Glacial abrupt climate changes and Dansgaard-Oeschger ... paper.pdf · Glacial abrupt climate changes and Dansgaard-Oeschger oscillations in a coupled climate model Zhaomin Wang and

PALEOCEANOGRAPHY, VOL. ???, XXXX, DOI:10.1029/,

Glacial abrupt climate changes and

Dansgaard-Oeschger oscillations in a coupled

climate model

Zhaomin Wang and Lawrence A. Mysak

Earth System Modelling Group, Department of Atmospheric and Oceanic

Sciences, McGill University, Montreal, Quebec H3A 2K6, Canada

Z. Wang, Earth System Modelling Group, Department of Atmospheric and Oceanic Sciences,

McGill University, Montreal, Quebec H3A 2K6, Canada. ([email protected])

D R A F T November 3, 2005, 11:05am D R A F T

X - 2 WANG AND MYSAK: GLACIAL ABRUPT CLIMATE CHANGES AND D-O OSCILLATIONS

There are three fundamental features which characterize large glacial mil-

lennial (Dansgaard-Oeschger) oscillations: 1) the climatic transitions were

abrupt and large; 2) the lengths of both interstadials and stadials and the

period of Dansgaard-Oeschger oscillations were not uniform; and 3) there were

no large millennial oscillations during an early stage of a glacial period and

a peak glacial period. In this modelling study, we offer a consistent expla-

nation for these three features by employing an Earth system Model of In-

termediate Complexity. We demonstrate that a moderate global cooling forces

the Atlantic Meridional Overturning Circulation (MOC) into an unstable state

and hence causes the flip-flop of the Atlantic MOC between a strong mode

and a weak mode. The durations of both interstadials and stadials are mod-

ulated by the changing background climate, in qualitative agreement with

the observations. In a warm climate, the Atlantic MOC is strong and sta-

ble, with the deep water formed mainly by intense heat loss to the atmosphere.

In a cold climate, the Atlantic MOC is weak and stable, and this mode is

largely maintained by the process of sea-ice brine rejection. Hence we con-

clude that brine rejection plays a necessary role in the oscillations, confirm-

ing a hypothesis suggested in some proxy data studies.


WANG AND MYSAK: GLACIAL ABRUPT CLIMATE CHANGES AND D-O OSCILLATIONS X - 3

1. Introduction

Greenland ice core records reveal that much of the last glacial period was punctuated

by millennial-scale climatic fluctuations, which are termed Dansgaard-Oeschger (D-O)

oscillations [Dansgaard et al., 1993; NGICP, 2004]. A much longer North Atlantic sedi-

ment core record shows that D-O oscillations also occurred during previous glacial periods

[McManus et al., 1999]. There are three key features of D-O oscillations. Firstly, there

are abrupt and large climatic transitions in the oscillations. During one cycle of the

oscillations, there is a sudden warming over several decades or less, a gradual cooling

phase followed by a rapid cooling, and finally a gradual warming phase leading up to

the next sudden warming. Shifts of 9 to 16 oC were observed over Greenland during a

cycle [Schwander et al., 1997; Lang et al., 1999]. Secondly, the millennial oscillations are

quasi-periodic, with the period being modulated by the nature of the glacial background

climate: long interstadials (warm states) occurred during late (oxygen isotope) stage 5

and early stage 3 when the background glacial climates were relatively warm, while long

stadials (cold states) and short interstadials occurred during stages 4 and 2, when the

background glacial climates were relatively cold [Bond et al., 1999; Schulz, 2002]. The

D-O periods were thus relatively long when the background climate was either relatively

warm or cold; when the background climate state was at an intermediate phase, the pe-

riods were relatively short. Thirdly, the extensive North Atlantic Ocean sediment record

reveals that the millennial oscillations were significantly suppressed when the background

climate was either warm (interglacial and early stage of a glacial) or very cold (peak or

maximum glacial) [McManus et al., 1999].



