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Recovery rates of the thermohaline circulation after a freshwater pulse in the North Atlantic vary consider-ably depending on the background climate, as demon-strated in the Community Climate System Model. The recovery is slowest in a Last Glacial Maximum (LGM) climate, fastest in a modern climate, and intermedi-ate between the two in a greenhouse warming (4XCO 2) climate. Previously proposed mechanisms to explain thermohaline circulation stability involving altered hor-izontal freshwater transport in the North Atlantic are consistent with relative recovery rates in the modern and 4XCO 2 climates, but fail to explain the slow LGM recovery. Instead, sea ice expansion inhibits deep-water formation after freshening in the LGM climate by re-ducing heat loss to the atmosphere and providing addi-tional surface freshwater. In addition, anomalous verti-cal freshwater transport across ∼1km depth after fresh-ening is most effective at weakening the stratification in the modern case but is negligible in the LGM case.
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Rates of thermohaline recovery from freshwater pulses in
Modern, Last Glacial Maximum and Greenhouse Warming
C. M. Bitz1, J. C. H. Chiang2, W. Cheng3, and J. J. Barsugli4
Recovery rates of the thermohaline circulation after
a freshwater pulse in the North Atlantic vary consider-
ably depending on the background climate, as demon-
strated in the Community Climate System Model. The
recovery is slowest in a Last Glacial Maximum (LGM)
climate, fastest in a modern climate, and intermedi-
ate between the two in a greenhouse warming (4XCO2)
climate. Previously proposed mechanisms to explain
thermohaline circulation stability involving altered hor-
izontal freshwater transport in the North Atlantic are
consistent with relative recovery rates in the modern
and 4XCO2climates, but fail to explain the slow LGM
recovery. Instead, sea ice expansion inhibits deep-water
formation after freshening in the LGM climate by re-
ducing heat loss to the atmosphere and providing addi-
tional surface freshwater. In addition, anomalous verti-
cal freshwater transport across 1km depth after fresh-
ening is most effective at weakening the stratification in
the modern case but is negligible in the LGM case.
1. Introduction
Shutdown of the thermohaline circulation (THC) in
a modern climate from enhanced freshwater input of
roughly 10 Sv-yr (1 Sv = 106m3s1) to the North
Atlantic is usually only temporary in climate models:
Once the freshwater input returns to normal, the cir-
culation recovers in a few decades to a century at most
[Stouffer et al., 2006]. These models agree roughly with
the duration of climate change that followed meltwater
pulses approximately 8200 year ago, which resulted in
the largest global cooling in the Holocene [Ellison et al.,
2006]. However, meltwater pulses of a similar magni-
tude during the most recent glacial are hypothesized to
have caused the THC to remain collapsed or greatly re-
duced for much longer durations [see e.g., Clark et al.,
1Dept. of Atmospheric Sciences, University of Washington
2Dept. of Geography and Center for Atmospheric
Sciences, University of California, Berkeley
3JISAO/School of Oceanography, University of
4CIRES, University of Colorado
Copyright 2006 by the American Geophysical Union.
2002]. It remains to be seen how anthropogenic climate
change and the anticipated increase in precipitation and
runoff in the northern North Atlantic will influence the
THC response to freshwater pulses, say from Greenland.
Here we investigate the influence of the climate state
on THC recovery following a freshwater addition to the
North Atlantic using the Community Climate System
Model Version 3 (CCSM3). We add the same fresh-
water anomaly to modern, greenhouse warming, and
glacial climates to investigate the mechanisms of THC
recovery in a spectrum of background climates.
2. Model
We impose freshwater perturbations to three nearly
equilibrated background climate cases: a 1,000 yr inte-
gration with “modern” (i.e., 1990s) conditions [Collins
et al., 2006], a 440 yr Last Glacial Maximum LGM (with
ice-sheet topography, ocean bathymetry, orbital config-
uration, and greenhouse gases prescribed for 21 ka)
integration [Otto-Bliesner et al., 2006], and a 4XCO2
stabilization integration [Bryan et al., 2006] where CO2
is initially increased at the rate of 1% per year from
present-day to four times present-day levels and then
held fixed for 300 yr. These particular background
climates, or “controls”, were chosen to sample a wide
range of initial THC strengths and structures and sur-
face heat and freshwater fluxes.
We branched freshwater pulse experiments from each
control by instantaneously freshening the upper 970m
of the North Atlantic and Arctic Oceans from 55-90N,
90W-20E by an average of 2psu (higher at the top and
tapering with depth). This is equivalent to adding 16
Sv-yr of freshwater. Our method is similar to Vellinga
et al. [2002] and contrasts with that used in the inter-
comparison study described by Stouffer et al. [2006],
where freshwater is added over 100 yr to the surface.
The latter distributes the freshening over time, while
the former distributes it in depth. Neither is obvi-
ously more realistic, but instantaneous freshening is
more computationally efficient for studying THC re-
covery. We conducted seven freshened runs for a min-
imum of 20 yr each (three modern, three LGM, and
one 4XCO2). One member from each climate is longer
(75 yr for modern and 4XCO2and 135 yr for LGM) to
capture the decade to century time-scale recovery. Ad-
ditional information about the model and experiments
is given in the auxiliary materials.
3. Results
Freshening in the North Atlantic causes the THC to
immediately collapse with a cessation of North Atlantic
Deep Water (NADW) formation in all seven runs as
shown in Fig. 1. In no case is the collapse permanent,
but the recovery rates are strikingly different depending
on the background climate state. The range of recov-
ery rates within each three-member ensemble is clearly
much smaller. After a complete collapse and even re-
versal in year one, every run exhibits an initial rapid
increase to 2–5 Sv by year 4, and then weakens again in
years 5–10 (see Fig 1). After year 10 the modern THC
rapidly recovers at a rate of 0.2 Sv yr1. In contrast,
the LGM THC recovery rate is about 10 times slower,
or 0.02 Sv yr1. The 4XCO2recover rate is 0.07 Sv
yr1 intermediate to the other two. According to Fig
1, the THC recovery takes about 40 yr in the mod-
ern case and about 70 yr in the 4XCO2case, while
the LGM case appears as though it could take several
centuries. This dependence of the THC recovery rate
on the climate state is the main finding of this study.
In the rest of this section we evaluate several possible
One might expect the recovery rate to depend on
the degree to which the THC is controlled by thermal
versus haline forcing in the different climates. Using a
simple box model, Rahmstorf [1996] argued that the di-
rection of freshwater transport by the meridional over-
turning circulation, WTHC, at 33S indicates whether
the THC at equilibrium is exclusively thermally driven
(southward WTHC) or both thermally and haline driven
(northward WTHC). WTHC is computed from
SSo)dz, (1)
where ¯vis the zonally integrated northward velocity,
Sis the zonal mean salinity, and Sois the reference
salinity, taken here as the global mean salinity. Rahm-
storf [1996] explained that a southward WTHC salini-
fies the Atlantic, so at equilibrium the sum of all other
Atlantic freshwater contributers must brake the THC,
rather than drive it. In this case, a THC shutdown
eliminates a salinity source to the Atlantic, which sus-
tains the shutdown by providing a positive feedback.
