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JOURNAL OF CLIMATEVOLUME 22

North Pacific Climate Response to Freshwater Forcing in the Subarctic North Atlantic:
Oceanic and Atmospheric Pathways
YUKO M. OKUMURA, CLARA DESER, AND AIXUE HU
Climate and Global Dynamics Division, National Center for Atmospheric Research, Boulder, Colorado

AXEL TIMMERMANN AND SHANG-PING XIEInternational Pacific Research Center, University of Hawaii at Manoa, Honolulu, Hawaii(Manuscript received 6 March 2008, in final form 1 August 2008)
ABSTRACT
Sudden changes of the Atlantic meridional overturning circulation (AMOC) are believed to have caused
large, abrupt climate changes over many parts of the globe during the last glacial and deglacial period. This
study investigates the mechanisms by which a large freshwater input to the subarctic North Atlantic and an
attendant rapid weakening of the AMOC influence North Pacific climate by analyzing four different ocean–
atmosphere coupled general circulation models (GCMs) under present-day or preindustrial boundary conditions. When the coupled GCMs are forced with a 1-Sv (Sv [ 106 m3 s21) freshwater flux anomaly in thesubarctic North Atlantic, the AMOC nearly shuts down and the North Atlantic cools significantly. The South
Atlantic warms slightly, shifting the Atlantic intertropical convergence zone southward. In addition to thisAtlantic ocean–atmosphere response, all of the models exhibit cooling of the North Pacific, especially along
the oceanic frontal zone, consistent with paleoclimate reconstructions. The models also show deepening of
the wintertime Aleutian low.
Detailed analysis of one coupled GCM identifies both oceanic and atmospheric pathways from the Atlantic
to the North Pacific. The oceanic teleconnection contributes a large part of the North Pacific cooling: thefreshwater input to the North Atlantic raises sea level in the Arctic Ocean and reverses the Bering Strait
throughflow, transporting colder, fresher water from the Arctic Ocean into the North Pacific. When the
Bering Strait is closed, the cooling is greatly reduced, while the Aleutian low response is enhanced. Tropical
SST anomalies in both the Atlantic and Pacific are found to be important for the equivalent barotropic
response of the Aleutian low during boreal winter. The atmospheric bridge from the tropical North Atlantic
is particularly important and quite sensitive to the mean state, which is poorly simulated in many coupled
GCMs. The enhanced Aleutian low, in turn, cools the North Pacific by increasing surface heat fluxes andsouthward Ekman transport. The closure of the Bering Strait during the last glacial period suggests that theatmospheric bridge from the tropics and air–sea interaction in the North Pacific played a crucial role in the
AMOC–North Pacific teleconnection.

1. Introduction
In contrast to the relatively stable climate of the last
10 kyr, the recent glacial and deglacial period (;10–110
kyr ago) was marked by large, rapid climate variations
on millennial time scales, as revealed by isotope analyses of Greenland ice cores (Dansgaard et al. 1993; see
Weart 2003 for a historical account) and other proxies.
Surface air temperature in Greenland repeatedly unCorresponding author address: Dr. Yuko M. Okumura, Climate
and Global Dynamics Division, National Center for AtmosphericResearch, P.O. Box 3000, Boulder, CO 80307.
E-mail: yukoo@ucar.edu
DOI: 10.1175/2008JCLI2511.1
Ó 2009 American Meteorological Society

derwent periods of abrupt warming followed by more
gradual cooling, with amplitudes of 58–108C. Sometemperature jumps occurred in less than a century oreven a decade. Some of the cold periods in Greenland
are also marked by layers of ice-rafted debris in the
North Atlantic Ocean sediment cores that indicate increased discharge of continental ice into the ocean(Heinrich 1988; Bond and Lotti 1995).
Growing evidence from subsequent studies showsthat these rapid climate changes were not confined to
the North Atlantic but spread over many parts of theworld. During the Greenland cold phases, the Atlanticintertropical convergence zone (ITCZ) was displaced

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OKUMURA ET AL.

