrador Current, which typically mixes with the North At-
lantic Current and travels to the Nordic Seas (Lewis et al.,
2012). The model simulations presented here, as well as oth-
ers published by Born and Levermann (2010) and Spence et
al. (2008), tend to support some amount of freshwater trans-
port into the Nordic Seas.
Likewise, sea surface cooling is concentrated in the North
Atlantic in all simulations (Fig. 5). Mean anomalies across
the North Atlantic for the first fifty years following the MWP
are on the order of 1◦C, though they exceed 2◦C locally in
the CCSM and ModelE-R experiments. Maximum anomalies
in the LOVECLIM simulation are on the order of ∼0.5◦C
and are located in the far North Atlantic. ModelE-R shows
cooling on the order of several tenths of a degree Celsius
across most of the Southern Hemisphere. The other simula-
tions show little significant change south of 30◦N with the
exception of CCSMall, which has some significant warming
in the south Atlantic.
3.3 Barotropic streamfunction
A common model diagnostic for the ocean circulation, in-
cluding the strengths of the subtropical and subpolar gyres
in the North Atlantic, is the vertically-integrated mass trans-
port (barotropic) streamfunction. Values for this quantity are
available for three of the simulations (Fig. 6). In these three
simulations, transports in both the subtropical gyre and the
subpolar gyre weaken for a few decades following freshwa-
ter forcing. This result is consistent with the concept that re-
duction of deep convection in the core of the subpolar gyre,
as occurs briefly in these simulations in the Labrador and/or
Irminger Seas, weakens this circulation (e.g., H¨
akkinen and
Rhines, 2004). The barotropic streamfunction is not available
as standard output for the fourth simulation (LOVECLIM)
and is not easily calculated offline. Neither deep convection
(Sect. 3.1) nor upper-ocean velocities (top 100m, not shown)
in this model show a large or long-term change forced by the
MWP, though, suggesting that the response of the gyres is at
least qualitatively similar to the other models.
On the other hand, Born and Levermann (2010) found
a prolonged strengthening of the subpolar gyre circulation
in a simulation of the 8.2ka event with the CLIMBER-3α
model. In this model, a reduction of deepwater formation in
the Nordic Seas intensified the subpolar gyre and triggered
internal feedbacks to increase and maintain deep convection
Fig. 6. Control values for barotropic streamfunction (contour lines)
and streamfunction anomalies in the first fifty years following the
MWP relative to the control simulation, in Sv (colored contours).
The contour interval for the control values is 10Sv. Dashed lines
show negative streamfunction values, or a cyclonic circulation. Pos-
itive (negative) anomalies for a cyclonic (anticyclonic) circulation
indicate weakening of the transport. Stippling shows statistical sig-
nificance for anomalies at the 95% level according to a Student’s
ttest. Statistical tests were not performed for ModelE-R since only
decadal averages were available.
in the Labrador Sea. If true, this could explain the onset of
deepwater formation in the Labrador Sea around the time of
the 8.2ka event (e.g., Hillaire-Marcel et al., 2001). The dif-
ferent response in CLIMBER-3αmight be explained by the
fact that the freshwater perturbation had less of a direct im-
pact on the Labrador Sea convection region and instead had
greater advection to the Nordic Seas (Born and Levermann,
2010).
www.clim-past.net/9/955/2013/ Clim. Past, 9, 955–968, 2013
962 C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka
Fig. 7. Control values (contour lines) and anomalies of annual-mean sea ice area in the first fifty years following the MWP relative to the
control simulation (colored contours), in percent. The contour lines show values of 5%, 25 %, 50 % and 75%. Stippling shows statistical
significance at the 95% level according to a Student’s ttest. Statistical tests were not performed for ModelE-R since only decadal averages
were available.
3.4 Sea ice
All of the simulations have areas of significantly expanded
sea ice following freshwater forcing, particularly in the
Labrador Sea and in the Norwegian and/or Barents Sea
(Fig. 7). Generally, these changes for the first fifty years fol-
lowing the MWP are on the order of 5–10%, although they
can be as large as 20–25% in some areas. Sea ice changes in
the Southern Ocean have a heterogeneous spatial pattern and
generally are not statistically significant.