If the shutdown is only temporary, then positive feed-
backs slow the recovery. A parallel argument suggests
a northward WTHC yields a negative feedback, which
aids recovery. de Vries and Weber [2005] found this rea-
soning worked well to describe a series of experiments
in an intermediate complexity model where they var-
ied WTHC at 33S by applying surface freshwater flux
Table 1. THC recovery rates and Atlantic northward fresh-
water fluxes
Background Recovery Control Perturbed
Climate Rate WTHC 33S Wgyre 33N
LGM 0.02 Sv yr10.26 Sv -0.47 Sv
Modern 0.2 Sv yr10.03 Sv -0.33 Sv
4XCO20.07 Sv yr1-0.02 Sv -0.30 Sv
2nd decade after freshening
0 20 40 60 80 100
Sinking Rate − Sv
Year Since Freshening
Figure 1. Annual mean thermohaline circulation in-
dex (solid lines = modern, dashed = LGM, dot-dash =
4XCO2). The light gray lines show the mean index for
the corresponding control runs with error bar indicat-
ing plus/minus one standard deviation. The index is the
sinking flux across 1022 m depth from 60-65N, which
emphasizes changes in NADW formation rate, as sug-
gested by Gent [2001]. Positive (negative) values indi-
cate sinking (upwelling). An index of the THC using the
maximum of the Atlantic meridional overturning stream-
function yields the same general conclusions (see auxil-
iary materials).
In our runs, only the 4XCO2control has a southward
WTHC (see Table 1). Because freshening the 4XCO2
climate does not cause a permanent THC collapsed,
other negative feedbacks must be operative. However,
if we assume these other feedbacks are about the same
strength for all three climate states, so that WTHC de-
termines the relative rate of recovery, then the LGM
case should recover fastest, with the modern case next,
and the 4XCO2case last. Because the LGM case is
actually the slowest to recover, we can conclude that
other feedbacks must control the LGM recovery.
Vellinga et al. [2002] argued that the THC recovery
after freshening was controlled by anomalous salt ad-
vection from the North Atlantic subtropical gyre. The
salinity flux by the gyre is typically expressed in terms
of a freshwater flux:
Wgyre =
SoZv0S0dz, (2)
where S0and v0are deviations from the zonal means and
the overbar indicates a zonal integral. After freshening,
Vellinga et al. [2002] found Wgyre at 25N decreased
owing to a southward shift of the Atlantic ITCZ. The
nature of the ITCZ shift after freshening depends on the
mean climate state (Cheng et al, submitted). The ITCZ
shift serves as a negative feedback on the THC, such
that a larger decrease in Wgyre in the subtropical North
Atlantic promotes a faster recovery. However, Wgyre
at 33N in the Atlantic decreases more after freshening
the LGM climate than the modern or 4XCO2climates
(see Table 1), and thus cannot explain the slow LGM
The role of surface buoyancy forcing on NADW for-
mation is another potential factor influencing the THC
recovery rate. The presence of the sea ice strongly al-
ters the surface buoyancy forcing in the subpolar seas,
so we compare the sea ice cover in relation to the sur-
face density in the freshened region for the three control
climate states. Figure 2a shows that in the LGM con-
trol sea ice of 30–80% concentration covers the densist
outcroppings, where NADW forms (also roughly where
the mixed layer reaches its maximum depth, see auxil-
iary material.), suggesting a larger haline influence on
the THC in the LGM climate. This is in agreement
with paleoclimate evidence of brine rejection contribut-
ing to glacial NADW formation [Vidal et al., 1998]. In
contrast, NADW formation occurs south of the sea ice
edge in the modern and 4xCO2 controls (Fig. 2b and
c). Freshening causes the sea ice cover to expand in all
cases, but little sea ice reaches the NADW formation
sites in the modern and 4XCO2climates. This disparity
in sea ice cover among the climates changes the nature
of feedbacks involving the surface fluxes that drive the
THC recovery through deep-water formation.
To quantify the effect of surface buoyancy flux forc-
ing on watermass formation [see Speer and Tziperman,
1992], we first compute components of the surface den-
sity flux, ραFTand ρβFS, where ρis the seawater
density, αand βare seawater thermal and haline con-
traction coefficients, and FTand FSare surface heat
and freshwater fluxes. The density flux components
are integrated over density outcroppings in the Atlantic
north of 55N, south of the Arctic Ocean, and west of
the Barents Sea (henceforth called the “North Atlantic
freshened region”) to arrive at a watermass transfor-
mation function τS,T . Finally we take the derivative,
∂τS,T /∂ρ, to give the surface watermass formation rate
(WFR) as a function of ρ(with units Sv(kg/m3)1) for
each component. As such, the only other sources of
buoyancy forcing to a given density range are from di-
apycnal mixing and boundary flow. If we imagine plac-
ing a bottom “boundary” on isopycnals at 1022 m depth
a) b) c)
5d) LGM−haline
1.026 1.028 1.030
g cm−3
5e) LGM−thermal
1.026 1.028 1.030
g cm−3
5f) Modern−thermal
1.026 1.028 1.030
g cm−3
5g) 4XCO2−thermal
1.026 1.028 1.030
g cm−3
Figure 2. (a-c) Annual mean surface density in the controls in g cm3with the 15% sea ice concentration contour in
the controls (solid lines) and in the second decade after freshening (dashed lines). (d-g) Watermass formation rate in Sv
(kg/m3)1for the North Atlantic freshened region in the control (decade 0) and after freshening (decades 1–5). Haline
components for the modern and 4XCO2are negligible and not shown.
in the sinking region, then the flow across this boundary
(i.e., our index of the THC in Fig. 1) is equal to the
WFR from surface fluxes summed over all isopycnals
that cross this boundary plus the diapycnal mixing into
these isopycnals between the surface and 1 km depth.
Figure 2d-g shows WFR for the North Atlantic
freshened region for all non-negligible thermal and ha-
line components in our runs. The rate that surface
fluxes further densify the heaviest outcroppings (cre-
ating NADW) is indicated by the magnitude of the
rightmost maxima. The thermal component dominates
NADW formation in all three controls and also after
freshening in the modern and 4XCO2climates. Only
the LGM case develops a substantial haline component
after freshening. Freshening shifts the whole pattern
to the left, as outcroppings are made lighter. Then as
the freshwater anomaly erodes, the pattern travel right-
ward, back towards its position in the control. (How-
ever, the physical location of the heaviest outcroppings
and deepest mixed layer sites are not altered by fresh-
ening, see auxiliary material.) In the modern case, sur-
face heat fluxes create dense watermasses at an even
higher rate than in the control within two decades af-
ter freshening, despite the shift in density (Fig. 2f ).