southward (Peterson et al. 2000; Wang et al. 2004;
Schmidt et al. 2006), Indian and Asian summer monsoons were weakened (Wang et al. 2001; Altabet et al.
2002; Ivanochko et al. 2005), and the eastern equatorial
Pacific was colder (Koutavas et al. 2002; Kienast et al.
2006) with the Pacific ITCZ displaced southward(Leduc et al. 2007) and the Walker circulation weakened (Stott et al. 2002). Outside the tropics, large SST
variations closely related to the Greenland records arereported in different parts of the North Pacific (Thunelland Mortyn 1995; Hendy and Kennett 1999; Kienast and
McKay 2001) and the Okhotsk Sea (Harada et al. 2006).
Enhanced ventilation during the cold periods is alsosuggested in the analyses of ocean sediments from theSanta Barbara Basin (Behl and Kennett 1996) and thenortheast Pacific (Lund and Mix 1998).
The Atlantic meridional overturning circulation(AMOC; a.k.a. thermohaline circulation) is considered
to have played a central role in the millennial climate
variations during the last glacial period. In present-day
climate, the ocean surface current carries warm, saline
tropical water toward the extratropical North Atlantic
where the water cools, sinks, and returns southward at
depth, transporting a large amount of heat into highlatitudes. Models with various complexity suggest that
the AMOC is sensitive to a change in extratropicalsurface buoyancy, particularly salinity fluxes (e.g.,
Stommel 1961; Bryan 1986; Manabe and Stouffer 1988;
Mikolajewicz and Maier-Reimer 1994; Rahmstorf 1995;
Schiller et al. 1997; Yin and Stouffer 2007). A large
freshwater flux anomaly due to melting icebergs or
glacial discharge increases the surface buoyancy andcan inhibit deep-water formation in the North Atlantic.
As a result, the AMOC shuts down temporally orreaches another state of equilibrium. Rooth (1982) and
Broecker et al. (1985) are among the first studies topropose the connection between rapid climate changes
and the AMOC. Their idea has been supported by anumber of paleoclimate studies indicating that reduced
production of North Atlantic Deep Water was accompanied by stadial conditions in the North Atlantic region (Boyle and Keigwin 1987; Oppo and Lehman 1995;
McManus et al. 2004).
While it is straightforward to understand the impactof an AMOC shutdown on surface temperatures in theextratropical North Atlantic via oceanic heat transport(Krebs and Timmermann 2007), it is not fully understood how climate in distant regions can be affected.
Ocean models successfully simulate the North Atlantic
cooling when the AMOC is forced to collapse, but surface temperature anomalies are mostly confined to theNorth Atlantic (e.g., Mikolajewicz and Maier-Reimer1994), suggesting an important role of the atmospheric

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circulation in conveying an AMOC change beyond theAtlantic basin. Indeed, the pioneering coupled GCM
study by Manabe and Stouffer (1988) showed that theclimatic impact of an AMOC shutdown can be global. In
response to a shutdown of the AMOC, the extratropicalNorth Atlantic cools up to 278C and the South Atlantic
warms slightly. Associated with this north–south dipoleof temperature anomalies, the Atlantic ITCZ shifts southward, consistent with paleoclimate records. The cooling is
not confined to the North Atlantic but spread over theentire Northern Hemisphere. These features are also
robust in subsequent coupled GCM studies, although thedegree of AMOC stability is strongly model dependent(Schiller et al. 1997; Dong and Sutton 2002; Vellinga andWood 2002; Zhang and Delworth 2005; Stouffer et al.
2006; Timmermann et al. 2007; Wu et al. 2008).
For a quantitative comparison of the climate response
to freshwater water perturbations and resultant AMOC
changes among different coupled GCMs, coordinated‘‘water hosing’’ experiments were conducted as part of
the Coupled Model Intercomparison Project/Paleoclimate
Modeling Intercomparison Project (CMIP/PMIP) (Stouffer
et al. 2006). In response to a 1 Sv (Sv [ 106 m3 s21)
freshwater forcing in the subarctic North Atlantic, theAMOC slows down rapidly in all of the models andassociated climatic anomalies exhibit a large degree ofsimilarity, especially in the Atlantic basin: a north–south
dipole pattern of surface temperature anomalies andsouthward displacement of the ITCZ. While a change inoceanic heat transport is fundamental for this interhemispheric surface temperature dipole (Yang 1999;
Knutti et al. 2004; Wu et al. 2008), atmospheric teleconnections from the extratropics and tropical air–sea
interaction also play an important role (Chiang and Bitz
2005; Krebs and Timmermann 2007; Chiang et al. 2008).
The AMOC-induced climate anomalies in the Atlantic further influence the other tropical basins throughatmospheric teleconnections. In response to a substantial weakening of the AMOC, the eastern Pacific ITCZ
is displaced southward (Dong and Sutton 2002; Vellinga
and Wood 2002; Zhang and Delworth 2005; Timmermann
et al. 2007; Wu et al. 2008). The tropical North Atlantic
cooling induces anomalous high pressure that extends
westward into the eastern tropical North Pacific, thereby
intensifying northeasterly trades across the Central
American isthmus. These cross-isthmus winds advect
anomalously cold and dry air from the Atlantic anddecrease SST in the eastern tropical North Pacific byenhancing surface turbulent heat fluxes, mixing, andupwelling with a distinct seasonality (Zhang and Delworth
2005; Xie et al. 2008). As a result of subsequent air–sea
interactions in the tropical Pacific, some models develop
an El Niño– or La Niña–like SST anomaly pattern with

