Model support for forcing of the 8.2 ka event by meltwaterfrom the Hudson Bay ice dome

Amy J. Wagner • Carrie Morrill • Bette L. Otto-Bliesner •

Nan Rosenbloom • Kelsey R. Watkins

Received: 26 July 2012 / Accepted: 14 February 2013 / Published online: 14 March 2013

� Springer-Verlag Berlin Heidelberg 2013

Abstract Previous model experiments of the 8.2 ka event

forced by the drainage of Lake Agassiz often do not pro-

duce climate anomalies as long as those inferred from

proxies. In addition to the Agassiz forcing, there is new

evidence for significant amounts of freshwater entering the

ocean at 8.2 ka from the disintegration of the Laurentide

ice sheet (LIS). We use the Community Climate System

Model version 3 (CCSM3) to test the contribution of this

additional meltwater flux. Similar to previous model

experiments, we find that the estimated freshwater forcing

from Lake Agassiz is capable of sustaining ocean and

climate anomalies for only two to three decades, much

shorter than the event duration of *150 years in proxies.

Using new estimates of the LIS freshwater flux (*0.13 Sv

for 100 years) from the collapse of the Hudson Bay ice

dome in addition to the Agassiz drainage, the CCSM3

generates climate anomalies with a magnitude and duration

that match within error those from proxies. This result is

insensitive to the duration of freshwater release, a major

uncertainty, if the total volume remains the same. An

analysis of the modeled North Atlantic freshwater budget

indicates that the Agassiz drainage is rapidly transported

out of the North Atlantic while the LIS contribution gen-

erates longer-lasting freshwater anomalies that are also

subject to recirculation by the subtropical gyre back into

the North Atlantic. Thus, the meltwater flux originating

from the LIS appears to be more important than the

Agassiz drainage in generating 8.2 ka climate anomalies

and is one way to reconcile some model-data discrepancies.

Keywords 8.2 ka event � Atlantic Meridional

Overturning Circulation � Freshwater forcing � Coupled

climate model � Abrupt climate change

1 Introduction

Given the likelihood of future reductions in the strength of

the Atlantic Meridional Overturning Circulation (AMOC;

Schmittner et al. 2005), it is important to document how

AMOC variations altered climate in the past and to assess

the skill of coupled climate models in reproducing past

AMOC changes. Of past abrupt changes in the AMOC, the

8.2 ka event provides a particularly useful case study

because its duration (*150 years; Thomas et al. 2007),

estimated magnitude of AMOC reduction (*40 %;

LeGrande et al. 2006) and background climate state are

closest to changes in the North Atlantic expected in the

future due to increased sea surface temperatures (SST) and

ice melt. The hypothesized cause of the 8.2 ka event,

A. J. Wagner � C. Morrill

Cooperative Institute for Research in Environmental Sciences,

University of Colorado, Boulder, CO 80309, USA

A. J. Wagner � C. Morrill

Paleoclimatology Branch, NOAA’s National Climatic Data

Center, Boulder, CO 80305, USA

A. J. Wagner (&)

Center for Marine Science, University of North Carolina

Wilmington, 601 South College Road, Wilmington,

NC 28403, USA

e-mail: wagnera@uncw.edu

B. L. Otto-Bliesner � N. Rosenbloom

Climate and Global Dynamics Division, National Center

for Atmospheric Research, Boulder, CO 80305, USA

K. R. Watkins

Department of Atmospheric and Oceanic Sciences,

University of Wisconsin-Madison,

Madison, WI 53706, USA

123

Clim Dyn (2013) 41:2855–2873

DOI 10.1007/s00382-013-1706-z

haline forcing from the drainage of proglacial Lake

Agassiz-Ojibway (Fig. 1, hereafter Lake Agassiz; Barber

et al. 1999) into the Hudson Bay, is not a perfect analog to

the predominantly thermal forcing of the AMOC that is

predicted for the future (Gregory et al. 2005). Nonetheless,

the 8.2 ka event offers a test of model sensitivity to North

Atlantic surface buoyancy anomalies that has precise dat-

ing, quantified forcing, and a duration short enough to

make simulations with state-of-the-art coupled climate

models feasible (Schmidt and LeGrande 2005; Thomas

et al. 2007; Kobashi et al. 2007).

Previous model experiments of the 8.2 ka event forced

by drainage of Lake Agassiz generally do well in gener-

ating global climate anomalies of the appropriate direction

and magnitude as compared to proxy records (Wiersma and

Renssen 2006; LeGrande and Schmidt 2008). These sim-

ulations produce events as long as the actual 8.2 ka event

only when run with a weaker-than-modern background

AMOC, however. This apparently is the case because an

initially weak North Atlantic circulation, in which there is

slower vertical and horizontal transport, prolongs fresh-

water anomalies that originate from a short meltwater pulse

(Wiersma et al. 2006). However, proxy reconstructions of

ocean transport do not fully support the idea that the

AMOC was substantially reduced during the early Holo-

cene compared to today. In general, grain size analyses and

d13C of benthic foraminifera suggest little difference in the

strength of the early Holocene AMOC compared to mod-

ern, although high variability throughout the Holocene is

observed (Praetorius et al. 2008; Hall et al. 2004; Bianchi

and McCave 1999; Oppo et al. 2003). In addition, the

relatively constant flux of 231Pa/230Th to Atlantic sediments

indicates no significant difference between modern and

early Holocene AMOC (McManus et al. 2004).

In addition to freshwater forcing (FWF) due to the

drainage of Lake Agassiz, there is proxy evidence from sea

level reconstructions (Carlson et al. 2008; Cronin et al.

2007; Tornqvist et al. 2004) and sediment geochemistry

(Carlson et al. 2009) for significant amounts of freshwater

entering the ocean at the time of the 8.2 ka event from the

collapse of the Laurentide ice sheet (LIS) over the Hudson

Bay. Several previously-published experiments examined

the effects of larger volumes of FWF at 8.2 ka, and came to

contradictory conclusions. In some of these sensitivity

experiments, larger volumes of FWF did not always pro-

duce longer and stronger climate responses (LeGrande and

Schmidt 2008), while in others the volume of the fresh-

water pulse was the main predictor of the magnitude of the

AMOC response (Wiersma et al. 2006). An additional set

of experiments found that introducing the LIS forcing

as icebergs rather than as freshwater caused somewhat

more weakening of the AMOC and cooling in Greenland

(Wiersma and Jongma 2010). New quantitative estimates

of the LIS forcing (Carlson et al. 2009) now allow us to

build upon these previous sensitivity studies and to test this

forcing’s contribution to the 8.2 ka event. In this study, we

present simulations using the Community Climate System

Model version 3 (CCSM3) to argue that meltwater from the

collapse of the ice dome over Hudson Bay coincident with

the drainage of Lake Agassiz was an essential forcing of

the 8.2 ka event.