In the 4XCO2case, surface heat fluxes create dense
watermasses after freshening at about 70% of the con-
trol rate (Fig. 2g). In contrast, the LGM case has
greatly reduced dense water formation after freshening
(Fig. 2d and e) owing to expanded sea ice cover over
deep-water outcroppings. Sea ice in the LGM climate
is transported further south and east, where it reduces
heat loss and provides a new source of surface fresh-
water after freshening that destroys outcroppings with
ρ1.029gcm3. There is some compensating increased
brine rejection along with reduced precipitation (dis-
cussed in the auxiliary material) over less dense waters,
with ρ1.028gcm3(Fig. 2d). Such a mechanism
involving sea ice was proposed by Stocker et al. [2001].
The reversal of the effective salt flux over the densest
water outcroppings, counteracting the thermal forcing
after freshening, is the signature of a fundamental dif-
ference between the LGM climate and the modern and
The WFR we computed only depends on surface
buoyancy forcing, yet the THC strength and struc-
ture also depends on the internal ocean stratification.
Among our control simulations, the modern ocean is
the least stable below about 900 m (see Fig. 3a), con-
sistent with its higher sinking rate in Fig. 1 and the
greater sinking depths reached by its NADW compared
to the other controls [Otto-Bliesner et al., 2006; Bryan
et al., 2006]. Comparably greater stratification in the
LGM control is due mainly to salinity, in broad agree-
ment with paleo evidence [Labeyrie, 1992], while in the
4XCO2control it is due mainly to temperature (not
shown). Weaker stratification below the freshened layer
in the modern case allows faster vertical dissipation of
the freshwater anomaly, as can be seen in the potential
density anomaly five decades after freshening in Fig.
3b. The downward anomalous freshwater flux at 1022
m and 1501 m depths is largest in the modern case and
smallest in the LGM case (Fig.3c). Anomalous down-
ward freshwater transport erodes the high stratification
that was imposed by freshening and assists in the THC
recovery. This salinity homogenization after freshening
is most effective in the modern case and least effective
in the LGM case. After five decades, LGM and 4XCO2
anomaly profiles below 1km depth are clearly more sta-
bilizing than the modern anomaly profile (Fig. 3b).
The strong implication is that the originally more sta-
ble density profiles below the surface layer in the LGM
and 4XCO2controls inhibit dissipation of the salinity
anomaly by transport [as noted by Tziperman et al.,
1994] and by convection [as noted by Lenderink and
Haarsma, 1993].
1.025 1.03
Depth − km
g cm−3
−1 −0.5 0
10−3 g cm−3
1 2 3 4 5
0.1 c)
Dedade After Freshening
Figure 3. Density profiles in the (a) controls and (b)
anomalies five decade after freshening and anomalous
vertical freshwater flux (c) through 1022m (black) and
1501m (grey). All quantities are averaged over the North
Atlantic freshened region. The reference salinity used to
compute the freshwater flux in (c) is the average salin-
ity above 3.125 km depth in the region. (Solid line =
modern, dashed = LGM, dot-dash = 4XCO2.)
We also examined the horizontal freshwater transport
into and out of the northern North Atlantic (integrated
over the entire water column in the freshened region),
as well as the surface freshwater flux. The total rate
at which the freshwater anomaly dissipates from the
water column depends little on the climate state (see
auxiliary material). Hu et al. [2007] found the THC
recovers more rapidly when the Bering Strait is open
(compared to closed) in a modern climate in CCSM ver-
sion 2 after adding freshwater at the rate of 1 Sv for 100
yr. Because their freshwater was added at the surface,
the anomaly remains more surface bound. In turn, a
sizeable fraction escapes through the 50 m deep Bering
Strait when it is open. In contrast, when we prescribe
less than 1/6 of the total freshwater anomaly that was
used by Hu et al and we distribute it over the upper
970 m, we find only about 1/16 of this anomaly escapes
through Bering Strait in our modern case. Thus Fresh-
water need not escape through Bering Strait in CCSM3
for the THC recovery to be rapid. It is also noteworthy
that the THC recovery in our 4XCO2freshened case is
much less rapid than in the modern case, despite Bering
Strait being open in both.
4. Summary and Conclusions
We have shown that the recovery rate from a tempo-
rary collapse of the thermohaline circulation after fresh-
ening in the north Atlantic depends on the background
climate state in CCSM3. Clearly, the recovery rate from
a given freshwater pulse in models with present day con-
ditions is not a good indication of the recovery rate from
a pulse of the same magnitude and location in a 4XCO2
or glacial climate.
We analyzed recovery mechanism that are influenced
by horizontal freshwater transport by the mean merid-
ional circulation at 33S and by the subtropical gyre in
the North Atlantic. If these were the dominant feedback
mechanism in our model, then the LGM case would
have the fastest recovery followed by the modern, with
the 4XCO2case being the slowest.
A mechanism that slows the recovery in the LGM
case is the expansion of sea ice in the NADW forma-
tion regions. NADW formation via surface fluxes in the
LGM freshened case alone is inhibited by sea ice ex-
pansion, which effectively reduces surface heat loss and
supplies meltwater over deep-water outcroppings. In
contrast, surface heat loss drives NADW formation in
the modern and 4XCO2climates within two decades af-
ter freshening. Sea ice in poised to play a greater role in
the LGM climate because NADW production sites are
partially covered with sea ice even before freshening.
After freshening the combination of the anomalous
horizontal freshwater transport and surface freshwater
flux in the North Atlantic freshened region are roughly
the same in all three cases. Therefore the dissipation
rate of the total freshwater anomaly through the sur-
face and lateral boundaries cannot explain the differ-
ing recover rates. Instead, we find that the anomalous
downward freshwater transport across 1km depth is
effective at weakening the stratification in the modern
case after freshening, while the same flux is smaller in
the 4XCO2and nearly zero in the LGM case.
Acknowledgments. We gratefully acknowledge support
from the National Science Foundation through grant ATM-
0502204. Computational facilities were provided by the National
Center for Atmospheric Research (NCAR).
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Auxiliary Materials - online only
The Community Climate System Model (CCSM3)
used in this study has an atmosphere component with
approximately 2.8horizontal resolution (T42 spectral
truncation) and 26 vertical levels. The ocean and sea
ice have zonal resolution of 1.125and the meridional
resolution of 0.54, except in the subtropics and trop-
ics where the meridional resolution is finer. CCSM3
is well documented and may be downloaded freely from Present-day forcing is fixed
at 1990 levels for anthropogenic greenhouse gases (CO2
= 355ppm, CH4= 1714ppb, and N2O= 311ppb), ozone,
and aerosols. and has LGM greenhouse gas concen-
trations (CO2= 185ppm, CH4= 350ppb, and N2O=
200ppb) estimated from icecores. Ozone and aerosol
forcing are set to pre-industrial estimates for the LGM
and to 1990 estimates for the modern and 4XCO2sim-
Individual ensemble members were branched from
different times in the controls to sample a range of ini-
tial conditions. Results for the first 20 yr are averaged
over the three ensemble members for the modern and
LGM freshened cases.