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JOURNAL OF CLIMATEan equatorial maximum (Dong and Sutton 2002; Wu
et al. 2008), whereas other models favor a meridional
dipole pattern similar to that in the tropical Atlantic(Vellinga and Wood 2002; Zhang and Delworth 2005).
The associated reduction in meridional SST gradient
suppresses the annual cycle in the eastern equatorial Pacific, which may lead to an enhanced El Niño–Southern
Oscillation (ENSO) via nonlinear frequency entrainment (Timmermann et al. 2007). Zhang and Delworth(2005) also suggest that an adjustment of the Walker
circulation leads to weakened Indian and Asian monsoons. These tropical responses are in general agreement with the paleoclimate studies mentioned earlier.
Beyond the Atlantic basin, coupled GCMs also exhibita robust response over the North Pacific (Timmermann
et al. 2007). When the AMOC nearly shuts down, theNorth Pacific cools throughout the basin with maximum
amplitude of 38–58C along the oceanic frontal zone, andwesterly winds intensify, associated with deepeningof the wintertime Aleutian low. This North Pacific response appears in many earlier coupled GCM studies(e.g., Manabe and Stouffer 1988; Mikolajewicz et al.
1997; Schiller et al. 1997; Dong and Sutton 2002; Vellinga
and Wood 2002; Zhang and Delworth 2005) and agrees
with paleoclimate records, suggesting a strong connection between the AMOC and North Pacific climate.
However, few studies address the mechanisms of thisinterbasin teleconnection. Manabe and Stouffer (1988)
explain the North Pacific cooling by thermal advectionof the North Atlantic temperature anomalies by prevailing westerly winds. The advected cold air enhances
upward surface heat fluxes in the North Pacific and theresultant SST cooling, in turn, intensifies the westerly
winds by thermal wind balance. If thermal advection isthe dominant mechanism for the North Pacific cooling,
it is expected that the magnitude of surface temperature
anomalies decreases toward the east across the Eurasiancontinent and North Pacific. In most of the models, however, the North Pacific cooling is larger than the surface airtemperature anomalies in the Far East by a few degrees.
The deepened Aleutian low appears to be the key tounderstand this North Pacific response in addition tothermal advection: anomalous northerly winds advect
colder Arctic air in the western basin, while enhanced
westerlies increase upward surface heat fluxes andsouthward Ekman transport in the central basin; furthermore, the positive wind stress curl may act to shift
the oceanic front to the south, all intensifying the SST
cooling. In the most comprehensive study on this subject
to date, Mikolajewicz et al. (1997) show that the atmospheric circulation anomaly in their water-hosing experiment projects strongly on the leading internal modeover the North Pacific sector, which represents changes

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in the strength and position of the Aleutian low. Theysuggest that the Aleutian low response is a combinationof remote influence from the Atlantic and local air–sea
interactions in the North Pacific. Using a ‘‘partial coupling’’ technique in a coupled GCM, Wu et al. (2008)
argue that the extratropical North Atlantic cooling is
the major driver for the North Pacific response. In theirmodel, the strong North Atlantic cooling accelerates
westerly winds over the North Pacific by triggering a
positive phase of the annular mode (Thompson andWallace 2000). However, the low pressure and westerly
wind anomalies associated with the annular mode patterns are limited to north of the subarctic North Pacific,
and thus this mechanism does not explain the stationary
wave feature of the Aleutian low that extends into the
midlatitude North Pacific. The large tropical heating
anomalies associated with changes in the Atlantic andPacific ITCZs are likely to influence the extratropicalatmospheric circulation by exciting barotropic Rossby
waves (Hoskins and Karoly 1981), but their role has notbeen investigated in association with the AMOC change
explicitly.
Given the rapid rate of North Pacific SST changes
associated with the North Atlantic climate anomalies(Thunell and Mortyn 1995; Hendy and Kennett 1999;
Kienast and McKay 2001), atmospheric teleconnections
are likely to have played a crucial role in the AMOC–
North Pacific linkage during the last glacial period.
Oceanic teleconnections, however, cannot be dismissed.
Hu and Meehl (2005) and Hu et al. (2007, 2008) showthat a part of the freshwater flux anomaly applied to theNorth Atlantic is transported into the North Pacificthrough the Bering Strait by reversing the hydraulicallycontrolled throughflow. This inflow of cold and fresh
Arctic water may have a significant impact on the heatand salinity balance in the North Pacific. With a depth of
30–50 m and a width of 85–150 km, the Bering Strait wasclosed during the last glacial period owing to snow/ice
accumulation over the continents and resultant decrease in sea level (Chappell 2002; Siddall et al. 2003).
Therefore, such an oceanic teleconnection is not relevant to the millennial-scale climate variations during thelast glacial period, but its impact needs to be evaluatedin models since most of previous coupled GCM studies
use the present-day boundary condition (i.e., an open
Bering Strait). The open Bering Strait experiment may
also be applicable to the short-lived cold event that
occurred approximately 8.2 kyr ago during the Holocene, although the associated freshwater discharge of
5–10 Sv lasted less than a year (Clarke et al. 2003).
Slow adjustment of the global thermohaline circulationand transmission of oceanic waves influence the North
Pacific as well (e.g., Manabe and Stouffer 1997;

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OKUMURA ET AL.