2 Methods

CCSM3 is a global, coupled ocean-atmosphere-sea ice-

land surface climate model coupled without flux correc-

tions (Collins et al. 2006). The atmospheric component is

the Community Atmosphere Model version 3 at a hori-

zontal resolution of T42 (an equivalent grid spacing of

approximately 2.8� in latitude and longitude) and 26 hybrid

vertical levels. The land model, which includes river

Fig. 1 Area of Lake Agassiz at

8.5 ka (blue) and the Laurentide

ice sheet at 8.5 ka (white) and

8.0 ka (stippling). Also shown

are the likely routings of the

Early Holocene background ice

melt flux down the St. Lawrence

River and the flux of freshwater

from Lake Agassiz to the

Labrador Sea (blue arrows).

Striping shows the area of

freshwater addition in the

CCSM3 lake and ice drainage

experiments. Lake and ice sheet

areas are from Dyke (2004)

2856 A. J. Wagner et al.

123

routing and plant functional types, uses the same grid as the

atmospheric component. The ocean model is a version of

the Parallel Ocean Program model with a nominal grid

spacing of approximately 1� in latitude and longitude, with

greater resolution in the tropics and North Atlantic

(approximately 0.3� in latitude). The vertical resolution is

40 levels extending to a depth of 5.5 km. The dynamic-

thermodynamic sea ice model uses the same horizontal grid

and land mask as the ocean component.

We first completed two control simulations with

boundary conditions appropriate to 8.5 ka BP. The first

control simulation (CTL_OG) incorporated early Holocene

orbital parameters and atmospheric greenhouse gas con-

centrations, while the second (CTL_ALL) also included the

LIS and glacial runoff down the St. Lawrence River. As

discussed further below, CTL_ALL has a weaker AMOC

than CTL_OG, allowing us to compare the effects of pre-

cursor AMOC strength on the response to FWF. We used

8.5 ka conditions rather than 8.2 ka due to availability of a

LIS reconstruction at that time interval; any difference in

our specified boundary conditions between these two times

would have been minimal regardless. Atmospheric green-

house gas concentrations were adjusted to early Holocene

values derived from Antarctic ice cores (Fluckiger et al.

2002; Monnin et al. 2004), with CO2 = 260 ppm,

CH4 = 660 ppb, and N2O = 260 ppb. Orbital parameters

were prescribed from the astronomical calculations of

Berger (1978) for 8.5 ka BP (eccentricity = 0.019199,

obliquity = 24.22�, longitude of perihelion = 319.50�),

with higher insolation values in the northern hemisphere in

boreal summer and lower insolation values in boreal winter

compared to present-day values (Fig. 2). The 8.5 ka LIS

was prescribed from ICE-5G (Peltier 2004). Prior to the

final drainage of Lake Agassiz, there was a small baseline

flow of glacial runoff down the St. Lawrence River. Based

on numerical reconstructions of the LIS during the degla-

ciation, the flow is estimated to have been approximately

0.05 Sv (Licciardi et al. 1999). We add this additional

baseline flow as runoff to the ocean model at the mouth of

the St. Lawrence River. CTL_OG was initialized from the

mid-Holocene (6 ka) simulation described in Otto-Bliesner

et al. (2006) and run for over 100 years until it reached

quasi-equilibrium. CTL_ALL was initialized from

CTL_OG and run for over 400 years until North Atlantic

surface salinity stabilized. Global mean ocean salinity

decreases slowly throughout the CTL_ALL simulation due

to the background meltwater flux, a trend that parallels

observed freshening during the late glacial and early

Holocene.

We next completed five experiments with FWF, three

examining the response to Lake Agassiz drainage alone

and two examining the response to forcing from both Lake

Agassiz and the LIS (Table 1). For all of these

experiments, we chose the region 50�–65�N, 35�–70�W to

distribute the FWF into the Labrador Sea (Fig. 1). The

actual path of the freshwater exiting the Hudson Bay

through the Hudson Strait would likely have been along the

Labrador Coast and restricted to the continental shelf until

joining the Gulf Stream and being routed to the northeast

across the Atlantic Ocean (Keigwin et al. 2005; Hillaire-

Marcel et al. 2007; Wunsch 2010). We applied FWF over a

larger region of the Labrador Sea for two reasons. First, the

ocean model component has a fixed volume so freshwater

is added as a negative salinity flux. To avoid negative

salinities when adding large fluxes of freshwater, the FWF

must be added over many model grid cells. Second, the

modeled Gulf Stream is too wide and too zonal compared

to observations, so attempts to inject FWF into just the

boundary current does not route freshwater anomalies

Fig. 2 Time versus latitude diagram of zonally-averaged incoming

solar radiation at the top of the atmosphere (W/m2) for pre-industrial

(top), as well as the difference 8.5 ka minus pre-industrial (bottom)

Model support for forcing of the 8.2 ka event 2857

123

across the North Atlantic in a realistic manner. Deficiencies

in simulating western boundary currents are very common

in models of this resolution (Large and Danabasoglu 2006).

Even though the initial geographic distribution of fresh-

water anomalies is not entirely realistic in our model, as

shown in Sect. 3.2.1 the transport across the North Atlantic

appears reasonable compared with proxies.

For the three drainage experiments, we added the

equivalent of 2.5 Sv of freshwater (Clarke et al. 2004) for

1 year to 8.5 ka control simulations described above. Two

of these experiments, Lake_strong (initial AMOC

*22.4 Sv) and Lake_medium (initial AMOC *19.4 Sv),

were branched from different years of CTL_OG, and one,

Lake_weak (initial AMOC *16.4 Sv), was branched from

CTL_ALL. The difference in initial AMOC between the

two experiments branched from CTL_OG reflects inter-

annual variability in AMOC, both values are within 2r of

the long-term mean of the control (20.0 Sv). In each case,

the system was then allowed to recover following FWF,

unperturbed, for 49 years.

Two additional FWF experiments were performed to

simulate forcing from the collapse of the Hudson Bay ice

dome following the Lake Agassiz drainage. These two

experiments were branched from the same year as

Lake_weak. We chose to use CTL_ALL because this

control experiment provided more realistic patterns of

North Atlantic SST for the early Holocene (Sect. 3.1.1).

The following forcing was applied after the 1-year Agassiz

flood: (1) 0.13 Sv of freshwater for 99 years (Lake ?