3.5 Surface air temperature
The North Atlantic region and the Arctic become signifi-
cantly colder in most simulations during the first fifty years
following the MWP, with mean annual temperatures in the
multi-model ensemble decreasing less than ∼0.5◦C over Eu-
rope and ∼1.0◦C over Greenland (Fig. 8). These results
hold for individual ensemble members, as well, for both
Europe (40◦N–60◦N, 10◦W–30◦E; anomalies are LOVE-
CLIM=0.0◦C, CCSMog = −0.3◦C, CCSMall = −0.5◦C,
ModelE-R= −0.6◦C) and Greenland (60◦N–80◦N, 60◦W–
20◦W; anomalies are LOVECLIM =0.0 ◦C, CCSMog =
−0.6◦C, CCSMall = −0.4◦C, ModelE-R= −0.8◦C). Tem-
perature changes are minimal in the tropics and the South-
ern Hemisphere. This spatial pattern agrees well with proxy
records, which clearly indicate colder conditions across the
Northern Hemisphere during the 8.2 ka event but suggest that
any Southern Hemisphere temperature changes were likely
regional (Fig. 8).
The magnitude of circum-North Atlantic temperature
changes inferred from proxies is somewhat larger than those
in the models. Temperature reconstructions from pollen and
δ18O in Europe consistently show anomalies of about −1.1 to
−1.2◦C in mean annual temperature during the 8.2 ka event,
although standard errors of these reconstructions are nearly
as large as the anomalies themselves (Veski et al., 2004; von
Grafenstein et al., 1998; Sarmaja-Korjonen and Sepp¨
a, 2007;
Feurdean et al., 2008). Nitrogen isotopes from Greenland in-
dicate temperatures decreased about 2.2 ◦C averaged over the
duration of the event, with an even larger decrease of 3.3 ◦C
during the most extreme 60yr period (Kobashi et al., 2007).
Anomalies over the North Atlantic in the LOVECLIM and
CCSM experiments are short-lived; generally, temperature
values are outside the range of natural variability (defined
as the mean ±2 standard deviations of the control) for less
than two decades (Fig. 9). Anomalies are longer-lived in the
ModelE-R simulation, lasting on the order of 100yr. These
longer-lived anomalies are a better match to high-resolution
proxy records from Europe and Greenland, which consis-
tently show an event duration of 100 to 150yr (Morrill et
al., 2013).
Clim. Past, 9, 955–968, 2013 www.clim-past.net/9/955/2013/
C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka 963
Fig. 8. (Top) Multi-model ensemble mean anomalies of annual-
mean 2-meter air temperature in the first fifty years following
the MWP relative to the control simulations, in degrees Celsius.
Stippling shows grid cells where at least three of the simulations
agree on the sign of the temperature anomaly. (Bottom) Qualita-
tive and quantitative mean-annual temperature anomalies relative to
the early Holocene background climate, in degrees Celsius, inferred
from proxy records for the 8.2ka event, as summarized by Morrill
et al. (2013).
3.6 Precipitation
Despite the noise inherent in precipitation, a number of fea-
tures are common among the model simulations for the fifty
years following the MWP. In all cases, the most important
changes are a reduction in precipitation over the North At-
lantic and Northern Hemisphere tropics, and an increase in
precipitation over the Southern Hemisphere tropics (Fig. 10).
The tropical pattern, consistent with a southward shift of
the mean position of the Intertropical Convergence Zone, is
clearest over the Atlantic Ocean (Fig. 11). Tropical proxy
records from both speleothem δ18O measurements and in-
dicators of lake water balance support this spatial pattern
(Fig. 10).
Several quantitative estimates of drying exist from prox-
ies in high northern latitudes; these include an ∼8% reduc-
tion in accumulation in central Greenland ice cores and an
∼17% reduction in rainfall inferred from pollen north of the
Mediterranean, although again the standard errors of these
reconstructions are nearly as large as the anomalies them-
selves (Feurdean et al., 2008; Pross et al., 2009; Hammer et
Fig. 9. Time series of annual-mean surface air temperature aver-
aged over the region 50–70◦N, 60◦W–10◦E in the North Atlantic,
expressed as anomalies in degrees Celsius from the long-term con-
trol average. The MWP of 2.5Sv for one year was added at Model
year 1. Vertical lines on the right show the 2-sigma range of inter-
annual variability in the control simulations, and are not shown for
ModelE-R since only 30-yr control averages are available.
al., 1997; Rasmussen et al., 2007). The model simulations
generally match the magnitude of drying in central Green-
land, but typically do not match either the direction or mag-
nitude of change in southeastern Europe. Additionally, ev-
idence for wetter conditions at 8.2ka from pollen and lake
geochemical records in northern Europe is not matched by
the freshwater experiments (Fig. 10).