Mikolajewicz et al. 1997; Timmermann et al. 2005a). It
is suggested that the intermediate/deep-water formation
tends to resume in the North Pacific when the Atlantic
MOC is suppressed (Saenko et al. 2004; Timmermann
et al. 2005b; Krebs and Timmermann 2007; Mikolajewicz
et al. 2007). This interbasin seesaw may explain the outof-phase temperature variations between the subarctic
North Pacific and Atlantic implied by a few paleoclimate studies (Kiefer et al. 2001; Sarnthein et al. 2006).
In an uncoupled ocean model experiment, however, achange in the North Pacific ventilation is reduced to
one-third of that in a coupled experiment (Mikolajewicz
et al. 1997), suggesting again a dominant role of theatmospheric teleconnections that can spread the Atlantic climate change much faster than the oceanic
processes can accomplish alone.
In the present study, we revisit the unresolved issues ofthe linkage between the AMOC and North Pacific climate. How consistent is the North Pacific climate response to a rapid weakening of the AMOC among
coupled GCMs in terms of its magnitude, seasonality,
and temporal evolution? What are the oceanic and atmospheric dynamical processes involved in the teleconnection besides atmospheric thermal advection? How
does a change in the Bering Strait throughflow affect theNorth Pacific? What are the mechanisms for the Aleutian
low response, and how does the enhanced Aleutian low
impact the North Pacific? To approach these problems,
we first analyze the CMIP/PMIP water-hosing experiments with four different coupled GCMs (section 3). By
identifying common features of the North Pacific response and its relation to changes in other regions, theroles of oceanic and atmospheric teleconnections canbe inferred. The hypothesized mechanisms are tested
in additional experiments with one particular coupled
GCM and its atmospheric component (section 4). The
oceanic teleconnection is examined by comparing two
sets of water-hosing experiments, one with an open
Bering Strait and one with a closed Bering Strait (section
4a). The origin of the atmospheric teleconnection is explored in a series of atmospheric GCM experiments
forced by SST anomalies simulated in the coupled GCM
water-hosing experiment (section 4b). The description ofmodels and experiments is provided in section 2. The lastsection summarizes the main results of this study anddiscusses their implications and remaining issues.

2. Models and experimentsa. CMIP/PMIP water-hosing experiments
A subset of the CMIP/PMIP water-hosing experiments is analyzed to identify common, robust features1427

of the AMOC–North Pacific teleconnection among
different models. In these coordinated experiments,
coupled GCMs are forced with a time-invariant freshwater flux anomaly in the subarctic North Atlantic (508–708N) for 100 years. The freshwater flux is evenly distributed over the broad forcing area to account formodel-to-model variations in the location of deep convection. In reality, the freshwater discharge was likely to
be more localized near the continental boundaries,
which may affect the climate response (Saenko et al.
2007). We examine the case of 1 Sv hosing for which the
AMOC nearly shuts down in all the models (Stouffer
et al. 2006). This rate will increase the global-mean sea
level by 9 m over a 100-yr period. Comparable sea levelchanges are estimated for the millennium climate variations during the last glacial period, although they took
place over a longer period (Chappell 2002; Siddall et al.
2003; Clark et al. 2004). The integration is continued
without freshwater forcing for an additional 100 years or
longer, allowing the climate system to recover. In thispaper, we focus on the coupled GCMs’ responses during
the first and last several decades of the 100-yr hosing
period. Although it takes longer than 100 years for the
global thermohaline circulation to equilibrate with the
freshwater forcing, the climate response develops very
rapidly during the first few decades due to atmosphericteleconnections, as will be shown in the following sections. The climatic impact of the freshwater forcing andresultant AMOC change is assessed by comparing thehosing run to climatologies of a 200-yr control simulation for each model. All the models use present-day or
preindustrial boundary conditions and therefore theBering Strait is open, connecting the Atlantic and Pacific through the Arctic Ocean.
Four different coupled GCMs are presented: the
National Center for Atmospheric Research Community
Climate System Model, version 2 (CCSM2) (Kiehl andGent 2004), Max Planck Institute for Meteorology
atmosphere–ocean model (ECHAM5-OM1) (Jungclaus
et al. 2006), Geophysical Fluid Dynamics Laboratory
Climate Model version 2.1 (GFDL CM2.1) (Delworth
et al. 2006), and the Hadley Centre Coupled Model version
3 (HadCM3) (Gordon et al. 2000). These state-of-the-art
climate models comprise ocean–atmosphere GCMs andland surface/sea ice models and maintain stable climates
without flux adjustment in the control simulations. The
ECHAM5-OM1, GFDL CM2.1, HadCM3, and the subsequent version of CCSM2 (CCSM3) have been used
in a recent assessment report by the Intergovernmental
Panel on Climate Change (IPCC), and their simulationsof mean climate are highly ranked among other models(Reichler and Kim 2008), thus well representing thecurrent generation of coupled GCMs. The horizontal

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resolutions of ocean/atmosphere GCMs are 1.18 3(0.278–0.68)/T42 (;2.88) in CCSM2, 38 3 38/T31 (;3.758)
in ECHAM5-OM1, 18 3 (1/38–18)/2.58 3 28 in GFDL
CM2.1, and 1.258 3 1.258/3.758 3 2.58 in HadCM3.
Note that CCSM2 and HadCM3 use virtual salt flux
instead of real water flux in the water-hosing experiments without a global salinity compensation. The next
subsection describes CCSM2 in more detail. Further
information on the other three models and the experimental setting are provided in Stouffer et al. (2006),
Timmermann et al. (2007), and the above references forindividual models.