Ice_100 yrs) for a total of 15.37 Sv*yrs and (2) 2.5 Sv for

5 years (Lake ? Ice_6 yrs) for a total of 15 Sv*yrs. In

both cases, the model was run for a total of 150 years

following the first addition of freshwater to allow the

system to stabilize and recover. In all of the FWF experi-

ments, the glacial runoff of 0.05 Sv down the St. Lawrence

River was turned off, as this flux was rerouted north out the

Hudson Strait and is incorporated in our FWF (Licciardi

et al. 1999).

The volume and duration of FWF we use for the

collapse of the Hudson Bay ice dome was estimated

using published sea level records (Carlson et al. 2008;

Cronin et al. 2007; Tornqvist et al. 2004) and U/Ca data

from sediment cores in Hudson Bay (Carlson et al.

2009). The volume of Lake Agassiz is estimated at

0.8 9 1014 m3, which is equivalent to approximately

0.2 m of sea level rise (Clarke et al. 2004). Sea level

estimates show the rise in sea level associated with the

8.2 ka event to be between 0.8 and 2.2 m if exclusively

from Lake Agassiz and between 0.5 and 1.4 m if

exclusively from the Hudson Bay ice dome, the differ-

ence between these two estimates being due to the dif-

fering effects of the lake and ice dome on the Earth’s

gravitational field (Li et al. 2012). In addition to sea

level estimates, Carlson et al. (2009) used U/Ca to cal-

culate a total increase in freshwater discharge through the

Hudson Strait after the opening of Hudson Bay of

0.13 ± 0.03 Sv. Carlson et al. (2009) were unable to

provide an exact duration for this enhanced freshwater

flux, but suggest that it lasted less than 500 years. This

flux includes the effects of both enhanced ice sheet

disintegration and the re-routing of rivers, which carry

runoff derived from precipitation–evaporation (P–E) as

well as ice sheet meltwater, from the St. Lawrence River

to the Hudson Bay. Since our 8.5 ka control simulations

predict the change in freshwater discharge to Hudson

Bay associated with river re-routing of P–E to be neg-

ligible (approximately 0.01 Sv), we assumed that nearly

the entire Carlson et al. (2009) estimate is due to ice

sheet disintegration and the northward re-routing of

meltwater. From these estimates, we calculated a forcing

duration of approximately 100 years. These estimates

then form the forcing of our Lake ? Ice_100 yrs exper-

iment, which has a total freshwater addition of

4.85 9 1014 m3 over 100 years that is equivalent to a

1.35 m eustatic sea level rise, and is regarded to be the

most realistic scenario for the collapse of the Hudson

Bay ice dome (Table 1). Our Lake ? Ice_6 yrs sensi-

tivity experiment extends the 2.5 Sv forcing used to

simulate lake drainage for an additional 5 years to obtain

a similar total freshwater addition (4.73 9 1014 m3;

1.31 m sea level rise) over the shortest duration feasible

with our rigid-lid ocean model.

Table 1 FWF experiments

Simulation Flux and duration

of freshwater added

Volume of freshwater

added (m3)

Equivalent

sea level rise (m)

Initial strength

of AMOC (Sv)

Lake_strong 2.5 Sv 9 1 year 0.79 9 1014 0.22 22.4

Lake_medium 2.5 Sv 9 1 year 0.79 9 1014 0.22 19.4

Lake_weak 2.5 Sv 9 1 year 0.79 9 1014 0.22 16.4

Lake ? Ice_100 yrs 2.5 Sv 9 1 year followed

by 0.13 Sv 9 99 years

4.85 9 1014 1.35 16.4

Lake ? Ice_6 yrs 2.5 Sv 9 6 years 4.73 9 1014 1.31 16.4

2858 A. J. Wagner et al.

123

3 Results

3.1 Validation of control simulation

In order to have confidence in the FWF experiments, the

model must exhibit realistic control conditions. To test this,

we compared differences between our 8.5 ka control sim-

ulations and a pre-industrial CCSM3 simulation to pub-

lished proxy data for summer SST and sea surface salinity

(SSS), AMOC and mixed layer depth. We chose only

proxy records with precise age models for this comparison

(radiocarbon age uncertainty \100 year), in order to

achieve the cleanest model-data comparisons.

3.1.1 North Atlantic SST and SSS

Quantitative proxy data for summer SST at 8.5 ka (Table 2)

are limited, but the most robust patterns (i.e., those

corroborated by more than one core) are warming in the

East Greenland Current by up to several degrees Celsius

compared to the pre-industrial (Andersen et al. 2004a, b;

Keigwin et al. 2005) and somewhat cooler temperatures in

the Irminger Sea (\1 �C; Ellison et al. 2006; Farmer et al.

2008; Thornalley et al. 2009). The former is consistent with

increased summer insolation during the early Holocene

(e.g., Marchal et al. 2002), while the latter has been

attributed to the effects of meltwater input to the central

North Atlantic region (Ellison et al. 2006; Thornalley et al.

2009). The difference in June–July–August SST in the

North Atlantic between CTL_ALL and the pre-industrial

simulation has the same general pattern of warming in the

East Greenland Current and cooling in the Irminger Basin

(Fig. 3). The strongest cooling between CTL_ALL and the

pre-industrial simulation is observed in the central North

Atlantic; however no proxy data is currently available for

comparison to the model simulation in this region. Com-

parison to the CTL_OG control indicates that warmer SSTs

in the North Atlantic during the early Holocene can be

explained by orbital forcing, while colder SSTs compared

to the pre-industrial are generally the result of meltwater

input (Fig. 3).

The same analysis was done with summer SSS values

(not shown). There are far fewer proxy records available

for comparison with the model results, but based on the

data available, SSS was lower (fresher) at the mouth of the

St. Lawrence and throughout the North Atlantic during the

early Holocene because of the increase in the baseline flow

of meltwater runoff (deVernal and Hillaire-Marcel 2006;

Solignac et al. 2006; Thornalley et al. 2009). Model results

show similar results with the exception of increased (more

salty) SSS in areas of increased sea ice formation.

3.1.2 AMOC

Proxy evidence for the state of the AMOC during the early

Holocene comes from multiple sources. In general, early

Holocene grain size analysis and benthic foram d13C sug-

gests little difference in the strength of the AMOC com-

pared to modern values although high variability

throughout the Holocene is observed (Praetorius et al.

2008; Hall et al. 2004; Bianchi and McCave 1999; Oppo

et al. 2003). In addition, the relatively constant flux of231Pa/230Th to Atlantic sediments indicates no significant

difference between modern and early Holocene AMOC

(McManus et al. 2004).