Figure A shows a timeseries of the maximum of the
meridional overturning streamfunction in the North At-
lantic, which is an often-used measure of the THC
strength and is an alternative to the index used in
the main text. The maximum overturning occurs at
about 35N at a depth of 850m in the modern con-
trol in CCSM3. The average value of the maximum
overturning for the three controls is approximately dou-
ble the sinking flux across 1022 m depth from 60-65N,
the measure of THC strength used in the main text.
The transient behavior of the THC after freshening ap-
pears qualitatively similar in the two measures. The
recovery rate in the modern climate is fastest, at least
in the first 40 years after freshening, and the recovery
rate in the 4XCO2modern climate is intermediate be-
tween the modern and LGM rates. There are subtle
differences in the details of the recovery as measured by
the two THC indices. Such differences in the various
measures of the THC strength during transient changes
have been described previously and are expected [Gent,
2004]. In fact, these subtle differences are evidence that
more than one measure of the THC should be consid-
ered when analyzing North Atlantic climate. However,
none of the results and conclusions in the main text are
influenced substantially by the choice of THC index.
Figure 2 shows the annual mean mixed layer depths
in the controls and their anomalies after freshening.
This figure shows that the locations with annual mean
mixed layer depths over 200 m roughly coincide with
the locations of the heaviest density outcroppings at
the surface in Fig 2. The mixed layer becomes shal-
lower in most places after freshening, but the location
of the maxima do not shift in any remarkable way. The
maximum mixed layer depths remain underneath the
sea ice before and after freshening in the LGM climate
and south of the sea ice edge before and after freshening
in the modern and 4XCO2climates.
0 20 40 60 80 100
Sinking Rate − Sv
Year Since Freshening
Figure A. Alternative index of the annual mean thermo-
haline circulation shown in Fig 1. Here the index is the
maximum of the Atlantic meridional overturning stream-
function north of 20N and below 418m depth. (Solid line
= modern, dashed = LGM, dot-dash = 4XCO2).
Figure 3a shows horizontal (depth integrated) fresh-
water fluxes into the northern North Atlantic freshened
region across 55N and at the interface of the Arctic
and Barents Seas and the freshwater flux into the sur-
face of the region. These individual components of the
freshwater budget depend on the climate state. How-
ever, the total freshwater dissipation rate (Fig. 3b) is
roughly the same for all three climates, so the total
freshwater dissipation does not provide an explanation
for the large differences in the THC recovery rates. Be-
cause the horizontal transports are integrated vertically,
the important influence of the vertical freshwater flux
(see Fig 3 and the main text) is not apparent. Nonethe-
less the routes of horizontal freshwater dissipation are
worth mentioning. In the first decade after freshening,
Fig. 3a shows that there is anomalous freshwater ex-
port to the north and south in all cases. The anomalous
southward export at 55N, which is comparable in the
first decade in all three cases, is due to the export of the
freshened, near-surface waters by the Labrador Current.
The anomalous northward export into the Arctic and
Barents Sea is at least three times larger in the first
decade in the modern and 4XCO2cases, which have
Bering Strait open, than in the LGM case. Instead, the
LGM has a considerable reduction in precipitation on
average over the freshened region, which is unseen in the
modern and 4XCO2cases. The reduction in precipita-
tion in the LGM climate occurs over a large area, while
there is increase in the convergence of sea ice with even
greater magnitude but localized over deep-water out-
croppings in the LGM (see Fig 2d). In all three cases
the total freshwater dissipation rate is about an order
of magnitude larger in the first decade after freshening
that in subsequent decades.
Figure 2. Annual mean mixed layer depths in the controls (upper panels) and the anomalies two decades after freshening
(lower panels) with the 15% sea ice concentration contour in the controls (solid black line) and two decades after freshening
(dashed, lower panels only).
1 2 3 4 5
Dedade After Freshening
Figure 3. Anomalous freshwater flux into the northern
North Atlantic (a) across 55N on average (black lines),
into the Arctic Ocean and Barents Sea (gray lines), and
downward into the top surface in the region inbetween
(red lines) and (b) the net flux (black plus red minus gray
curves). (solid line = modern, dashed = LGM, dot-dash
= 4XCO2). The transport is computed on the native
model grid, which is not aligned with true latitudes and
longitudes, so the transport across the southern bound-
ary actually spans true latitudes from 53.2–58.6N, de-
pending on the longitude.
Auxiliary Reference
Gent, P. R, and G. Danabasoglu, Heat uptake and
the thermohaline circulation in the Community Climate
System Model, Version 2, J. Clim.,17, 4058–4069, 2004.
... While the timing of freshwater input and DO cycles is still debated, and while freshwater hosing may not be the cause of the AMOC weakening (Barker et al., 2015), these studies provide useful information to study DO cycles. For example, Weber and Drijhout (2007) and Bitz et al. (2007) show that the recovery time of the AMOC is longer under glacial conditions (Last Glacial Maximum, LGM) compared with pre-industrial (PI) conditions. These studies suggest that a larger expansion of sea ice over the North Atlantic in the LGM causes an increase in the recovery time of the AMOC. ...
... These studies suggest that a larger expansion of sea ice over the North Atlantic in the LGM causes an increase in the recovery time of the AMOC. Extensive sea ice covers the original deepwater formation regions and suppresses atmosphereocean heat exchange in the deepwater formation region (Oka et al., 2012(Oka et al., , 2021Sherriff-Tadano and Abe-Ouchi, 2020), which makes it difficult for the AMOC to recover after freshwater hosing is ceased (Bitz et al., 2007;Weber and Drijhout, 2007). In contrast, Gong et al. (2013) compare the recovery time of the AMOC under PI, mid-glacial and LGM conditions in a comprehensive climate model and find that the recovery time is shortest in the mid-glacial case and longest in the PI case. ...
... After the cessation of freshwater hosing and the replacement of the surface wind, the sea surface salinity as well as the subsurface ocean temperature increase gradually in PC-MIS3H_wind and PC-MIS3H_windwater as in PC-MIS3-5aiceH. However, the thick sea ice over the deepwater region prevents the initiation of deepwater formation and maintains the weak AMOC (Loving and Vallis, 2005;Bitz et al., 2007;Oka et al., 2012;Sherriff-Tadano et al., 2021a). This result shows that the cooling effect of the MIS3 ice sheet plays a role in increasing the recovery time of the AMOC by increasing sea ice over the deepwater formation region. ...