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ice anomalies averaged over the last 30 years of the 100yr hosing period are then added in different regions: theextratropical North Atlantic (408–75N), tropical Atlantic (258S–308N), entire Atlantic (258S–758N), or entire
Atlantic (258S–758N) and tropical Pacific (158S–158N)
combined. Outside these regions, a linear smoothing isapplied to 108 latitude bands. Each of these CAM2 experiments is integrated for 30 years after a 4-monthspinup. Since year-to-year variability in these experiments is purely due to internal atmospheric dynamics,
each year’s simulation is treated as an independent
sample in tests of statistical significance.

b. CCSM2 and CAM2
The oceanic and atmospheric teleconnection processes are examined through additional experimentswith CCSM2 and its atmospheric component, Community Atmosphere Model version 2 (CAM2) (a model
description is available online at http://www.ccsm.
ucar.edu/models/atm-cam/). CAM2 is a global GCM
with an Eulerian spectral dynamical core and a comprehensive physical package. Both stand-alone CAM2
and CCSM2 simulations use a version with triangulartruncation at T42 (;2.88) and 26 hybrid vertical levels.
The ocean component is based upon the Parallel Ocean
Program (POP) developed at Los Alamos National
Laboratory. POP has 40 vertical levels and a horizontal
resolution of ;1.18 3 (0.278–0.68) with the northernpole displaced into Greenland. CCSM2 also includes
the Community Land Model version 2 (CLM2) and theCommunity Sea Ice Model version 4 (CSIM4), whichhave horizontal grids identical to those of CAM2 andPOP, respectively. These four components are linked bymeans of a flux coupler.
In addition to the CMIP/PMIP experiment describedin the previous subsection, another water-hosing experiment is conducted with CCSM2. This experiment is
identical to the previous one except that Bering Strait isclosed, preventing water mass exchange between the
North Pacific and the Arctic Ocean. The control simulation is also repeated with the strait closed. The AMOC
in these two experiments is analyzed in detail by Hu andMeehl (2005) and Hu et al. (2007, 2008). By comparing
the North Pacific response in the open and closed Bering Strait cases, the roles of oceanic and atmosphericteleconnections can be assessed.
To examine the atmospheric teleconnection processes in more detail, a series of CAM2 simulations isconducted by prescribing SST and sea ice fraction fields
simulated in the closed Bering Strait experiment. First,
CAM2 is forced with monthly climatologies of SST and
sea ice from the 200-yr control run of CCSM2 in theclosed Bering Strait configuration. Monthly SST and sea

3. Coupled GCMs’ response to 1-Sv freshwater
forcing in the subarctic North Atlantic
The AMOC intensity, defined here as the maximum
meridional overturning streamfunction value in the
North Atlantic, ranges from 15 Sv to 23 Sv in the control
runs of the four coupled GCMs (Stouffer et al. 2006;
Timmermann et al. 2007). The AMOC is vigorous andpresents pronounced multidecadal variations in GFDL
CM2.1 and HadCM3 while both the mean intensity andmultidecadal variations are relatively weak in CCSM2
and ECHAM5-OM1. In response to 1 Sv freshwater
forcing in the North Atlantic, all the models exhibit a
rapid weakening of the AMOC during the first few
decades. The AMOC reaches a near-equilibrium, collapsed state within 50 yr with the intensity reduced to
2.5–6 Sv.
Figure 1 shows annual-mean SST, surface windstress, and precipitation anomaly fields averaged overthe last 30 years of the hosing period for the four
coupled GCMs. The Atlantic response is very robust
and consistent among the models. The extratropicalNorth Atlantic SST exhibits strong cooling of 88–128C
that extends from the eastern basin into the subtropical North Atlantic. The South Atlantic warms slightly
with maximum amplitude of 18–38C off the Angola–
Namibia coast. Associated with this interhemispheric
dipole pattern of SST anomalies, anomalous northerly
winds blow across the equator and the Atlantic ITCZ
is displaced toward the warmer Southern Hemisphere.
The Coriolis force acting on these northerlies induces
an anomalous easterly (westerly) component north(south) of the equator, enhancing (reducing) the easterly trade winds and associated latent heat loss from thesea surface. This tropical Atlantic anomaly pattern bears
a close resemblance to decadal–interdecadal variationsin the present-day climate in which thermodynamic
ocean–atmosphere interaction among wind, evaporation, and SST plays an important role (Xie and Carton2004).

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OKUMURA ET AL.

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FIG. 1. Ocean–atmosphere anomalies simulated by four different coupled GCMs in the 1-Sv water-hosing experiment: SST (shading,
8C), surface wind stress (vectors, N m22), and precipitation (green contours . 1 mm day21 and orange contours , 21 mm day21 atintervals of 1 mm day21) anomalies are averaged over the last 30 yr of the 100-yr hosing period.