In the CCSM3 model simulations, we define the strength

of the AMOC as the maximum of the overturning

streamfunction in the North Atlantic Ocean excluding the

shallow wind-driven overturning. The average AMOC for

the early Holocene control simulations are 16.7 ± 1.1 Sv

and 20.0 ± 1.2 Sv for CTL_ALL and CTL_OG, respec-

tively. The latter is nearly identical to the average AMOC

in a CCSM3 pre-industrial simulation (19.1 ± 0.9 Sv;

Otto-Bliesner et al. 2006). The former can be explained by

the increased freshwater flux (0.05 Sv) from the St. Law-

rence River, and even though it is lower than the simulated

pre-industrial AMOC, it is within the variability and pre-

cision of the proxy data.

Table 2 SST proxy data used in proxy-model comparison

Core Proxy DSST (�C), 8.5 ka—PI Reference

26GGC Mg/Ca 2.6 ± 1.3 Keigwin et al. (2005)

26GGC Alkenones 7.1 ± 0.1 Keigwin et al. (2005)

CR 19/5 Diatom assemblage 0.3 ± 0.8 Andersen et al. (2004a)

LO09-14 Diatom assemblage 0.5 ± 0.7 Andersen et al. (2004b)

MD99-2251 Mg/Ca -0.7 ± 1.2 Farmer et al. (2008)

MD99-2251 Foram assemblage -0.6 ± 2.6 Ellison et al. (2006)

MD99-2269 Diatom assemblage 3.0 ± 0.9 Andersen et al. (2004a)

RAPiD-12-1 K Mg/Ca -0.1 ± 1.9 Thornalley et al. (2009)

PI pre-industrial, defined for this analysis as 1000–1850 A.D. Uncertainty estimates are the 2r differences in mean between 8.5 ka and PI

Model support for forcing of the 8.2 ka event 2859

123

3.1.3 Mixed layer depth

Hillaire-Marcel et al. (2001) suggest that Labrador Sea

convection was not present prior to 7 ka based on dinocyst

and foram data in three sediment cores. This observation

might be reconciled with close-to-modern strength of the

AMOC during the early Holocene in one of several ways.

First, the area of Labrador Sea convection could have been

centered in a different location prior to 7 ka and, therefore,

not detectable in the Hillaire-Marcel et al. (2001) cores.

Second, if deep-water formation was negligible in the

Labrador Sea but the overall strength of the AMOC was

not significantly different, convection must have been

greater than modern in another region of the North

Atlantic. Paleocurrent flow speeds based on sortable silt

(SS) mean grain size data from four North Atlantic Deep

Water (NADW) formation locations suggests convection

may have been increased in the Irminger region or Nordic

Seas during the early Holocene (Hall et al. 2010). This

conclusion is further supported by an increase in fine ter-

rigenous eNd values from the Laurentian Fan around 8 ka

that indicates Labrador Sea Water formation was estab-

lished then at modern values (Hall et al. 2010).

Figure 4 shows the annual-mean mixed layer depths

(MLD) in the North Atlantic for CTL_OG and CTL_ALL

with the locations of the Hall et al. (2010) and Hillaire-

Fig. 3 Summer (JJA) SST

difference between the 8.5 ka

control simulations (CTL_OG,

top, and CTL_ALL, bottom)

and a pre-industrial simulation.

Points indicate locations of

proxy data sites discussed in

text. Stippling indicates

significant difference at the

95 % level according to a

student’s t test

2860 A. J. Wagner et al.

123

Marcel et al. (2001) cores. In the model, areas of strong

convection during the early Holocene exist in the Green-

land–Iceland–Norwegian (GIN) Seas, the Irminger Basin

and east of the Labrador Sea. Given the difficulty and

uncertainties in reconstructing convection, model and

proxy results do not necessarily contradict each other.

The addition of the Laurentide Ice Sheet and the 0.05 Sv

background meltwater flux down the St. Lawrence River in

CTL_ALL compared to CTL_OG does not cause the MLD

to shoal in the Labrador Sea (Fig. 4). Convection in both

8.5 ka controls is also no weaker than in the pre-industrial

control (not shown). This result might seem at odds with

the results of Wiersma et al. (2006), in which the baseline

flow prevents Labrador Sea convection. These authors use

a much larger flux of 0.17 Sv, however.

3.2 Freshwater forcing (FWF) experiments

In the next several sections, we will present the results

from our five FWF experiments and, where applicable,

review the proxy evidence of the 8.2 ka event. We then

compare the proxy results to the model FWF experiment

results and discuss which FWF experiments most closely

match the proxies in duration and magnitude of response.

Fig. 4 Annual mixed layer

depth in the 8.5 ka control

simulations (CTL_OG, top, and

CTL_ALL, bottom). Circle

indicates location of Hall et al.

(2010) sediment core and boxes

indicate Hillaire-Marcel et al.

(2001) sediment core locations

Model support for forcing of the 8.2 ka event 2861

123

3.2.1 Salinity

Measurements of the d18O of seawater show that SSS

decreased around 8.2 ka at several locations in the North

Atlantic (Fig. 5). The proxies show decreased surface and

subsurface salinity values in the Labrador Sea (Hoffman

et al. 2012; Carlson and Winsor 2012) at the Laurentian

Fan (Keigwin et al. 2005), and in the Irminger region

(Came et al. 2007; Thornalley et al. 2009; Ellison et al.

2006). The greatest freshening is in the western Labrador

Sea (Hoffman et al. 2012) and at Gardar Drift (Ellison et al.

2006), with decreases in salinity of 0.8 ppt and 1 ppt,

respectively.

All of the freshwater experiments show a freshening

over most of the North Atlantic (Fig. 5), with the

Lake ? Ice experiments showing a greater magnitude and

more widespread significant differences from the control

than the Lake only experiments. Even though we did not

add freshwater to the model exclusively in the most likely

location for the 8.2 ka event (i.e., along the Labrador

coast), it is reassuring that our experiments nevertheless

yield reasonable magnitudes and areal extents of salinity

anomalies compared to proxy records. In the model, the

area at the mouth of the St. Lawrence River has greater

SSS following FWF. This is due to the re-routing of the

baseline meltwater from the St. Lawrence to the Hudson

Strait at the time of the Lake Agassiz drainage.

3.2.2 Mixed layer depth

The modeled changes in salinity have measurable impacts

on North Atlantic convection, as shown by changes in

MLD. The greatest shoaling of MLD is in the Lake ? Ice

experiments in the GIN Seas, south of Iceland and south

of Greenland (Fig. 6). The Lake only experiments

(Lake_strong and Lake_medium not shown) do have sig-

nificant reductions in MLD in the GIN Seas and south of

Greenland, but only about 2/3 the magnitude of the

reductions seen in the Lake ? Ice experiments. The initial

strength of the AMOC at the time of the FWF in the Lake

Fig. 5 Annual SSS averaged for entire length (50 or 150 years) of

FWF experiments. a Control 8.5 ka simulation (CTL_ALL). b Dif-

ference between Lake_weak and control. c Difference between

Lake ? Ice_100 yrs and control. d Difference between

Lake ? Ice_6 yrs and control. Stippling indicates significant differ-

ence at the 95 % level according to a student’s t test. Colored circles

correspond to the locations of proxy sites. Circle colors correspond to

color scale for model results

2862 A. J. Wagner et al.

123

only experiments does not affect the changes in MLD

while the more important factor appears to be the addi-

tional FWF associated with the Lake ? Ice experiments.