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Glacial periods undergo frequent climate shifts between warm interstadials and cold stadials on a millennial timescale. Recent studies show that the duration of these climate modes varies with the background climate; a colder background climate and lower CO2 generally result in a shorter interstadial and a longer stadial through its impact on the Atlantic Meridional Overturning Circulation (AMOC). However, the duration of stadials is shorter during Marine Isotope Stage 3 (MIS3) than during MIS5, despite the colder climate in MIS3, suggesting potential control from other climate factors on the duration of stadials. In this study, we investigate the role of glacial ice sheets. For this purpose, freshwater hosing experiments are conducted with an atmosphere–ocean general circulation model under MIS5a and MIS3 boundary conditions, as well as MIS3 boundary conditions with MIS5a ice sheets. The impact of ice sheet differences on the duration of the stadials is evaluated by comparing recovery times of the AMOC after the freshwater forcing is stopped. These experiments show a slightly shorter recovery time of the AMOC during MIS3 compared with MIS5a, which is consistent with ice core data. We find that larger glacial ice sheets in MIS3 shorten the recovery time. Sensitivity experiments show that stronger surface winds over the North Atlantic shorten the recovery time by increasing the surface salinity and decreasing the sea ice amount in the deepwater formation region, which sets favorable conditions for oceanic convection. In contrast, we also find that surface cooling by larger ice sheets tends to increase the recovery time of the AMOC by increasing the sea ice thickness over the deepwater formation region. Thus, this study suggests that the larger ice sheet during MIS3 compared with MIS5a could have contributed to the shortening of stadials in MIS3, despite the climate being colder than that of MIS5a, because surface wind plays a larger role.
... Results of EMICs (Rahmstorf et al., 2005) and coupled GCMs also suggest that AMOC may have multiple equilibrium states under present or glacial climate conditions (Hawkins et al., 2011;Hu et al., 2012). Experiments with climate models provide evidence that the sensitivity of the AMOC to freshwater perturbation is larger for glacial boundary conditions than for interglacial conditions (Swingedouw et al., 2009) and that the recovery time scale of the AMOC is longer for LGM conditions than for the Holocene (Bitz et al., 2007). ...
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International Ocean Discovery Program Expedition 382, Iceberg Alley and Subantarctic Ice and Ocean Dynamics, investigated the long-term climate history of Antarctica, seeking to understand how polar ice sheets responded to changes in insolation and atmospheric CO2 in the past and how ice sheet evolution influenced global sea level and vice versa. Five sites (U1534–U1538) were drilled east of the Drake Passage: two sites at 53.2°S at the northern edge of the Scotia Sea and three sites at 57.4°–59.4°S in the southern Scotia Sea. We recovered continuously deposited late Neogene sediments to reconstruct the past history and variability in Antarctic Ice Sheet (AIS) mass loss and associated changes in oceanic and atmospheric circulation. The sites from the southern Scotia Sea (Sites U1536–U1538) will be used to study the Neogene flux of icebergs through “Iceberg Alley,” the main pathway along which icebergs calved from the margin of the AIS travel as they move equatorward into the warmer waters of the Antarctic Circumpolar Current (ACC). In particular, sediments from this area will allow us to assess the magnitude of iceberg flux during key times of AIS evolution, including the following: The middle Miocene glacial intensification of the East Antarctic Ice Sheet, The mid-Pliocene warm period, The late Pliocene glacial expansion of the West Antarctic Ice Sheet, The mid-Pleistocene transition (MPT), and The “warm interglacials” and glacial terminations of the last 800 ky. We will use the geochemical provenance of iceberg-rafted detritus and other glacially eroded material to determine regional sources of AIS mass loss. We will also address interhemispheric phasing of ice sheet growth and decay, study the distribution and history of land-based versus marine-based ice sheets around the continent over time, and explore the links between AIS variability and global sea level. By comparing north–south variations across the Scotia Sea between the Pirie Basin (Site U1538) and the Dove Basin (Sites U1536 and U1537), Expedition 382 will also deliver critical information on how climate changes in the Southern Ocean affect ocean circulation through the Drake Passage, meridional overturning in the region, water mass production, ocean–atmosphere CO2 transfer by wind-induced upwelling, sea ice variability, bottom water outflow from the Weddell Sea, Antarctic weathering inputs, and changes in oceanic and atmospheric fronts in the vicinity of the ACC. Comparing changes in dust proxy records between the Scotia Sea and Antarctic ice cores will also provide a detailed reconstruction of changes in the Southern Hemisphere westerlies on millennial and orbital timescales for the last 800 ky. Extending the ocean dust record beyond the last 800 ky will help to evaluate dust-climate couplings since the Pliocene, the potential role of dust in iron fertilization and atmospheric CO2 drawdown during glacials, and whether dust input to Antarctica played a role in the MPT. The principal scientific objective of Subantarctic Front Sites U1534 and U1535 at the northern limit of the Scotia Sea is to reconstruct and understand how intermediate water formation in the southwest Atlantic responds to changes in connectivity between the Atlantic and Pacific basins, the “cold water route.” The Subantarctic Front contourite drift, deposited between 400 and 2000 m water depth on the northern flank of an east–west trending trough off the Chilean continental shelf, is ideally situated to monitor millennial- to orbital-scale variability in the export of Antarctic Intermediate Water beneath the Subantarctic Front. During Expedition 382, we recovered continuously deposited sediments from this drift spanning the late Pleistocene (from ~0.78 Ma to recent) and from the late Pliocene (~3.1–2.6 Ma). These sites are expected to yield a wide array of paleoceanographic records that can be used to interpret past changes in the density structure of the Atlantic sector of the Southern Ocean, track migrations of the Subantarctic Front, and give insights into the role and evolution of the cold water route over significant climate episodes, including the following: The most recent warm interglacials of the late Pleistocene and The intensification of Northern Hemisphere glaciation.
... A similar modeling technique has recently been applied to study the atmospheric impacts of AMOC slowdown under the RCP8.5 warming scenario (Liu et al. 2020). This type of experiments is motivated by the commonly adopted water hosing technique in the context of paleoclimate (Manabe and Stouffer 1995;Zhang and Delworth 2005;Stouffer et al. 2006;Bitz et al. 2007;Cheng et al. 2007;Hu et al. 2008;Kageyama et al. 2013;Liu et al. 2014). Similar experiments with perturbed CO 2 concentration and freshwater flux together have been reported in previous studies but with different focuses (Smith et al. 2014;Wen et al. 2018). ...
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This study examines global patterns of net ocean surface heat flux changes (ΔQnet) under greenhouse warming in an ocean–atmosphere coupled model based on a heat budget decomposition. The regional structure of ΔQnet is primarily shaped by ocean heat divergence changes (ΔOHD): excessive heat is absorbed by higher-latitude oceans (mainly over the North Atlantic and the Southern Ocean), transported equatorward, and stored in lower-latitude oceans with the rest being released to the tropical atmosphere. The overall global pattern of ΔOHD is primarily due to the circulation change and partially compensated by the passive advection effect, except for the Southern Ocean, which requires further investigations for a more definitive attribution. The mechanisms of North Atlantic surface heat uptake are further explored. In another set of global warming simulations, a perturbation of freshwater removal is imposed over the subpolar North Atlantic to largely offset the CO2-induced changes in the local ocean vertical stratification, barotropic gyre, and the Atlantic meridional overturning circulation (AMOC). Results from the freshwater perturbation experiments suggest that a significant portion of the positive ΔQnet over the North Atlantic under greenhouse warming is caused by the Atlantic circulation changes, perhaps mainly by the slowdown of AMOC, while the passive advection effect can contribute to the regional variations of ΔQnet. Our results imply that ocean circulation changes are critical for shaping global warming pattern and thus hydrological cycle changes.