4 Discussion and conclusions
To summarize, the models generally do a good job in repro-
ducing large-scale patterns of temperature and precipitation
changes at 8.2 ka inferred from proxy records. These patterns
include cooling across most of the Northern Hemisphere and
a southward shift of the Intertropical Convergence Zone. The
models have less success in matching the magnitude and
duration of climate anomalies. Temperature changes in the
multi-model ensemble are about half the size of those of
quantitative proxy records from Europe and Greenland. For
all but one of the simulations, the duration of the 8.2ka cli-
mate anomalies is on the order of several decades rather than
the ∼150yr observed in proxy records. Also, there are dis-
crepancies between model and data for some regional-scale
anomaly patterns, including precipitation changes in Europe.
These patterns are less well-constrained by proxy evidence,
however.
The background climate state of the early Holocene, and
the location of convection areas in the North Atlantic more
specifically, might explain some of the differences we see
between models and proxy data. The ModelE-R simulation
has the best match to proxies for event duration, and it has
www.clim-past.net/9/955/2013/ Clim. Past, 9, 955–968, 2013
964 C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka
Fig. 10. (Top) Multi-model ensemble mean anomalies of annual-
mean precipitation in the first fifty years following the MWP rela-
tive to the control simulations, in % change from control. Stippling
shows grid cells where at least three of the simulations agree on the
sign of the precipitation anomaly. (Bottom) Qualitative and quanti-
tative annual-mean precipitation anomalies, in % change from early
Holocene background climate, inferred from proxy records for the
8.2 ka event, as summarized by Morrill et al. (2013).
been previously demonstrated for this model that the lack
of Labrador Sea convection is essential for this response
(LeGrande and Schmidt, 2008; LeGrande et al., 2006). Previ-
ous work with the ECBilt-CLIO model also supports this in-
terpretation; when Labrador Sea convection is weakened by
the background meltwater flux, the ocean’s ability to trans-
port freshwater anomalies away from the North Atlantic is di-
minished and the response to freshwater forcing is prolonged
(Wiersma et al., 2006). On the other hand, lack of convection
in the Labrador Sea does not lead to a long-lived climate re-
sponse in this model’s successor, LOVECLIM. While the ex-
act reasons for this have yet to be determined, the background
meltwater flux used in the LOVECLIM experiment is less
than that in the ECBilt-CLIO experiments (0.05 vs. 0.17Sv;
Wiersma et al., 2006; Li et al., 2009), and it seems that
LOVECLIM is also less sensitive to freshwater perturbations
than its predecessor. Also adding uncertainty to the impor-
tance of convection strength in the Labrador Sea, proxies in-
dicate that AMOC strength was not too different from today
during the early Holocene. In this case, some other convec-
tion area, perhaps in the Irminger Basin (between Greenland
Fig. 11. (Top) Annual-mean precipitation zonally-averaged across
the Atlantic (90◦W–40◦E) in the control simulation, in cmyr−1.
(Bottom) Anomalies of Atlantic annual-mean precipitation for the
first fifty years following the MWP relative to the control simula-
tion, in cm yr−1.
and Iceland), might have been stronger in the early Holocene
to offset the weaker Labrador Sea convection (Hall et al.,
2010). If this was true, the strengthened convection areas
elsewhere might be able to compensate for decreased fresh-
water divergence in the Labrador Sea.
Another factor in the model-data mismatch could be the
size or the complexity of the MWP. The model simula-
tions were forced with 2.5Sv for one year, which was the
best estimate of the flood hydrograph simulations of Clarke
et al. (2004). As these authors point out, though, the total
volume of Lake Agassiz could have generated twice this
forcing and more complex multipulse patterns are possible
(Teller et al., 2002). Their flood model generates a stable
drainage channel that prohibits complete drainage, but this
result might be unlikely for an outburst flood from Lake
Agassiz. In addition, uncertainties in the reconstructed po-
sition of the ice-margin on the northern side of Lake Agas-
siz translate into a range of possible lake volumes spanning
45–200% of the best estimate (Tornqvist and Hijma, 2012).