Numerous modelling studies have investigated the possible causes of millennial oscilla-

tions, with the model complexity ranging from box [Birchfield et al., 1994] to comprehen-

sive general circulation models [Tziperman, 2000], and from stand-alone ocean circulation

models [Marotzke, 1989] to coupled multi-component climate models [Ganopolski and

Rahmstorf, 2001; 2002; Knutti et al., 2004]. There have also been many proxy data stud-

ies of D-O oscillations [Broecker and Denton, 1990; Bond et al., 1999; Dokken and Jansen,

1999; van Kreveld et al., 2000], and a variety of mechanisms have been proposed for these

oscillations. Early studies suggested that D-O oscillations can be generated by internal

oceanic processes [Broecker and Denton, 1990; Weaver, 1999]. Recently, in coupled cli-

mate models under the forcing of glacial boundary conditions, an external weak periodic

millennial-scale freshwater forcing with noise excitation [Ganopolski and Rahmstorf, 2002]

or strong freshwater forcing [Sakai and Peltier, 1997; Timmermann et al., 2003; Knutti et

al., 2004] triggered D-O oscillations. Ice sheet-thermohaline circulation interactions have

also been proposed as a cause for D-O oscillations [Birchfield et al., 1994; van Kreveld et

al., 2000; Wang and Mysak, 2001]. Northern North Atlantic sea-ice brine rejection has

been found to play an important role during the oscillations in some proxy data studies

[Dokken and Jansen, 1999; van Kreveld et al., 2000]. Routing switch of river discharge

has also been recently hypothesized as a potential mechanism for the D-O oscillations

[Clark et al., 2001]. In Gildor and Tziperman [2003] and Kaspi et al. [2004], sea ice

switch-like behavior driven by variability in a weak ocean circulation causes large temper-

ature changes during a glacial period. In their box model studies, Sima et al. [2004] and



Olsen et al. [2005] have simulated that the oscillations are modulated by slowly changing

background forcing.

In this study, we invoke the stability of the glacial Atlantic MOC in a slowly changing

background climate in an Earth system Model of Intermediate Complexity (EMIC) to

answer the following three major questions: 1) Why were the glacial climatic transitions

abrupt and large? 2) Why were the lengths of both interstadials and stadials and the

period of D-O oscillations not uniform? 3) Why were there no large millennial oscillations

during an early stage of a glacial period and a peak glacial period?

While early modelling studies [Ganopolski and Rahmstorf, 2001; Schmittner et al., 2002]

have shown that the stability of the glacial Atlantic MOC was reduced, the thermal modu-

lation effects on the stability of this circulation by the slowly changing glacial background

climate have not been thoroughly studied in state-of-the-art coupled climate system mod-

els. We will show that the Atlantic MOC is not stable in a background climate which is

undergoing a period of moderate cooling. This Atlantic MOC instability results in D-O

oscillations, which are characterized by flip-flops between strong and weak modes. The

Atlantic MOC is stable when the background climate is either relatively warm or rela-

tively cold, with the circulation being strong for the relatively warm climate and weak for

the relatively cold climate. We will also show that the durations of the interstadials and

stadials vary with the background climate, and also that the period of the oscillations

is not uniform throughout the glacial period. Thus in this modelling study, we offer a

consistent explanation for the above three features of D-O oscillations using an EMIC. For

the first time, we will also show theoretically that sea ice brine rejection plays a necessary



role in maintaining the weak Atlantic MOC mode of these oscillations, which corroborates

a hypothesis suggested in some proxy data analysis studies [Dokken and Jansen, 1999;

van Kreveld et al., 2000; Clark et al., 2002].

2. Model and experimental design

In this study, the McGill Paleoclimate Model-2 (MPM-2) [Wang, 2005] is employed.

The MPM-2 is an extension of Wang and Mysak [2000, 2002], and Wang et al. [2005],

with several major improvements. The variables in the MPM-2 are sectorially or zonally

averaged across each continent and ocean basin, with variables in or over North America

and Eurasia downscaled to a 5 by 5 degree resolution in order for us to couple a two-

dimensional dynamic ice sheet model [Marshall and Clarke, 1997] to the model [Wang and

Mysak, 2002]. A new parameterization for the solar energy disposition, which concerns the

amount of solar energy absorbed by the atmosphere and by the surface and the amount

of solar energy reflected to space, has been employed [Wang et al., 2004], and a global

dynamic vegetation model (VECODE: VEgetation COntinuous DEscription) [Brovkin et

al., 2002] has been coupled to the model [Wang et al., 2005]. In the MPM-2, the model

domain is extended from (75o S, 75o N) to (90o S, 90o N), surface winds are parameterized

[Petoukhov et al., 2000], and there are no oceanic heat and freshwater flux adjustments

[Wang, 2005]. The MPM-2 is now a global climate model which consists of atmosphere,

ocean, sea ice, land surface, continental ice and vegetation components.