... A number of other mechanisms have been proposed to explain the AMOC variability. The expansion of North Atlantic sea ice has been shown to impact the AMOC through its control over surface fluxes, and by affecting the locations where deep convection occurs (Stocker et al. 2001;Bitz et al. 2007;Sévellec and Fedorov 2015). The link between wind forcing and the AMOC has been explored, in particular in the North Atlantic and the Southern Ocean. ...
North Atlantic meridional density gradients have been identified as a main driver of the Atlantic meridional overturning circulation (AMOC). Due to the cabbeling effect, these density gradients are increasingly dominated by temperature gradients in a warming ocean, and a direct link exists between North Atlantic mean temperature and AMOC strength. This paper quantifies the impact of this mechanism in the Stommel and Gnanadesikan models. Owing to different feedback mechanisms being included, a one degree warming of North Atlantic mean ocean temperature strengthens the AMOC by 3% in the Gnanadesikan model and 8% in the Stommel model. In the Gnanadesikan model that increase is equivalent to a 4% strengthening of Southern Hemisphere winds, and can compensate for a 14% increase in the hydrological cycle. Furthermore, mean temperature strongly controls a freshwater forcing threshold for the strong AMOC state, suggesting that the cabbeling effect needs to be considered to explain past and future AMOC variability.
... To test the sensitivity of the circulation response to the background climate state, we perform an additional experiment where the GSR is introduced in a colder climate. The main goal is to test if the model responds differently under cold climate conditions (when sea ice is present), where the climate system might be more sensitive to perturbations (e.g., Bitz et al. 2007). This is achieved by reducing the solar constant in noridge by 6 W m 22 , from the default S 0 5 1366 W m 22 (warm state) to S 0 5 1360 W m 22 (cold state), mimicking a reduction in atmospheric pCO 2 (equivalent to slightly less than a factor of 2 reduction). ...
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Changes in the geometry of ocean basins have been influential in driving climate change throughout Earth’s history. Here, we focus on the emergence of the Greenland-Scotland Ridge (GSR) and its influence on the ocean state, including large-scale circulation, heat transport, water mass properties and global climate. Using a coupled atmosphere-ocean-sea ice model, we consider the impact of introducing the GSR in an idealized Earth-like geometry, comprising a narrow Atlantic-like basin and a wide Pacific-like basin. Without the GSR, deep-water formation occurs near the North Pole in the Atlantic basin, associated with a deep meridional overturning circulation (MOC). By introducing the GSR, the volume transport across the sill decreases by 64%, and deep convection shifts south of the GSR, dramatically altering the structure of the high-latitude MOC. Due to compensation by the subpolar gyre, the northward ocean heat transport across the GSR only decreases by ∼30%. As in the modern Atlantic ocean, a bidirectional circulation regime is established with warm Atlantic water inflow and a cold dense overflow across the GSR. In sharp contrast to the large changes north of the GSR, the strength of the Atlantic MOC south of the GSR is unaffected. Outside the high-latitudes of the Atlantic basin, the surface climate response is surprisingly small, suggesting that the GSR has little impact on global climate. Our results suggest that caution is required when interpreting paleoproxy and ocean records, which may record large local changes, as indicators of basin-scale changes in the overturning circulation and global climate.
... How sensitive the AMOC at the full glacial was to the freshwater perturbation compared with that at the interglacial (viz., how much freshwater is needed to bring the AMOC to collapse) has been intensively studied but the answer remains unclear. Some proxies indicate the cold on is less sensitive compared with the warm on with variant explanations ( Galaasen et al., 2014;Lynch-Stieglitz et al., 2014), while other proxies ( Mokeddem et al., 2014;Thornalley et al., 2013) and most models ( Bitz et al., 2007;Swingedouw et al., 2009; suggest the opposite. ...
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The notion that the Atlantic Meridional Overturning Circulation (AMOC) can have more than one stable equilibrium emerged in the 1980s as a powerful hypothesis to explain rapid climate variability during the Pleistocene. Ever since, the idea that a temporary perturbation of the AMOC—or a permanent change in its forcing—could trigger an irreversible collapse has remained a reason for concern. Here we review literature on the equilibrium stability of the AMOC and present a synthesis that puts our understanding of past and future AMOC behavior in a unifying framework. This framework is based on concepts from Dynamical Systems Theory, which has proven to be an important tool in interpreting a wide range of model behavior. We conclude that it cannot be ruled out that the AMOC in our current climate is in, or close to, a regime of multiple equilibria. But there is considerable uncertainty in the location of stability thresholds with respect to our current climate state, so we have no credible indications of where our present‐day AMOC is located with respect to thresholds. We conclude by identifying gaps in our knowledge and proposing possible ways forward to address these gaps.
... where S and y represent salinity and meridional current, respectively; S 0 is a reference salinity; the prime and angle bracket denote a deviation from the zonal mean and a zonal integration, respectively. (Rahmstorf 1996;de Vries and Weber 2005;Bitz et al. 2007;Liu and Liu 2013). Positive W gyre at the southern boundary represents freshwater import into the basin, and vice versa. ...
The Pacific meridional overturning circulation (PMOC) is not well known compared to the Atlantic meridional overturning circulation (AMOC), due to its absence today. However, considering PMOC development under different climate conditions shown by proxy and modeling studies, a better understanding of PMOC is appropriate to properly assess the past and future climate change associated with global ocean circulation. Here, the PMOC response to freshwater forcing in the North Atlantic (NA) is investigated using an Earth system model of intermediate complexity under glacial (i.e., Last Glacial Maximum) and interglacial [i.e., preindustrial with/without inflow through Bering Strait (BS)] conditions. The water hosing over NA led to the shutdown of the AMOC, which accompanied an active PMOC except for the preindustrial condition with the opening BS, indicating that the emergence of the PMOC is constrained by the freshwater inflow through the BS, which hinders its destabilization through enhancing ocean stratification. However, the closure of the BS itself could not explain how the sinking motion is maintained in the North Pacific. Here we found that various atmospheric and oceanic processes are involved to sustain the active PMOC. First, an atmospheric teleconnection associated with the collapsed AMOC encouraged the evaporation in the sinking region, causing buoyancy loss at the surface of the North Pacific. Second, the strengthened subpolar gyre transported saltier water northward, enhancing dense water formation. Finally, the vigorous upwelling in the Southern Ocean enabled a consistent mass supply to the sinking region, with the aid of enhanced westerlies.