The models also exhibit a robust response in the North
Pacific (Fig. 1). Maximum SST cooling of 38–58C occurs
along 408–458N, near the climatological oceanic frontal
zone in these models. In CCSM2 strong cooling east of
the Kamchatka Peninsula extends southeastward andalong the oceanic front east of 1608E. The GFDL CM2.1
and ECHAM5-OM1 show a sharp meridional SST gradient extending from the east coast of northern Japan
across the basin, indicative of a southward shift of theoceanic front. Indeed, the position of the zero wind stresscurl line is displaced southward in these models (notshown). Westerly winds strengthen over the cooling region in association with a cyclonic circulation anomaly.
The largest wind anomalies are collocated with the
maximum SST cooling in ECHAM5-OM1, GFDL CM2.1,
and HadCM3 but displaced slightly to the north in
CCSM2. These ocean–atmosphere anomalies are reminiscent of the pattern associated with the North Pacificinterdecadal variability (e.g., Trenberth and Hurrell
1994; Mantua et al. 1997; Deser et al. 2004b). However,while intensified southerly winds induce positive temperature anomalies over the Gulf of Alaska in the interdecadal variations, both air temperature and SST
anomalies are weakly negative in the water-hosing ex-

periments. The overall cooling of the Northern Hemisphere is likely to mask the regional warming in theeastern North Pacific in the water-hosing experiments.
The tropical Pacific displays a generally consistent
response despite some differences among the models(Fig. 1). All of the models simulate strengthening of thenortheasterly trades across the Central American isthmus. These enhanced northeasterly winds cool thenortheast tropical Pacific by advecting anomalously
cold, dry air from the North Atlantic and triggering
local air–sea interactions (Xie et al. 2008). The CCSM2,GFDL CM2.1, and HadCM3 develop a meridional dipole pattern of SST anomalies in the eastern basin withanomalous cross-equatorial northerly winds and southward displaced ITCZ, much like that in the tropicalAtlantic. The ECHAM5-OM1, on the other hand,shows La Niña–like SST anomalies that peak at theequator. Precipitation anomalies in this model are muchsmaller compared to the other three models, which may
be due to the lower resolution of its atmospheric component (T31). In the western basin, all models show
suppressed deep convection over the Maritime Continent, suggesting a weakening of the Walker circulation.
Peculiar to HadCM3 is a large precipitation increase in

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the western-central tropical North Pacific associatedwith weak SST warming.
The consistency of the North Pacific response in these
coupled GCMs motivates us to study its mechanisms.
We begin by examining the temporal evolution ofocean–atmosphere anomalies. Figure 2 presents latitude–
time sections of annual-mean SST, surface wind stress,
and precipitation anomalies averaged zonally in the
Atlantic and Pacific and smoothed with a 7-yr runningmean filter. The extratropical North Atlantic cools very
rapidly after the hosing starts, reaching 70%–90% amplitude of the equilibrium response by the end of thefirst decade. The North Pacific cooling is as rapid as theNorth Atlantic cooling, but delayed by one decade, withthe largest change occurring during the second decade.
The rapidity and distinct spatial structure of this remoteresponse imply that dynamical processes are involved
beside atmospheric thermal advection that undergoes
diffusion and damping while traveling across the Eurasian continent. The North Pacific cooling is not preceded by westerly wind anomalies and they tend todevelop simultaneously, except in ECHAM5-OM1 forwhich westerly wind anomalies associated with multidecadal variations appear to lead the SST cooling. A
close examination of SST and wind anomaly maps
shows that the frontal cooling begins before the westerly
wind and cyclonic circulation anomalies become organized in all the models (not shown), suggesting an important role for oceanic processes in initiating the rapid
SST cooling. An ensemble simulation is necessary to
fully assess the nature of the causality between North
Pacific SST and surface wind changes owing to thepresence of atmospheric ‘‘noise’’ associated with ENSO
and internal atmospheric variability in a single model
realization.
The wind anomalies are likely to enhance the North
Pacific cooling through changes in surface heat fluxes
and oceanic currents. These wind anomalies may be remotely forced by changes in tropical atmospheric convection or extratropical surface cooling. Both tropicalAtlantic and Pacific ocean–atmosphere anomalies develop during the second decade, coincident with theNorth Pacific response (Fig. 2). The Atlantic anomalies
remain nearly steady after year 20–30 except in HadCM3
for which SST, surface wind, and precipitation anomalies all gradually increase throughout the hosing period.
The cross-equatorial winds and weak warming in thetropical South Atlantic appear soon after the tropicalNorth Atlantic begins to cool, in support of interhemispheric air–sea interaction via wind–evaporation–SST
feedback (Chiang and Bitz 2005; Chiang et al. 2008).
The tropical Pacific response continues to intensify until
year 50–60 except in ECHAM5-OM1, suggesting a slow