In the control experiment of this study, the standard LGM boundary conditions used

by the PMIP project [Joussaume and Taylor, 2000] are employed to obtain a glacial

background climate, i.e., the orbital forcing is for 21 kyr BP, the ice sheets are prescribed



as at the LGM [Peltier, 2004], the atmospheric CO2 level is at 200 ppm, and the global

salinity is increased one p.p.t. to take into account the global sea-level drop. The sea-ice

meridional advection velocity is prescribed: its annual mean value increases from 0.15 cm/s

in the Arctic Ocean to 3.0 cm/s at 40o N. In the Southern Ocean, the prescribed sea-ice

velocity increases from 1 cm/s at 70o S to 5 cm/s at 40o S. A white noise freshwater forcing

is also applied over the 45 to 75o N latitude band of the North Atlantic. The standard

deviation of the integrated white noise freshwater flux is 0.03 Sv and the decorrelation

time step is one year.

3. Abruptness, periodicity and Atlantic MOC modes

Under the forcing of the glacial boundary conditions and the imposed white noise fresh-

water forcing over 45-75o N of the North Atlantic, the model was run for 20 thousand

years in the control experiment. After the spin-up phase, a periodic oscillation persists

throughout the whole model run, with a period of approximately 1,400 years. It is argued

in Winton and Sarachik [1993], Ganopolski and Rahmstorf [2002] and Timmerman et al.

[2003] that ocean dynamics mainly accounts for this time scale .

Fig. 1 shows the time series of the Atlantic MOC (we use the maximum streamfunction

below the Ekman layer in the Atlantic to represent the intensity of the Atlantic MOC),

Atlantic sea surface salinity at 57.5o N, surface air temperature (SAT) at 62.5o N over the

North Atlantic and SAT at 62.5o S for the time interval from 10 to 15 thousand model

years. During one cycle, the Atlantic MOC jumps abruptly from a relatively weak state

to a peak one and then slowly weakens; however it stays at a strong state for several

hundred years. When the Atlantic MOC weakens to a critical value, it abruptly switches



to a weak state and then it takes almost 1000 years to recover until the next abrupt jump

(Fig. 1a). Similar behavior of the Atlantic MOC was also simulated in Ganopolski and

Rahmstorf [2001] by prescribing a periodic and weak freshwater forcing. The evolution of

the sea surface salinity in the North Atlantic has the same abrupt and periodic behavior,

with an amplitude of 2 p.p.t. (Fig. 1b). This is similar to what is found in Ganopolski

and Rahmstorf [2001] and it is in agreement with proxy data reconstructions [van Kreveld

et al., 2000].

The change of zonally averaged Atlantic SAT over 60-65o N clearly shows rapid northern

climate transitions, either from a cold state (stadial) to a warm state (interstadial) or from

a warm state to a cold state on a multi-decadal time scale (Fig. 1c). After each rapid

warming or cooling, there is a gradual cooling or warming, respectively. The maximum

temperature change is up to 6 oC. The zonally averaged SAT changes over North America

and Eurasia at the same latitude band are only 2.5 oC and 1.4 oC respectively, and the

zonally averaged Pacific SAT change over 55-60o N is just 1.8 oC , indicating that the

North Atlantic is the origin of the warming and cooling events during D-O oscillations.

The strong cooling and warming events are mainly caused by the Atlantic MOC mode

switches, but these changes are significantly amplified by the concurrent large North

Atlantic sea-ice extent changes. The northern hemisphere sea-ice area decreases from a

peak stadial value of 16.0 × 106 km2 to a peak interstadial value of 11.6 × 106 km2. This

change happens mainly in the North Atlantic.

The zonally averaged SAT in the southern high latitudes also shows a periodic, delayed

out-of-phase behavior as compared with the North Atlantic SAT (Fig. 1d). However, the



amplitude of these oscillations is small (around 0.5 oC) because the oceanic heat transport

change is small (see the discussion for Fig. 3). Note that the atmospheric CO2 level is

prescribed. If a variable atmospheric CO2 were used in the model, these Southern Ocean

oscillations would perhaps be larger. Observations show CO2 rises of less than 10 ppm

during the cool phase of less pronounced D-O oscillations and up to 20 ppm rises during

long lasting cool phases of large oscillations [Stauffer et al., 1998; Stocker and Marchal,

2000]. However, the existence of the SAT changes over the Southern Ocean corresponding

to less pronounced D-O oscillations in the proxy data is less evident [Blunier and Brook,

2001; Clark et al., 2002;].