Clim. Past, 9, 955–968, 2013 www.clim-past.net/9/955/2013/
C. Morrill et al.: Model sensitivity to North Atlantic freshwater forcing at 8.2ka 965
Reconstructions of sea level rise at 8.2ka support the idea
of a larger freshwater drainage. Using peat deposits from the
Mississippi River delta (US), Li et al. (2012) reconstructed a
total eustatic sea level rise of 0.8 to 2.2m at 8.2 ka. Another
reconstruction from the Rhine-Meuse delta (the Netherlands
and Belgium) implies a sea level rise of 3.0±1.2 m (Hi-
jma and Cohen, 2010). These are significantly larger than the
forcing of 2.5Sv for one year (∼0.2m sea level equivalent)
or even than the best estimate of the entire volume of Lake
Agassiz (∼0.4m sea level equivalent). Recent model simu-
lations suggest that the collapse of the Laurentide ice-sheet
saddle around 8.2ka provided this larger volume of fresh-
water (Gregoire et al., 2012), and that this forcing results in
a cooling event that matches many proxy records (Wiersma
and Jongma, 2010; Wagner et al., 2013).
The difference in boundary conditions between the control
simulations does not obviously account for divergent model
responses. As shown in the comparison of the two CCSM
simulations, CCSMog and CCSMall, the addition of a rem-
nant Laurentide Ice Sheet and a background meltwater flux
does not alter the model response to freshwater forcing, ei-
ther in magnitude or duration. It is worth noting, however,
that these boundary conditions were important in previous
experiments with ECBilt-CLIO for prolonging the AMOC
response to Lake Agassiz drainage (Wiersma et al., 2006).
Thus, the effects of these boundary conditions might be very
model-dependent. Differences between early Holocene and
preindustrial orbital forcing and greenhouse gas concentra-
tions are relatively minor, and are not expected to have an
important influence. This should be verified, though, with ad-
ditional model experiments.
Another explanation for the model-data discrepancies is
that the models are not sensitive enough to freshwater per-
turbations. If true, this finding would have important impli-
cations for future climate projections, particularly as models
suggest that continued melting of the Greenland Ice Sheet at
its current rate will have a significant impact on the AMOC
(Hu et al., 2009). There are few model intercomparisons to
determine whether the sensitivity of these three models to
freshwater perturbations is representative of coupled climate
models as a whole. For hosing experiments of 0.1Sv for
100 yr under modern boundary conditions, earlier versions of
the CCSM3 (CCSM2) and LOVECLIM (ECBilt-CLIO) have
AMOC and surface air temperature responses close to the
multi-model ensemble mean (Stouffer et al., 2006). For hos-
ing experiments in a Last Glacial Maximum climate, how-
ever, AMOC decreases somewhat less in the CCSM3 and
LOVECLIM compared to the multi-model ensemble mean
(Kageyama et al., 2012). Improved constraints on the size
of freshwater forcing and its location with respect to early
Holocene convection areas are necessary to rule out the pos-
sibility of inadequate model sensitivity.
Acknowledgements. We thank Lauren Gregoire and an anonymous
reviewer for their helpful comments. Funding for the CCSM
simulations was provided by grants from the U.S. National Science
Foundation, Office of Polar Programs, to CM (ARC-0713951)
and BLO-B (ARC-0713971), and supercomputer time was
provided by a grant from the National Center for Atmospheric
Research (NCAR) Computational Information Systems Laboratory
(CISL). CM and BLO-B thank Nan Rosenbloom for running the
CCSMall simulations, Ellen Ward for assistance with figures, and
Esther Brady and Amy Wagner for helpful discussions. ANL
thanks NASA GISS for institutional support. This is Past4Future
contribution no. 37. The research leading to these results has
received funding from the European Union’s Seventh Framework
programme (FP7/2007–2013) under grant agreement no 243908,
“Past4Future. Climate change – Learning from the past climate.”
Edited by: M. Kageyama
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