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oceanic adjustment process. There is some indicationthat the North Pacific response intensifies concurrently
with the tropical Pacific response.
The atmospheric circulation response in the North
Pacific displays a distinct seasonal dependence with acyclonic circulation anomaly strongest in boreal winterand absent in summer (Fig. 3). An examination of theindividual monthly responses shows that it typicallyappears in September–October and persists throughMarch–April, the season when a semipermanent low
pressure center develops over the Aleutian Islands (notshown). Thus, the seasonal Aleutian low becomesstronger when the AMOC slows down in the coupled
GCMs. The extratropical atmospheric circulation anomalies are largest during the winter season, not only in the
North Pacific but also in the North Atlantic and theSouthern Hemisphere (Fig. 3). The North Pacific cooling, in contrast, does not show much seasonality withwinter-to-summer differences less than 18C. The tropical Atlantic and Pacific anomalies also exhibit someseasonal variations. The northeasterly winds across theCentral American isthmus are strongest in summer–
early fall when the climatological ITCZ moves over the
Caribbean Sea (Sutton and Hodson 2007; Xie et al.
2008). The zonal bands of precipitation anomalies movenorth and south following the seasonal migration of the
ITCZ in both basins (not shown). The eastern Pacificprecipitation anomalies become most asymmetric about
the equator during winter.

4. Atlantic–North Pacific teleconnection in CCSM2
a. Oceanic teleconnection through the Bering Strait
The North Pacific is connected to the Arctic Ocean
and, ultimately, to the North Atlantic through theBering Strait, a narrow and shallow pass between Alaska
and Siberia. In the present-day climate, the Bering Strait
throughflow is hydraulically controlled and directed
northward, transporting warmer and fresher water fromthe North Pacific into the Arctic Ocean (Woodgate et al.
2005). In the CCSM2 water-hosing experiment, partof the freshwater applied in the North Atlantic flows
into the Arctic Ocean, raising the sea level there andweakening the throughflow. The Bering Strait throughflow is eventually reversed after a decade and about 22%of the total added freshwater is discharged into theNorth Pacific by the end of the 100-yr hosing period (Huand Meehl 2005; Hu et al. 2007, 2008). Approximately
29% of this freshwater is transported as sea ice, increasing sea ice coverage along the Kamchatka Peninsula and in the Sea of Okhotsk by 30%–40% in theannual mean and 60%–70% in winter (not shown).

15 MARCH 2009

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FIG. 2. Time evolution of ocean–atmosphere anomalies simulated by four different coupled GCMs in the 1-Sv
water-hosing experiment. SST (shading, 8C), surface wind stress (vectors, N m22), and precipitation [green
contours . 1 mm day21 and orange contours , 21 mm day21 at intervals of 1 mm day21; minor contours at
10.5 (20.5) mm day21 in dashed green (orange) lines] anomalies are zonally averaged in (left) the Pacific(.258N, 1608E–1508W; ,258N, 1508–1008W) and (right) the Atlantic (608W–108E) and smoothed with a 7-yr
running mean filter.

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JOURNAL OF CLIMATEVOLUME 22

FIG. 3. As in Fig. 1 but averaged for (left) December–February and (right) June–August.

The reversal of the Bering Strait throughflow significantly impacts the heat and salinity balance in the North
Pacific. Figure 4 shows the SST, precipitation, and sea
level pressure anomaly fields in the open and closed
Bering Strait experiments and their differences averaged over the last 30 yr of the hosing period for October–March, the season of the Aleutian low response. Whenthe Bering Strait is closed and no cold freshwater istransported from the Arctic Ocean, the magnitude of
the North Pacific cooling is greatly reduced from 248 to
218C. In the difference map, which indicates the impactof the throughflow changes, negative SST anomalies

15 MARCH 2009

OKUMURA ET AL.

FIG. 4. Ocean–atmosphere anomalies simulated by CCSM2 inthe (top) open and (middle) closed Bering Strait 1-Sv water-hosing
experiments and (bottom) their differences. SST (shading, 8C), sealevel pressure (purple contours at intervals of 1 hPa; negative
contours dashed), and precipitation (green contours . 1 mm day21
and orange contours , 21 mm day21 at intervals of 1 mm day21)
anomalies are averaged in October–March over the last 30 yr of
the 100-yr hosing period.

extend from the eastern subarctic North Pacific to thewest coast of the United States and further into thesubtropical basin. Associated with this subtropicalcooling, the Pacific ITCZ shifts southward, contributing
to more than half of the tropical Pacific precipitation
changes in the open Bering Strait experiment. The Atlantic response is similar between the two experiments,
although the magnitude of the northern cooling is reduced by 18–28C in the open Bering Strait experiment1433