The streamfunction in the Atlantic Ocean basin is shown in Fig. 2a and 2b for both

the strong and weak modes; these plots are snapshot outputs at model year 10,430 and

10,660, respectively. Deep water forms in the subpolar region of the North Atlantic in both

the strong and weak modes, with a shallower and less vigorous overturning circulation

for the weak mode. Although some modelling results suggest a southward shift of the

deep water convection site [Ganopolski et al., 1998; Knorr and Lowmann, 2003], proxy

data support the concept that the deep water formed more or less continuously in the

same subpolar region of the North Atlantic during the last glacial period [Vidal et al.,

1998; Dokken and Jansen, 1999; Sarnthein et al., 2000; Clark et al., 2002]. The UVic

intermediate complexity climate model also simulates the deep water formation in the

subpolar region under LGM boundary conditions [Weaver et al., 1998]. Our subsequent

sensitivity experiments show that northern sea-ice brine rejection is responsible for the

continuous maintenance of the subpolar deep water formation in this sea-ice covered region



(see Fig. 6c and the related discussion), which supports the hypothesis based on proxy

data reconstructions [Vidal et al., 1998; Dokken and Jansen, 1999].

The annual mean oceanic heat transport contrast between the strong and the weak

mode is much larger in the Atlantic (Fig. 3a) than in the Indo-Pacific (Fig. 3b). At

40o N in the Atlantic the northward heat transport for the strong mode is 0.4 PW larger

than that for the weak mode, while in the Indo-Pacific, the heat transport difference

is small in the northern middle and high latitudes. This demonstrates again that the

large northern climate change during the oscillations is mainly induced by the Atlantic

MOC mode switch. The cooling and warming that occur over the Pacific are principally

induced by atmospheric heat transports. In the southern hemisphere, the change of the

heat transport in the Atlantic is opposite to and larger than the change in the Indo-Pacific.

Consequently, the zonal integral of the oceanic southward heat transport is slightly larger

for the weak mode than for the strong mode. The southern high latitude SAT thus has

a cooling (warming) corresponding to the warming (cooling) in the northern hemisphere.

The opposite changes of the oceanic heat transport in the Atlantic and Indo-Pacific lead to

a small difference in the global southward heat transport between the two Atlantic MOC

modes, which is one of the reasons for a small simulated zonally averaged SAT change at

62.5o S (Fig. 1d). This also indicates a complicated spatial pattern for the global effect

of the Atlantic MOC mode switch.

In Gildor and Tziperman [2003] and Kaspi et al. [2004], sea ice switch-like behavior

driven by a small Atlantic MOC variability is responsible for large temperature changes

during a glacial period. Although the Atlantic MOC jumps to about 40 Sv in this study,



the sea ice extent (figure not shown) and temperature (see Fig. 1c) corresponding to this

state are just slightly different from those corresponding to the Atlantic MOC state shown

in Fig. 2a. This is because the oceanic heat transport at the peak state of the Atlantic

MOC is just slightly larger than that at the state as shown in Fig. 2a. Fig. 3a shows that

the maximum oceanic heat transport change between the two modes is 0.4 PW at 40o N,

which is close to the maximum heat transport change of 0.3 PW in Kaspi et al. [2004].

4. Modulation by a changing background climate

Under the prescribed standard PMIP LGM boundary conditions [Joussaume and Tay-

lor, 2000], the global annual mean SAT drop relative to the present-day SAT is 4.1 oC

averaged over one cycle in our model. This drop is at the low end of the range found by

most PMIP1 investigators (i.e., 4-6 oC) [Joussaume and Taylor, 2000] and smaller than

that obtained in Ganopolski et al. [1998] (6.2 oC). Note, however, that some radiative

forcings are missing in the MPM-2, such as CH4 and dust, and that the increased con-

tinental area due to sea level drop [Broccoli, 2000], which affects the surface albedo, is

neglected. At this stage, it is difficult to include all radiative forcings with enough relia-

bility and accuracy. In order to investigate how the millennial oscillations are modulated

by the changing glacial background climate state, in the next experiment the atmospheric

CO2 level is gradually decreased from 240 ppm at a rate of 1 ppm per thousand years; at

the same time, we maintain the other conditions and the white noise freshwater forcing

as in the control experiment. We note that the modulation effects of a changing glacial

background climate on millennial oscillations can also be obtained by 1) prescribing the

evolution of ice sheets between 60 and 30 kyr BP [Claussen et al., 2003], 2) imposing a



North Atlantic freshwater forcing which increases with ice volume [Sima et al., 2004], and

3) introducing a variable orbital forcing [Olsen et al., 2005]. The model is integrated for

80 thousand years after the spin-up phase. The final CO2 level of 160 ppm is thus reached

at the end of the model run. By changing the CO2 level from 240 to 160 ppm, a gradual

but fairly large cooling of the background climate is realized. The global annual mean

SAT changes from 10.9 oC (a 3.2 oC cooling relative to the present day) to 8.3 oC (a 5.8

oC cooling).