3.2.3 Sea ice

Sea ice coverage also changes significantly in the North

Atlantic, growing in area following the FWF (Fig. 7).

Similar to previously discussed results, much larger and

significant differences between the control and

Lake ? Ice experiments are observed while widespread

weak differences are seen in the Lake only experiments.

There is a greater than 60 % increase in annual sea ice

area during the 50 years of maximum AMOC response

to the FWF in the Lake ? Ice experiments. Sea ice

anomalies have a strong seasonality, with the greatest

changes occurring in the winter and spring (not shown).

Not surprisingly, the areas of greatest increase in sea ice

area are seen south of Greenland, in the GIN Seas, and

south of Svalbard. As follows with a decrease in surface

salinity and increase in stratification, a cool and

fresh ‘‘lid’’ on the ocean surface promotes sea ice

growth.

3.2.4 AMOC

While a quantitative proxy record of the AMOC during

the early Holocene is, as of yet, unavailable, several

qualitative proxy records show a slowdown of the

AMOC during the 8.2 ka event. SS data presented by

Ellison et al. (2006) indicate a decline in the near bottom

flow speed of the Iceland–Scotland Overflow Water

(ISOW) during the 8.2 ka event. The minimum flow

speeds inferred from the SS data lasted 100–200 years.

Kleiven et al. (2008) present benthic d13C data from

Eirik Drift that suggest that low-nutrient Lower North

Atlantic Deep Water (LNADW) was replaced by south-

ern-sourced deep water due to a shoaling or reduction in

AMOC during the 8.2 ka event. The maximum slow-

down of LNADW formation at Eirik Drift persists for

approximately 100 years.

The Lake only experiments (Lake_strong, Lake_

medium, Lake_weak) show a short, slightly significant

decrease in AMOC of about 10 % that recovers after

approximately 25 years (Fig. 8). Additionally, in the

CCSM3, the size and duration of the reduction in AMOC

due to the FWF does not appear to be dependent on the

initial state of the AMOC. The Lake ? Ice_100 yrs

experiment shows a fast, significant reduction (up to a

50 % decrease) in AMOC with the addition of freshwater

that levels out after about 50 years and begins to recover

after another 50 years (Fig. 8). The Lake ? Ice_6 yrs

simulation shows changes in AMOC similar in magnitude

and duration to the Lake ? Ice_100 yrs simulation. The

primary difference between the two Lake ? Ice simulations

is the AMOC decreases more quickly in the Lake ?

Ice_6 yrs experiment than the Lake ? Ice_100 yrs

Fig. 6 Annual mixed layer depth anomalies from an 8.5 ka control

simulation (CTL_ALL). a Difference between Lake_weak and

control for 20 years of maximum AMOC weakening (model years

5–24). b Difference between Lake ? Ice_100 yrs and control for

50 years of maximum AMOC weakening (model years 40–89).

c Difference between Lake ? Ice_6 yrs and control for 50 years of

maximum AMOC weakening (model years 40–89). Stippling indi-

cates significant difference at the 95 % level according to a student’s

t test

Model support for forcing of the 8.2 ka event 2863

123

experiment and then begins to recover in the second

decade following the FWF. However, instead of

remaining near control strength as the Lake only exper-

iments do, the strength of the Lake ? Ice_6 yrs experi-

ment rapidly decreases again and begins a slow recovery

approximately 100 years into the experiment. Further

discussion of the differences between the Lake ? Ice

experiments follows in Sect. 3.3. Overall, though, the

Lake ? Ice experiments match the proxy records more

closely than the Lake only experiments for the duration

of the AMOC slowdown.

Ocean heat transport (OHT) changes parallel changes in

the AMOC (not shown). On average, across the whole

Atlantic basin, OHT decreases by nearly 15 % in the

Lake ? Ice_100 yrs experiment with as much as a 25 %

decrease in the highest north latitudes. The Lake only

experiments show a small basin-wide decrease in OHT with

maximum decreases between 40�–50�N and north of 75�N.

The decrease in OHT is partially, but not entirely, com-

pensated for by an increase in atmospheric heat transport.

3.2.5 Surface temperature and precipitation

Most proxy records for the 8.2 ka event estimate surface

temperature changes. It has been estimated that at the

peak of the cooling associated with the 8.2 ka event,

surface temperatures in Greenland were 3.3 ± 1.1 �C

cooler than the period immediately preceding the event

(Kobashi et al. 2007). Thomas et al. (2007) estimated the

duration of the 8.2 ka event to be approximately

160 years, with the period of maximum cooling only

60–70 years long. While the Lake_strong and Lake_me-

dium experiments produce a significant decrease in

Greenland surface temperature, the duration of the

cooling is less than a decade in length, and the Lake_

weak experiment shows no significant change (Fig. 9). In

contrast, the Lake ? Ice experiments have significant

departures from the control simulation in Greenland

surface temperature that last between 100 and 120 years

in length, much closer to the estimates provided by

Thomas et al. (2007).

Fig. 7 Annual sea ice area (%). a 8.5 ka control experiment

(CTL_ALL). b Difference between Lake_weak and control experi-

ment for 20 years of maximum AMOC weakening (model years

5–24). c Difference between Lake ? Ice_100 yrs and control for

50 years of maximum AMOC weakening (model years 40–89).

d Difference between Lake ? Ice_6 yrs and control experiment for

50 years of maximum AMOC weakening (40–89). Stippling indicates

significant difference at 95 % level according to a student’s t test

2864 A. J. Wagner et al.

123

Proxies also suggest a cooling of 1.0–1.6 �C in Europe

(von Grafenstein et al. 1999; Veski et al. 2004; Sarmaja-

Korjonen and Seppa 2007; Feurdean et al. 2008) and about

0.6 �C in the Canary Current off the west coast of Africa

(Kim et al. 2007). The Lake only experiments show small

surface temperature anomalies in response to the FWF

over, yet these anomalies are too small in regions where

proxy data are available for comparison (Fig. 10a). The

spatial distributions of surface temperature anomalies for

the three Lake only FWF experiments are very similar and,

therefore, only the Lake_weak results are shown in Fig. 10.