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Transient simulations of the last deglaciation have been increasingly performed to better understand the processes leading to both the overall deglacial climate trajectory as well as the centennial- to decadal- scale climate variations prevalent during deglaciations. The Paleoclimate Modelling Intercomparison Project (PMIP) has provided a framework for an internationally coordinated effort in simulating the last deglaciation (~20 – 11 ka BP) whilst encompassing a broad range of models. Here, we present a multi-model intercomparison of 17 simulations of the early part of the last deglaciation (~20 – 15 ka BP) from nine different climate models spanning a range of model complexities and uncertain boundary conditions/forcings. A main contrasting element between the simulations is the method by which groups implement freshwater fluxes from the melting ice sheets and how this forcing then impacts ocean circulation and surface climate. We find that the choice of meltwater scenario heavily impacts the deglacial climate evolution, but the response of each model depends largely on the sensitivity of the model to the freshwater forcing as well as to other aspects of the experimental design (e.g., CO2 forcing or ice sheet reconstruction). There is agreement throughout the ensemble that warming begins in the high latitudes associated with increasing insolation and delayed warming in the tropics aligned with the later increases in atmospheric CO2 concentration. The delay in this warming in the tropics is dependent on the timescale of the CO2 reconstruction used by the modelling group. Simulations with freshwater forcings greater than 0.1 Sverdrup (Sv) after 18 ka BP experience delayed warming in the North Atlantic, whereas simulations with smaller freshwater forcings begin deglaciating sooner. All simulations show a strong correlation between North Atlantic temperatures, atmospheric CO2 concentrations, and the AMOC. In simulations with a freshwater forcing greater than 0.1 Sv, North Atlantic temperatures correlate strongly with changes in the AMOC. Simulations with a smaller freshwater forcing show stronger correlations with atmospheric CO2. This indicates that the amount of meltwater strongly controls the climate trajectory of the deglaciation. Comparing multiple simulations run by the same model demonstrate model biases by showing similar surface climate spatial patterns despite the use of different ice sheet reconstructions and/or meltwater flux scenarios. Simulations run with different models, but similar boundary conditions, have provided insight into the sensitivity of individual models to particular forcings, such as the amount freshwater forcing, which has been highly debated in previous studies. This debate has stemmed from the so-called ‘meltwater paradox’ that exists in choosing how much meltwater to input into simulations of the last deglaciation (i.e., large and geologically inconsistent meltwater forcings that successfully produce abrupt climate events versus glaciologically realistic meltwater fluxes that do not). The results of this research highlight how important this decision is.
Paleo proxy records indicate that abrupt changes in thermohaline circulation (THC) were induced by rapid meltwater discharge from retreating ice sheets. Such abrupt changes in the THC have been understood as a hysteresis behavior of nonlinear system. Previous studies, however, primarily focused on a near-static hysteresis under fixed or slowly varying freshwater forcing (FWF), reflecting the equilibrated response of the THC. This study aims to improve the current understanding of transient THC responses under rapidly varying forcing and its dependency on forcing timescales. The results simulated by an Earth system model suggest that the bifurcation is delayed as the forcing timescale is shorter, causing the Atlantic meridional overturning circulation collapse (recovery) to occur at higher (lower) FWF values. The delayed shutdown/recovery occurs because bifurcation is determined not by the FWF value at the time but by the total amount of freshwater remaining over the THC convection region. The remaining freshwater amount is primarily determined by the forcing accumulation (i.e., time-integrated FWF), which is modulated by the freshwater/salt advection by ocean circulations and freshwater flux by the atmospheric hydrological cycle. In general, the latter is overwhelmed by the former. When the forced freshwater amount is the same, the modulation effect is stronger under slowly varying forcing because more time is provided for the feedback processes.
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Plain Language Summary The Last Glacial Maximum (LGM, about 21,000 years before present) was a period when extensive continental ice sheets covered parts of North America and northern Europe. A climate model simulates enhanced multidecadal variability of the Atlantic meridional overturning circulation (AMOC), a major oceanic current system with global climatic impacts, during the LGM relative to modern times. The main driver of multidecadal AMOC variability under LGM conditions is the surface heat flux variability associated with the East Atlantic pattern, a mode of internal atmospheric variability. Under modern conditions, the main driver of the multidecadal AMOC variability is the heat flux variability associated with the North Atlantic Oscillation, another mode of internal atmospheric variability. The difference in AMOC forcing is basically due to the change in topography linked to the extensive ice sheets during the LGM.
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Ocean heat uptake and the thermohaline circulation are analyzed in present-day control, 1% increasing CO2, and doubled CO2 runs of the Community Climate System Model, version 2 (CCSM2). It is concluded that the observed 40-yr trend in the global heat content to 300 m, found by Levitus et al., is somewhat larger than the natural variability in the CCSM2 control run. The observed 40-yr trend in the global heat content down to a depth of 3 km is much closer to trends found in the control run and is not so clearly separated from the natural model variability. It is estimated that, in a 0.7% increasing CO2 scenario that approximates the effect of increasing greenhouse gases between 1958 and 1998, the CCSM2 40-yr trend in the global heat content to 300 m is about the same as the observed value. This gives support for the CCSM2 climate sensitivity, which is 2.2°C.Both the maximum of the meridional overturning streamfunction and the vertical flow across 1-km depth between 60° and 65°N decrease monotonically during the 1% CO2 run. However, the reductions are quite modest, being 3 and 2 Sv, respectively, when CO2 has quadrupled. The reason for this is that the surface potential density in the northern North Atlantic decreases steadily throughout the 1% CO2 run. In the latter part of the doubled CO2 run, the meridional overturning streamfunction recovers in strength back toward its value in the control run, but the deep-water formation rate across 1-km depth between 60° and 65°N remains at 85% of the control run value. The maximum northward heat transport at 22°N is governed by the maximum of the overturning, but the transport poleward of 62°N appears to be independent of the deep-water formation rate.
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The North Atlantic Ocean thermohaline circulation is examined in three experiments using the Climate System Model. They are a control integration for 1870 conditions, and particular emission scenarios for the 20th and 21st centuries. It is found that the strength of the thermohaline circulation does not change significantly over the 21st century. This is in contrast to several other recent studies, which have projected a significant reduction over the 21st century. The reason for the difference is that the Northwest Atlantic becomes warmer and more saline in the Climate System Model. These changes combine to make little change to the surface ocean density in this region, and hence to the rate of deep water formation. Caveats about the Climate System Model and other coupled climate models are then discussed.
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Here we investigate the role of the Bering Strait (BS) in the thermohaline circulation (THC) response to added freshwater forcing (hosing) in the subpolar North Atlantic, through analyzing simulations of a fully coupled climate model with an open and closed BS. Results show that the THC declines similarly with an open and closed BS during hosing. However, the recovery of the THC is delayed by about a century in the closed BS simulation than in the open BS one after the hosing is off. The closed BS prevents the added freshwater being transported from the Atlantic into the Pacific via the Arctic as in the open BS case. Further, the freshwater supply is elevated significantly after the hosing by exporting the freshwater stored in the Arctic during hosing, as sea ice, back to the North Atlantic. This stabilizes the surface stratification there and suppresses the recovery of the deep convection.