The proxy data from the NGRIP (North Greenland Ice Core Project) ice core [NGICP,

2004] (Fig. 4a) suggest that millennial oscillations are suppressed during the early part of

the glacial and at the LGM. A much longer ocean sediment core record [McManus et al.,

1999] confirms that this type of behavior was also the case for the past 0.5 million years,

which covers the past five glacial-interglacial cycles. These observations thus strongly

suggest that a moderate global cooling favors the millennial oscillations. Also, the glacial

millennial oscillations are not of uniform structure: they have a much longer period during

late stage 5 and early stage 3 (warmer background climate) and during stages 4 and 2

(colder background climate) (see Fig. 4a). When the background climate is warmer, the

interstadial phases are much longer, whereas when the background climate is colder, the

stadial phases are much longer [Bond et al., 1999; Schulz, 2002]. During a ’cooling’ Bond

cycle, which has smaller ice sheets and higher atmospheric CO2 level [Stauffer et al., 1998;

Stocker and Marchal, 2000] at the beginning, the interstadial phase is longer for the first

millennial cycle; the stadial phase is longer and the interstadial phase is much shorter for

the last millennial cycle [Bond et al., 1993]. We also note that the long stadial for the



last cycle is usually accompanied by a massive iceberg discharge event (Heinrich event)

several hundred years after the rapid cooling [Bond and Lotti, 1995].

Fig. 4b shows that the Atlantic MOC gradually slows down when the global annual

mean SAT drops (see Fig. 4c). Once a threshold value T1 of the SAT (T1 = 10.5 oC

when CO2 is 223.4 ppm) is reached, millennial oscillations appear. When another SAT

threshold value T2 (T2 = 9.7 oC when CO2 is 190.7 ppm) is reached, the millennial

oscillations disappear. Thus in our simulation the millennial oscillations exist for the

SAT in the range of 10.5 to 9.7 oC and for CO2 in the range of 223.4 to 190.7 ppm under

prescribed LGM ice sheets. The period is longer when the SAT drops just beyond the first

threshold value T1 and each oscillation has a relatively long interstadial phase (marked

by red bar in Fig. 4d); the period also becomes longer when the SAT is approaching

the second threshold value T2 and each oscillation has a relatively long stadial phase and

short interstadial phase (marked by blue bar in Fig. 4d). This thermal modulation effect

on the durations of interstadial and stadial phases by the background climate was also

simulated in a simple atmosphere-ocean model [Winton, 1997]. These simulated features

of the glacial millennial oscillations are consistent with paleoclimate data (see Fig. 4a,

Bond et al. [1999] and McManus et al. [1999]).

Fig. 4 demonstrates that a warm climate favors a strong and stable Atlantic MOC and

a cold climate favors a weak and stable Atlantic MOC. An intermediate cold climate forces

the Atlantic MOC to oscillate on a millennial time scale. A relatively long interstadial

is produced by a relatively warm climate. The opposite is true for a relatively cold



climate. Although a massive iceberg discharge event could also produce a long stadial,

the modulation by the background climate should be taken into account.

Sensitivity studies have been also carried out to investigate the role of the imposed

stochastic freshwater forcing in the above runs. In the experiment without the stochastic

freshwater forcing, millennial oscillations still occur. However, the background climate

range in which millennial oscillations occur becomes narrower. By doubling the stochastic

forcing intensity, the background climate range widens. Also the imposed stochastic

freshwater forcing slightly shortens the periods.

Based on the modelling results presented above, we highlight, in Fig. 5, an impor-

tant new mechanisms for the large glacial millennial oscillations. After the early stage

of a glacial cooling period, further global cooling could weaken the Atlantic MOC by

decreasing the equator-to-pole oceanic temperature contrast (and hence the horizontal

oceanic density gradient) in the upper layer [Prange et al., 1997; Wang et al., 2002] and

increasing the vertical stratification in the North Atlantic deep water formation region

[Winton, 1997; Wang et al., 2002]. When the cooling reaches a critical state as indicated

by T1 in Fig. 5, the Atlantic MOC is forced into an unstable state (the green dashed

line) between S and W. This means that this thermal-forcing induced Hopf bifurcation

point is passed and a limit cycle can be obtained under constant external forcing, with the

Atlantic MOC oscillating between the two solid green curves. Thus a periodic oscillation

occurs in the control experiment with constant glacial boundary conditions. When the

background climate is further cooled down, another critical state as indicated by T2 is

reached. The Atlantic MOC evolves into a weak but stable state after this point. The



oscillations hence disappear. In other words, large millennial oscillations can only occur

in an intermediate cold glacial state. This mechanism has been hypothesized in Alley et

al. [1999], Ganopolski and Rahmstorf [2002], and Shaffer et al. [2004].