The Lake ? Ice experiments produce much larger negative

surface temperature anomalies and show good agreement

with the proxy records (Fig. 10b, c), both in magnitude and

spatial extent. The greatest surface temperature anomalies

are seen in Greenland, northern Europe and the Norwegian

Sea in both of the Lake ? Ice experiments.

Individual high-resolution proxy sites were selected for

comparison with surface temperature results from the

Lake ? Ice_100 yrs experiment. The proxy data presented

are the raw data (not converted to temperature) and the 8.2 ka

event is considered present when three consecutive points

between 7.9 and 8.5 ka exceed the mean ± two sigma of the

early Holocene background climate as calculated from the

periods 7.4–7.9 and 8.5–9.0 ka (Morrill and Jacobsen

2005). A detectable event is seen in Ammersee, Germany

Fig. 8 Maximum Atlantic

meridional overturning

streamfunction of freshwater

experiments (3-year running

averages), expressed as

anomalies from the appropriate

control average. Shaded box

shows the 2-sigma (2r)

variability of the control

simulation

Fig. 9 Greenland surface

temperature (70�–80�N, 20�–

50�W) anomaly (3-year running

averages) compared to the

appropriate control average.

Shaded box shows the 2-sigma

(2r) variability of the control

simulation

Model support for forcing of the 8.2 ka event 2865

123

(von Grafenstein et al. 1999) from ostracod d18O measured in

lake sediment cores that corresponds to a significant decrease

in surface temperature observed in the model results from the

same region (Fig. 11a). A similar response in speleothem

d18O is seen in nearby Katerloch Cave, Austria (Fig. 11b;

Boch et al. 2009). Both the ostracod and speleothem d18O are

thought to reflect mean annual temperature, although con-

tributions from changes in the d18O of the precipitation

source region in the North Atlantic at 8.2 ka are also likely

(LeGrande and Schmidt 2008). The duration of the event as

defined by the statistical test (two sigma deviation from the

pre-event mean) is typically shorter in the proxy records than

in the model results; however, the difference is not large and

is affected by the choice of statistical threshold, as well as by

proxy resolution and age model.

The greatest climatic responses to the 8.2 ka event are

typically observed in and around the North Atlantic basin.

However, as more proxy records of temperature and pre-

cipitation from all over the globe are published, it is

becoming clearer there was a global response to the FWF

associated with the 8.2 ka event (e.g., Morrill and Jacobsen

2005; Cheng et al. 2009). Consistent with previous FWF

experiments and our dynamical understanding of climate

system response to freshening of the North Atlantic (e.g.,

Wiersma and Renssen 2006; LeGrande and Schmidt 2008),

the Lake ? Ice model results indicate the southern hemi-

sphere tends to be warmer and wetter during the 8.2 ka

event while the northern hemisphere shows varying

degrees of cooling and decreased precipitation (Fig. 12).

3.3 Freshwater removal from the North Atlantic

An analysis of the freshwater surface fluxes and transport

for the North Atlantic (35�N–80�N) provides insight into

differences between the various FWF experiments.

Because the CCSM3 ocean component uses a virtual salt

flux rather than freshwater volume flux for freshwater

exchange between the ocean and atmosphere, we calculate

the meridional ocean freshwater transport following Hu

et al. (2008), using meridional ocean velocity and ocean

salinity relative to a reference salinity. We present in

Fig. 13 all terms of the North Atlantic freshwater budget,

expressed in terms of cumulative anomalies relative to the

amount of added freshwater. These cumulative time series

track the progress of the ocean model in removing fresh-

water from the North Atlantic (e.g., 100 % marks complete

divergence of all added freshwater). Likewise, a freshwater

residual plot (Fig. 13g) shows the amount of freshwater

addition still remaining in the North Atlantic. There are

two main questions we try to address in this section: (1)

Why does the Lake only forcing fail to generate an event

comparable in duration to proxy records? and (2) Why

are the outcomes of the Lake ? Ice_100 yrs and Lake ?

Ice_6 yrs experiments so similar?

To the first order, all simulations have a similar temporal

evolution of the North Atlantic freshwater balance fol-

lowing FWF. Immediately after FWF commences, excess

freshwater begins to be efficiently diverged out of the

Fig. 10 Annual surface temperature anomalies from 8.5 ka control

(CTL_ALL) averaged for entire length (either 50 or 150 years) of

experiments. a Difference between Lake_weak and control. b Differ-

ence between Lake ? Ice_100 yrs and control. c Difference between

Lake ? Ice_6 yrs and control. Colored circles correspond to proxy

sites with an estimated change in surface temperature during the

8.2 ka event. Circle color corresponds to color scale for model

experiments

2866 A. J. Wagner et al.

123

North Atlantic (negative values in Fig. 13f). Northward

and southward freshwater transports, as opposed to changes

in surface freshwater fluxes, are the main drivers of the

divergence (Fig. 13c–e). One to two decades following the

onset of FWF, freshwater begins to be gained each year

rather than lost at the southern boundary (35�N; change

from negative to positive slope in Fig. 13c). This is due to

recirculation of freshwater in the subtropical gyre. In all

simulations except for Lake ? Ice_6 yrs any gain of

freshwater at the southern boundary is offset by continued

northward transport of freshwater at the northern boundary

(80�N; negative values in Fig. 13d). Freshwater transport at

the northern boundary is also subject to some recirculation,

in the form of a net southward transport of sea ice across

80�N that then melts in the North Atlantic (positive values

in Fig. 13a). When taking the sea ice melt into account, the

amount of freshwater transport northward is generally

equal to the amount of freshwater transport southward.

Surface fluxes of freshwater with the atmosphere are rel-

atively unimportant, although they do slightly add to

freshwater divergence in the Lake ? Ice experiments, as

precipitation is decreased more than evaporation following

FWF (negative values in Fig. 13b).

We argue that the minor response in the Lake only

experiments is mainly a function of the small volume and

duration of FWF (Fig. 7). Freshwater is quickly transported

Fig. 11 Proxy-model

comparison of surface

temperature for sites in northern

Europe: a Ammersee, Germany

(von Grafenstein et al. 1999)

and b Katerloch Cave, Austria

(Boch et al. 2009). Model time

series are decadal-averaged

anomalies from the early

Holocene control simulation

(250 years) using a 10� 9 10�box centered on the proxy

location. Error bars show 95 %

confidence interval for the

difference between the

Lake ? Ice_100 yrs freshwater

experiment and 8.5 ka control

simulation (CTL_ALL).