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Four high-resolution benthic isotope records from North Atlantic and south Norwegian Sea cores are compared for the past 60 kyr. The benthic records show light δ18O peaks simultaneously to or following the Heinrich events. The amplitude of these light peaks varies between 1‰ for the most northern core ENAM93-21 and 0.35‰ for the deep North Atlantic cores. The high δ18O gradient between the North Atlantic cores and the Pacific core V19-30 suggests a regional influence of δ18O-depleted meltwater to deeper depth during periods of iceberg discharges. The most likely process for transferring δ18O-depleted meltwater to depth is sea ice formation induced by brine formation. During Heinrich events 4 and 3 this mecanism was of major importance (compared to open ocean convection) in the southern Norwegian Sea and at the northern boundaries of the Atlantic, leading to intermediate to deep water formation. Other δ18O excursions appearing in all the records point to the contribution of both brine formation and global changes in the δ18O of seawater associated with the Heinrich event periods.
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The variability of the thermohaline circulation on decadal and centennial time scales is investigated within the context of a simple geostrophic three-layer ocean model with a rectangular closed basin. When slowly varying the atmospheric forcing the model response shows sudden transitions, characterized by local changes in convective activity. Many different equilibria were found. Next, the deep-water formation process was analyzed with a one-dimensional box model. In this box model four different regimes can be identified: convective, nonconvective, periodic, and a regime where both convection and no convection are possible with the same mixed boundary conditions. These regimes were identified in the circulation of the ocean model. -from Authors
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The climate sensitivity of the Community Climate System Model version 3 (CCSM3) is studied for two past climate forcings, the Last Glacial Maximum (LGM) and the mid-Holocene. The LGM, approximately 21 000 yr ago, is a glacial period with large changes in the greenhouse gases, sea level, and ice sheets. The mid-Holocene, approximately 6000 yr ago, occurred during the current interglacial with primary changes in the seasonal solar irradiance. The LGM CCSM3 simulation has a global cooling of 4.5°C compared to preindustrial (PI) conditions with amplification of this cooling at high latitudes and over the continental ice sheets present at LGM. Tropical sea surface temperature (SST) cools by 1.7°C and tropical land temperature cools by 2.6°C on average. Simulations with the CCSM3 slab ocean model suggest that about half of the global cooling is explained by the reduced LGM concentration of atmospheric CO2 (∼50% of present-day concentrations). There is an increase in the Antarctic Circumpolar Current and Antarctic Bottom Water formation, and with increased ocean stratification, somewhat weaker and much shallower North Atlantic Deep Water. The mid-Holocene CCSM3 simulation has a global, annual cooling of less than 0.1°C compared to the PI simulation. Much larger and significant changes occur regionally and seasonally, including a more intense northern African summer monsoon, reduced Arctic sea ice in all months, and weaker ENSO variability.
In this study, the Bering Strait's role in the thermohaline circulation (THC) response to freshwater forcing is investigated through analyzing two sets of simulations using National Center for Atmospheric Research Community Climate System Model version 2. Each set of the simulations includes a control run and a freshwater perturbation (hosing) run. In the hosing runs, an additional 106 m3s-1 freshwater flux is uniformly added into the sub-polar North Atlantic between 50 and 70°N for a century, then this additional freshwater flux is turned off. The Bering Strait (BS) is open in one set of the simulations, but closed in the other set. Results show that when the freshwater forcing is on, the THC behaves similarly in both hosing experiments with a sharp decline in the first 10 years and a slower decline thereafter. However, after the switch off of the freshwater forcing, the recovery of the THC is delayed by about a century in the closed BS run relative to the open BS run. This delayed recovery is caused by the prevention of the closed BS to transport freshwater from the Atlantic into the Pacific via the Arctic, inducing a slower removal of the added freshwater from the sub-polar North Atlantic, thus a delayed recovery of the deep convection there.
The Atlantic thermohaline circulation (THC) is an important part of the Earth's climate system. Previous research has shown large uncertainties in simulating future changes in this critical system. As an activity of WCRP CMIP/PMIP committees, the responses of the THC to idealized freshwater perturbations and the associated climate changes have been intercompared. This intercomparison is among models ranging from the Earth system models of intermediate complexity (EMICs) to the fully coupled atmosphere-ocean general circulation models (AOGCMs) in order to better understand the causes of the wide variations in the THC response. The robustness of particular simulation features has been evaluated across the model results. In response to 0.1 Sv freshwater input in the northern North Atlantic, the multi-model ensemble mean THC weakens by 30% after 100 years. All models simulate some weakening of the THC but no model simulates a shutdown. The multi-model ensemble indicates that the surface air temperature could present a complex anomaly pattern with a cooling south of Greenland and a warming over the Barents and Nordic Seas. The Atlantic ITCZ tends to shift southward. In response to 1 Sv freshwater input, the THC switches off rapidly. A large cooling occurs over the North Atlantic. The annual mean Atlantic ITCZ moves into the Southern Hemisphere. Models disagree in terms of the reversibility of the THC after its shutdown. In general, the EMICs and AOGCMs obtain similar THC responses and climate changes with more pronounced and sharper patterns in the AOGCMs.
In an experiment with the latest version of the Hadley Centre climate model the model response has been analyzed after the thermohaline circulation (THC) in the Atlantic Ocean has been suppressed. The suppression is induced by a strong initial perturbation to the salinity distribution in the upper layer of the northern North Atlantic. The model is then allowed to adjust freely. Salinity gradually increases and deep water formation in the Greenland and Norwegian Seas restarts, later also in the Labrador Sea. The meridional overturning recovers after about 120 yr. In the first few decades when the overturning is very weak surface air temperature is dominated by cooling of much of the Northern Hemisphere and weak warming of the Southern Hemisphere, leading to maximum global cooling of 0.9°C. The disruption to the atmosphere's radiation balance results in a downward flux anomaly at the top of the atmosphere, maximally 0.55 W m-2 in the first decade then decreasing with the THC recovery. The processes responsible for the recovery of the THC is examined in detail. In future model development this will help to reduce uncertainty in modeling THC stability. The recovery is driven by coupled ocean-atmosphere response. Northward salt transport by the subtropical gyre is crucial to the recovery of salinity in the North Atlantic. A southward shift of the ITCZ creates positive salinity anomalies in the tropical North Atlantic. This supports the northward salt transport by the subtropical gyre that helps to restart deep water formation and the THC.
Which physical processes effectively determine the stabil-ity regime of the Atlantic meridional overturning circulation (MOC) is not yet fully understood. We investigate the role of the oceanic freshwater transport into the basin, employing a coupled atmosphere/ocean/sea-ice model of intermediate complexity. By modifying the longitudinal variation of sur-face salinities near the southern border at 33 • S, the amount of salt flowing out of the Atlantic via the Brazil Current can be regulated. In turn, this will influence whether the MOC exports or imports salt water. The latter is associated with a basin-scale salinity-overturning feedback, which can be ei-ther positive or negative. Pulse experiments strongly sug-gest that its sign determines the existence of a monostable or bistable regime in our model.