A strong salinity perturbation or freshwater forcing can also push the Atlantic MOC

into an unstable state after passing a Hopf bifurcation point and hence cause the Atlantic

MOC oscillations [Tziperman et al.. 1994; Sakai and Peltier, 1997; Tziperman, 1997, 2000;

Timmermann, 2003]. But a strong freshwater forcing at the warm (interstadial) phase is

at odds with the evidence that abrupt cooling occurred before the large freshwater forcing

event (Heinrich event) [Bond and Lotti, 1995].

5. The role of sea-ice brine rejection

As mentioned above, the change of sea-ice extent significantly amplifies the warming

and cooling induced by the rapid Atlantic MOC mode switch. This is due to two thermal

features of sea ice: high surface albedo and low heat conductivity (insulation effect). We

next show that sea-ice brine rejection also plays a necessary role in the oscillations.

When sea ice forms, salt is rejected; when sea ice melts, freshwater is released. If sea ice

forms and melts in the same region over a certain period, the integrated freshwater forcing

is canceled out over this period. However, sea ice is generally transported away from a

region of formation to a region of melt. In addition, we note that the area-integrated

freshwater forcing is zero over sea-ice covered regions, except for the case in which the

total sea-ice volume changes due to climate changes. In the MPM-2, the rate of the total

sea-ice volume change is small in terms of freshwater forcing (on the order of 0.001 Sv or

less).



Fig. 6a shows the time series of the integrated freshwater flux over two latitude bands

in the northern North Atlantic. In Fig. 6b, we show the time series of integrated brine

rejection (freshwater equivalent) to the north of 60o N (red curve) and freshwater release

to the south of 60o N (blue curve). Fig. 6b clearly indicates that when the Atlantic MOC

is in a strong mode, both salt rejection to the north of 60o N and freshwater release to the

south of 60o N are very small, consistent with a very small sea-ice extent and small sea-ice

mass in the North Atlantic; when the Atlantic MOC is in a weak mode, both strong brine

rejection and freshwater release occur, which are consistent with a large sea-ice extent

and large sea-ice mass. The large changes of freshwater fluxes in Fig. 6a are mainly due

to the sea-ice brine rejection/freshwater release. The modulated freshwater flux change

can be as large as 0.1 Sv regionally. Note that the salt rejection to the north of 60o N is

nearly balanced by the freshwater release to the south of 60o N (Fig. 6b). Thus the sum

of freshwater fluxes over 45-60o N and 60-75o N in Fig. 6a is almost a constant.

In order to see the effect of sea-ice brine/freshwater release, we designed two other

experiments in which the same glacial boundary conditions and stochastic freshwater

forcing are employed as in the control experiment. In one experiment, starting at model

year 11,000, the brine rejection/freshwater release is fixed at its value at year 10,430. At

year 11,000, the Atlantic MOC is weak (this year is marked by the red vertical dash-dotted

line in Fig. 6c) and both the salt rejection to the north of 60o N and freshwater release

to the south of 60o N are large (Fig. 6b), while at year 10,430, the Atlantic MOC is

strong (see the blue vertical dash-dotted line in Fig. 6c) and both the salt rejection and

freshwater release are very small (Fig. 6b). After a large brine rejection at year 11,000 is



replaced by a small one, the weak mode Atlantic MOC goes to an even weaker and finally

a collapsed state (see the blue curve in Fig. 6c). In another experiment, starting at model

year 11,830, the brine rejection/freshwater release is fixed at its value at year 11,000. In

this experiment, a very large brine rejection/freshwater release value corresponding to a

weak mode of the Atlantic MOC is used to replace a small one corresponding to a strong

mode of the Atlantic MOC. The strong mode is then maintained throughout the end of

the model run, without any switch to a weak mode. The results of these two experiments

demonstrate the important and necessary role of sea-ice brine rejection processes in the

millennial oscillations of the Atlantic MOC.

Sensitivity experiments with no brine rejection taken into account or with very small

sea-ice meridional advection velocity show that only a very weak or collapsed Atlantic

MOC can be obtained under the forcing of the glacial boundary conditions. By increasing

(decreasing) the sea-ice advection velocity, the oscillation period is shortened (lengthened),

further demonstrating the important role of sea-ice brine rejection.