Shading indicates the 100 years

of FWF. Dashed lines show the

thresholds for event detection in

the proxy record, as described in

Morrill and Jacobsen (2005)

Model support for forcing of the 8.2 ka event 2867

123

out of the North Atlantic in all simulations, with typically

about 80 % of the cumulative freshwater addition gone

after a decade or two (Fig. 13f). For a small and short FWF

such as applied in the Lake only experiments, this leaves

little residual freshwater (1 Sv*yr or less) in the North

Atlantic to affect surface buoyancy and the AMOC

(Fig. 13g).

The Lake ? Ice_100 yrs and Lake ? Ice_6 yrs experi-

ments show strikingly similar results despite the difference

in length of FWF. Both experiments show a dramatic

decrease in AMOC strength with the addition of the FWF

(Fig. 8), but the Lake ? Ice_100 year experiment shows a

gradual decrease in AMOC strength for the first 50 years of

the experiment while the AMOC in the Lake ? Ice_6 yrs

experiment rebounds in the first decade following the FWF,

then sharply declines at the end of the second decade for the

following 30 years. The Lake ? Ice_6 yrs experiment

eventually reaches an AMOC strength about 8 Sv less than

the control, the same decrease as observed in the

Lake ? Ice_100 yrs experiment, and recovers at about the

same rate as the Lake ? Ice_100 yrs experiment. The

second decline in AMOC strength in the Lake ? Ice_6 yrs

experiment, which occurs from model years 25–60, is lar-

gely the result of an increase in the residual freshwater in

the North Atlantic (from *3 Sv*yrs to *6 Sv*yrs,

Fig. 13g) following recirculation of freshwater in the

Fig. 12 a Annual surface

temperature anomalies of

Lake ? Ice_100 yrs compared

to control (CTL_ALL) for entire

length of experiment

(150 years). b Annual

precipitation anomalies of

Lake ? Ice_100 yrs compared

to control for entire length of

experiment. Stippling indicates

significant difference at 95 %

level according to a student’s

t test

2868 A. J. Wagner et al.

123

subtropical gyre (change from negative to positive slope,

Fig. 13c). This, along with the influence of sea ice melt

(Fig. 13a), acts to prolong the AMOC response to the short

6-year forcing. At about model year 80, southward fresh-

water divergence at 35�N reaches a near-equilibrium with

northward freshwater recirculation at 35�N (near-zero slope

in Fig. 13c), and the North Atlantic freshwater residual

begins to decrease gradually due mainly to northward

transport at 80�N (Fig. 13d).

Residual freshwater in the North Atlantic evolves with a

similar temporal pattern in the Lake ? Ice_100 yrs

experiment, though with muted amplitude since freshwater

is added gradually (Fig. 13g). In Lake ? Ice_100 yrs, a

volume of freshwater (*2 Sv*yrs) nearly equivalent to the

initial Lake pulse exists as a residual during the hosing

period, when the amount of freshwater divergence is gen-

erally equal to the volume of the 0.13 Sv flux (near-zero

slope in Fig. 13g). Once the 0.13 Sv freshwater flux is

Fig. 13 Freshwater balance in

the North Atlantic, calculated as

cumulative anomalies from the

appropriate control simulation

and expressed relative to the

total amount of FWF added in

each experiment. a Cumulative

freshwater anomaly from sea ice

melt minus formation for 35�N

to 80�N, b cumulative surface

freshwater input anomaly from

precipitation minus evaporation

plus runoff for this region,

c cumulative freshwater

transport across 35�N,

d cumulative freshwater

transport across 80�N, e total

cumulative freshwater transport

out of the 35�N to 80�N region,

f Total cumulative freshwater

divergence (sum of sea ice,

atmospheric and transport

fluxes) out of the region,

g cumulative residual amount of

freshwater remaining in the

region. Positive (negative)

numbers indicate a freshwater

gain (loss) from the 35�–80�N

region. The latitude of 80�N is

defined as along the line of

80�N latitude from 90�W to

30�E and along the line of 30�E

from 80� to 70�N

Model support for forcing of the 8.2 ka event 2869

123

turned off at year 100, the remaining residual freshwater is

diverged from the North Atlantic over the next 50 years

(positive slope in Fig. 13g).

4 Discussion

A number of modeling experiments of the 8.2 ka event

have been previously performed (i.e., Wiersma et al. 2006;

LeGrande and Schmidt 2008; Clarke et al. 2009; Tindall

and Valdes 2011) using different boundary conditions,

model resolutions and freshwater fluxes. Several of these

experiments used freshwater fluxes comparable to our Lake

only simulations (2.5–5.0 Sv for 1 year) and yield similar

results when initial AMOC strength is relatively strong. For

these other models, the decrease in maximum AMOC

following FWF ranges from 15 to 50 % and lasts less than

60 years in all cases. Our CCSM3 results are comparable,

showing a decrease in maximum AMOC of about 10 %

that lasts less than 25 years.

Two modeling groups have performed additional FWF

experiments from a weakened initial state of the AMOC,

namely reductions of 45 and 25 % compared to their strong

states (Wiersma et al. 2006; LeGrande and Schmidt 2008).

In these cases, anomalies in maximum AMOC are slightly

smaller than for the strong states, but last up to 150 years

and are in much better agreement with the duration of the

8.2 ka event inferred from proxy records. In contrast, we

find little difference in response to the initial state of the

AMOC in our Lake only experiments. In all our cases, the

AMOC anomalies last only two to three decades. It is

worth noting, however, that our weakened state is not as

different from our strengthened state (\20 % reduction) as

in those previous simulations. Given that our Lake ? Ice

experiments began from the relatively weakened initial

state of the AMOC, we are unable to comment on the

effects of initial AMOC strength on the response to the

larger volume forcing. We expect based on the Lake only

experiments and the relatively small differences in AMOC

strength between our weak and strong states, however, that

a Lake ? Ice experiment begun from CTL_OG would not

yield substantially different results.

Past modeling experiments using larger volumes of

FWF were treated as sensitivity experiments and were

considered at the time to be unrealistic in the amount or

rate of FWF (i.e., LeGrande and Schmidt 2008; Wiersma

and Renssen 2006). Wiersma et al. (2006) presented results

from ECBilt-CLIO-VECODE experiments with forcing of

15 Sv*yr (4.67 9 1014 m3) added over 5 years, similar to

our Lake ? Ice_6 yrs simulation. In their simulations,

when this forcing was applied to a strong initial AMOC,

the resulting event lasting about 150 years and was in good

agreement with the duration inferred from proxy evidence.

The event was twice as long when begun from an initial

state with AMOC weaker by 25 %, however. A somewhat

smaller forcing of 10 Sv*yr (3.145 9 1014 m3) applied in

2 years also extended climate anomalies to *100–150

years duration in the GISS ModelE-R (LeGrande and

Schmidt 2008). Thus, our results are in line with previous

experiments and indicate robustness across models in their

sensitivity to FWF. Given current estimates of total FWF

around 8.2 ka from both lake drainage and ice sheet con-

tributions, however, initial states with relatively stronger

AMOC might yield more realistic results in terms of event

duration than those with weaker AMOC.

Our conclusion that the duration of FWF is less

important than the volume is also supported by previous

model results. Wiersma et al. (2006) came to a similar

conclusion for short release durations (5 years or shorter),

and Wiersma and Renssen (2006) showed that this was also

true up to a release duration of 20 years. Our results indi-

cate that even longer release durations up to 100 years,

which might be the most realistic in term of ice sheet

dynamics, yield ocean and climate responses similar to the

short release experiments. One caveat to this conclusion is

the fact that a 6-year forcing might have been distributed

over a somewhat smaller area of the North Atlantic than a

100-year forcing, which is not the case in our experiments.

An additional experiment could test whether the 6-year

duration FWF coupled with a smaller area of injection

would produce different results.

One possibly unrealistic aspect of our simulations is the

relatively large area over which freshwater is released. As

we argued above, however, the spatial patterns of SSS

anomalies between our FWF experiments and the control

simulation (Fig. 5) seem reasonable when compared to

proxy reconstructions of SSS for the 8.2 ka event. Also,

despite the large differences in forcing location among

8.2 ka simulations using other models, which range from a

few grid cells in the Labrador Sea to all grid cells across

the North Atlantic between 50�N and 70�N, the climate

outcomes are quite similar. Likewise, experiments com-

paring the effects of freshwater release evenly over the

Labrador Sea to a more precise injection along the western

boundary find no significant differences in response

(Spence et al. 2008). From these observations, we argue

that any uncertainty introduced through our choice of

forcing location is minimal.

Some previous studies have concluded that model

‘‘weather’’ is significant enough to affect the outcomes of

FWF and that an ensemble of model simulations with

different initial conditions is necessary (e.g., LeGrande

et al. 2006; Tindall and Valdes 2011), while another has

concluded the opposite (Wiersma et al. 2006). In models

exhibiting an unstable state of the AMOC, high-fre-

quency natural variability can make AMOC recovery

2870 A. J. Wagner et al.

123

unpredictable when the FWP is shorter than a decade or

two (Renssen et al. 2002). We were not able to complete

such an ensemble due to the high resolution and compu-

tational expense of our model simulations. However, the

three Lake only runs form an ensemble of sorts with dif-

ferent initial AMOC strengths. These simulations show

similar responses to the FWF, leading us to expect little

dependence on initial conditions if we were to run a true

ensemble. It is worth noting, however, that the CCSM3

appears to have less intrinsic variability in the control

simulation than has been reported for several of the other

models used in 8.2 ka simulations, and the appropriate

amount of natural variability for the early Holocene is

unknown.

We have introduced FWF associated with the break-up

of the Laurentide Ice Sheet as liquid water rather than ice,

which might be more realistic. A recent study by Wie-

rsma and Jongma (2010) compared the difference in

NADW export, sea ice area and Greenland temperature

between a freshwater release and an equivalent amount of

freshwater introduced to the North Atlantic as icebergs.

While the modeled response to the iceberg perturbation is

somewhat greater than that from the freshwater alone, the

duration of the response is no different between the two

experiments. The authors conclude that the effect of latent

heat of melting icebergs on SST and sea ice formation

should not be disregarded when performing these types of

freshwater hosing experiments. In this and other similar

sensitivity studies for the Last Glacial Maximum (Levine

and Bigg 2008), any differences in forcing location aris-

ing from perturbing the ocean with icebergs rather than

freshwater have little impact on the overall ocean and

climate response. Thus, the magnitude of the climate

response we find in the Lake ? Ice experiments should

perhaps be regarded as a minimum estimate, although our

good match to proxy evidence for event duration should

be robust.

Some proxy evidence points to the likelihood of multi-

ple freshwater pulses around 8.2 ka, separated by a century

or more and possibly caused by a multi-staged drainage of

Lake Agassiz (e.g., Ellison et al. 2006; Lewis et al. 2012).

In our experiments, as well as in most previous experi-

ments, the FWF occurs in one pulse. Wiersma and Jongma

(2010) simulate a two-stage forcing, in which a first lake

outburst is followed after 200 years by a second outburst

and discharge of icebergs from the LIS. They find that this

scenario can account for the multi-phased sequence of

events recorded in one North Atlantic sediment core

(Ellison et al. 2006), but does not produce the single period

of cooling recorded in Greenland ice cores. As the tem-

poral pattern of FWF around 8.2 ka becomes quantified

better and more high-resolution records of the climate

response become available, these will provide refined tar-

gets for further model simulations.

5 Summary and conclusions

In this study, we performed several experiments to sim-

ulate the 8.2 ka event, varying the initial strength of the

AMOC and the amount and duration of FWF. In the first

set of experiments, we added freshwater to the Labrador

Sea to simulate the drainage of Lake Agassiz (2.5 Sv for

1 year). The climate response to this forcing was quite

consistent across simulations and appears to be indepen-

dent of the initial strength of the AMOC. Nonetheless, our

results indicate that the volume of freshwater contained in

Lake Agassiz is not enough to generate a climate response

that matches proxy records of the 8.2 ka event in terms of

magnitude and duration. This finding, particularly

regarding event duration, is consistent with previous

simulations completed with other models. In two addi-

tional experiments, we used recently published estimates

of the total freshwater flux around 8.2 ka (lake drainage

plus melt from the Laurentide Ice Sheet) and found a good

match to proxy records. These experiments also demon-

strate that the volume, and not the duration, of FWF

appears to be the more important factor in generating

appropriate 8.2 ka anomalies. This finding highlights the

importance of estimates of sea level rise around the 8.2 ka

event in establishing the total volume of freshwater added

to the ocean during the event. We conclude that the FWF

from the Laurentide ice sheet was more important than the

Lake Agassiz flood itself in causing the 8.2 ka event to

persist and is one way to reconcile some model-data

discrepancies.

Acknowledgments We thank two anonymous reviewers for their

helpful comments and suggestions. This research was funded by

grants from the National Science Foundation, Office of Polar Pro-

grams, to C. M. (ARC-0713951) and B. O.-B. (ARC-0713971) and

the NOAA Hollings Scholar Program to K. W. Proxy data was har-

vested from the World Data Center for Paleoclimatology. This

manuscript benefited from insightful conversations with E. Brady,

A. Carlson, and Z. Liu. Supercomputer time was provided by a grant

from the National Center for Atmospheric Research (NCAR) Com-

putational Information Systems Laboratory (CISL). The National

Center for Atmospheric Research is sponsored by the National Sci-

ence Foundation.

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