In addition to the forcing by a moderately cold background climate, internal oceanic

feedbacks (for example, see Winton [1997] and Paul and Schulz [2002]) and the imposed

stochastic forcing, we have demonstrated that the process of sea-ice brine rejection during

sea-ice formation also plays a necessary role in the D-O oscillations. Fig. 7 illustrates a

negative feedback loop between the Atlantic MOC, North Atlantic climate and northern

sea-ice brine rejection under an appropriate background climate forcing. If we have a

strong (weak) Atlantic MOC initially, we will have a high (low) SST and SAT in the North

Atlantic and hence a warm (cold) North Atlantic. A warm (cold) North Atlantic leads



to reduced (enhanced) sea-ice extent and thickness, i.e., sea-ice mass. The southward

sea-ice mass transport must be reduced (increased). The sea-ice brine rejection in the

subpolar North Atlantic is consequently small (large). The small (large) brine rejection

does not favor (favors) a strong Atlantic MOC and thus a weak (strong) Atlantic MOC

may appear.

6. Discussion and conclusions

An extensive parameter sensitivity study of the D-O oscillations simulated here is not

presented in this paper, although some sensitivity experiments have been done. It is

well-known that model results are sensitive to the descriptions of mixing processes in

the ocean. Uncertainties in the simulations of global energy and hydrological cycles may

lead to uncertain surface forcing on the ocean circulation. However, the occurrence of a

weakened Atlantic MOC in a cold climate like the LGM appears to be a robust feature that

has been simulated by many climate models [for example see Winton, 1997; Ganopolski

et al., 1998; Kim et al., 2003; Knorr and Lohmann, 2003; Shin et al., 2003]. Our early

version of this model [Wang et al., 2002] also showed a weakened Atlantic MOC for a

very cold climate, and proxy data support this conclusion [Wang et al., 2002]. A recent

proxy data study shows that the LGM Atlantic MOC mode is the same as the Younger

Dryas mode which is a weak mode [Keigwin, 2004]. Therefore, there must be a transition

between the strong mode (corresponding to an interglacial and an early stage of a glacial)

and the weak mode (corresponding to a peak glacial) during a glacial period. We note

that whether or not the large millennial oscillations occur during the transition phase is



model dependent. In this study, we demonstrate that the simulated large glacial millennial

oscillations are sensitive to the process of sea ice salt rejection.

The role of active continental ice sheets is not addressed in this study. There appears to

be little agreement about the response of ice sheet mass balance to millennial oscillations

in current climate models. For example, in one modelling study [Gildor and Tziperman,

2003], cooling leads to the ice mass loss, while in another modelling study [Schmittner

et al., 2002], cooling reduces the ice mass loss. Future work is needed in this aspect.

In another sensitivity experiment with our model, we found that an active vegetation

component plays a minor role in the oscillations.

By employing an EMIC, in this study we have demonstrated that an intermediate

cold glacial background climate could force the Atlantic MOC into an unstable state and

that the Atlantic MOC oscillates on a millennial scale with rapid decadal-scale switches

between the strong mode and the weak mode. During these oscillations, deep water forms

continuously in the subpolar North Atlantic. For the first time, we have also demonstrated

that sea-ice brine rejection/freshwater release and sea-ice transport play an important

role in the maintenance of the weak mode. Our model results capture many important

features of the glacial millennial oscillations. The Atlantic MOC mode switch, along with

the amplification effect by sea ice, causes the large and abrupt northern climate changes.

The southern climate changes are very small and gradual. The longer D-O interstadials

occur in a warmer climate, while a colder climate favors a longer stadial. The net result is

that the period of the oscillations is longer when the background climate is either warmer

or colder and the period is shorter for an intermediate state.



Acknowledgments. This work was supported by an NSERC Discovery Grant and a

CFCAS Project Grant awarded to L. A. Mysak. We thank David Holland and Yi Wang

for their comments on this manuscript. The constructive comments of the reviewers

(Eli Tziperman and the anonymous reviewer), which helped to improve this paper, are

gratefully acknowledged.

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10 11 12 13 14 1510

20

30

40

50

Atla

ntic

MO

C (

Sv) (a)

10 11 12 13 14 1532

34

36

38

Sa

linity

(p

.p.t.) (b)

10 11 12 13 14 15−16

−14

−12

−10

−8

−6

SA

T (o

C)

(c)

10 11 12 13 14 15−12

−11

−10

SA

T (o

C)

Time (kyr)

(d)

Figure 1. The simulated D-O oscillations in the control experiment with the stan-

dard glacial boundary conditions used by PMIP [Joussaume and Taylor, 2000] and the

imposed white noise freshwater forcing. (a) Atlantic MOC (represented by the maximum

streamfunction below the Ekman layer). (b) Sea surface salinity at 57.5o N in the North

Atlantic. (c) The surface air temperature (SAT) at 62.5o N over the North Atlantic. (d)

The SAT at 62.5o S over the Southern Ocean.D R A F T November 3, 2005, 11:05am D R A F T


5

4

3

2

1

0

Dep

th (k

m)

2220

1816141210

86

42

161412

(a)

30oS EQ 30oN 60oN

5

4

3

2

1

0

Dep

th (k

m)

2

12108

6

4

0−1

4

(b)

30oS EQ 30oN 60oN

Latitude

Figure 2. The Atlantic MOC modes in the control experiment. (a) The streamfunction

in the Atlantic at model year 10,430, defined as the strong mode. (b) The streamfunction

in the Atlantic at model year 10,660, defined as the weak mode.



−2

−1

0

1

2P

W(a)

60oS 30oS EQ 30oN 60oN 90oN

−2

−1

0

1

2

PW

(b)


−2

−1

0

1

2

PW

(c)


Latitude

Figure 3. Oceanic heat transports in the control experiment. The oceanic heat

transports in the (a) Atlantic, (b) Indo-Pacific and (c) Global ocean. The red curve is for

the strong mode and blue for the weak mode.



90 80 70 60 50 40 30 20 10

−45

−40

−35

−30

OIS 5OIS 4

OIS 3OIS 2

δ 18O

(per

mil)

Age (kyr BP)

(a)

0 10 20 30 40 50 60 70 8010

20

30

40

50

Atla

ntic

MO

C (S

v) (b)

0 10 20 30 40 50 60 70 808

9

10

11

12

SAT

(o C)

(c)

10 20 30 40 50 60−17

−15

−13

−11

−9

−7

−5

SAT

(o C)

(d)

Time (kyr)

Figure 4. Ice core data and the simulated modulation of D-O oscillations by the

background climate in the experiment with a decreasing atmospheric CO2 level (from 240

to 160 ppm) and other conditions fixed as in the control experiment. (a) The time series

of δ18O from the NGRIP ice core [NGICP, 2004]. (b) The simulated Atlantic MOC. (c)

Global annual mean SAT. (d) SAT at 62.5o N over the North Atlantic. Oxygen isotope

stage (OIS) 5, 4, 3 and 2 are shown in (a) with red representing a relatively low ice volume

and a warm climate and blue a relatively large ice volume and a cold climate [Lambeck

and Chapell, 2001]. The red line and blue line in (d) mark the lengthy interstadials and

the lengthy stadials, respectively; the oscillations associated with these warm and cold

events evidently have larger periods than the oscillations in the middle of the time series.

Note that the time interval is from 10 to 60 kyr for (d).



Figure 5. Schematic diagram of the glacial Atlantic MOC intensity and stability versus

the background climate. T1 = 10.5 oC, T2 = 9.7 oC, S = 22 Sv and W = 16 Sv in this

modelling study.



10 11 12 13 14 15−0.1

0

0.1

0.2

0.3

0.4F

resh

wa

ter

flux

(Sv) (a)

10 11 12 13 14 15−0.2

−0.1

0

0.1

0.2

Brin

e r

eje

ctio

n (

Sv)

(b)

10 11 12 13 14 150

10

20

30

40

50

Atla

ntic

MO

C (

Sv)

Time (kyr)

(c)

Figure 6. (a) The integrated North Atlantic freshwater flux over 45 to 60o N (blue

curve) and 60 to 75o N (red curve) for the control experiment. (’positive’ means that

the ocean obtains freshwater.) (b) The integrated North Atlantic sea-ice brine rejection

(freshwater equivalent) and freshwater release to the south of 60o N (blue curve) and

to the north of 60o N (red curve) for the control experiment. (c) The simulated Atlantic

MOC in the experiment with the brine rejection/freshwater release after model year 11,000

fixed at the value at model year 10,430 (blue curve) and in the experiment with the brine

rejection/freshwater release after model year 11,830 fixed at the value at model year 11,000

(red curve). The black curve in (c) is the one from the control experiment.



Figure 7. The negative feedback loop involving the Atlantic MOC, North Atlantic

climate and sea-ice brine rejection.


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