owing to faster removal of the freshwater anomaly andresultant weakening of the upper-ocean stratification. Itshould be noted that the condition of Bering Strait also
affects the mean climates in the control runs (e.g.,
without freshwater forcing; not shown). When Bering
Strait is closed, the North Atlantic becomes slightly
warmer and the North Pacific cools up to 18C with the
Pacific ITCZ shifted southward. About one-third of the
response difference between the two experiments (Fig.
4, bottom panel) arises from the difference in the control runs.
Contrary to the SST response, the wintertime Aleutian low in the closed Bering Strait experiment is surprisingly similar to and even slightly stronger than that in
the open Bering Strait experiment (Fig. 4). Geopotential
height anomalies over this low pressure center display aquasi-barotropic structure with amplitudes increasing
with height (Fig. 5, solid line). As in the open Bering
Strait experiment, it appears only during the winterseason when prevailing mean westerly winds enhanceatmospheric teleconnections (cf. Fig. 12). The stronger
North Pacific cooling in the open Bering Strait experiment compared to the closed one, on the other hand,
induces a localized high pressure anomaly and thus
weakens the Aleutian low response (Fig. 4, bottom
panel). The associated geopotential height anomaly
difference between the open and closed Bering Strait
experiments shows an out-of-phase relationship between
the lower and upper troposphere (Fig. 5, dashed line),
indicative of a direct baroclinic response to diabatic
heating generated by local SST anomalies (Hoskins andKaroly 1981).
In the closed Bering Strait experiment, the rapid
North Pacific cooling seen in all four coupled GCMs isabsent and SST anomalies develop slowly during thefirst couple of decades (Fig. 6). Westerly wind anomalies
are generally weaker in the closed Bering Strait experiment, especially in the band of 508–608N, due to theabsence of a high pressure anomaly induced by strong
North Pacific cooling (Fig. 4). The tropical Pacific precipitation anomalies appear at about the same time in
the two experiments although they do not amplify after
the third decade in the closed Bering Strait experimentunlike the open Bering Strait experiment (Fig. 6).
How does a change in the Bering Strait throughflow
influence North Pacific SSTs in the open Bering Strait
experiment? Fig. 7 displays the time evolution of the
SST and sea surface salinity anomaly fields from year 10
through 21. By the end of the first decade, the North
Pacific is already cooled by 0.58–18C. This cooling islikely due to the weakening of the Bering Strait
throughflow and resultant reduction in the northwardtransport of heat and freshwater. In the closed Bering

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FIG. 5. Vertical profiles of October–March geopotential heightanomalies (m) over the Aleutian low region (308–708N, 1608E–
1408W) simulated by CCSM2 in the closed Bering Strait 1-Sv
water-hosing experiment (solid) and the difference between the
open and closed Bering Strait experiments (dashed).

Strait experiment, in which atmospheric thermal advection and upward surface heat flux anomalies induced
by atmospheric circulation changes are the major causeof North Pacific SST anomalies, basin-scale cooling
does not occur until the second decade (not shown).
Superimposed upon this overall cooling are large negative SST and salinity anomalies along the east coast of
the Kamchatka Peninsula. During the second decade,the anomalies extend southeastward into the interiorof the basin, following the mean ocean current shown inthe bottom panel of Fig. 7. Upon reaching the west coastof the United States near the end of the second decade,the anomalies spread to the north and south along thecoast with the southern branch penetrating into thesubtropical basin. Thus, advection of temperature andsalinity anomalies by the mean ocean current plays animportant role for the rapid North Pacific cooling in the
open Bering Strait experiment. The zonal ocean current
averaged over 358–508N, 1408E–1208W and in the upper
100 m is 4 cm s21, giving the estimated time of 8.4 yr tocross the basin, in agreement with the time scale inferredfrom the visual examination of SST and salinity anomaly
fields. While the advection of cold water directly affects
the mixed layer heat balance, the advection of freshwater indirectly contributes to the cooling by increasing
the upper-ocean stratification and reducing oceanic
convection. To the best of our knowledge, this sensitivity

FIG. 6. As in lhs of Fig. 2 but for anomalies simulated by CCSM2
in the (top) open and (bottom) closed Bering Strait 1-Sv waterhosing experiment; (top) is repeated from Fig. 2.

of North Pacific SSTs to changes in the Bering Strait
throughflow has not been described previously.
The importance of the Bering Strait throughflow for
the North Pacific cooling may be model dependent. By
the end of the 100-yr hosing period, ECHAM5-OM1
and GFDL CM2.1 exhibit SST and sea surface salinity
anomaly patterns very similar to Fig. 7 (see Fig. 4 ofStouffer et al. 2007 for GFDL CM2.1). The Bering
Strait throughflow is, indeed, reversed in the GFDL
CM2.1 water-hosing experiment (J. Yin 2008, personalcommunication). Note that CCSM2 (GFDL CM2.1) isforced by virtual salt (real water) flux and thus the response of the Arctic sea level and the Bering Strait
throughflow does not depend on the form of forcing. In
HadCM3, on the other hand, negative salinity anomalies are mostly confined to the coastal region of thenorthwest Pacific and the strong SST cooling does not15 MARCH 2009

OKUMURA ET AL.

FIG. 7. (top four) Time evolution of SST (shading; 8C) and sea surface
salinity (contours at intervals of 0.5 psu; negative contours dashed andzero contours omitted) anomalies simulated by CCSM2 in the open
Bering Strait 1-Sv water-hosing experiment. Three-year average maps
are shown from year 10–12 through year 19–21. (bottom) Mean ocean
current (m s21) in the control simulation of open Bering Strait experiment, averaged from the surface to 100-m depth over 200 yr.

1435

appear to be associated with the inflow of freshwater.
Among the four coupled GCMs, CCSM2 displays thelargest salinity anomalies over the North Pacific, suggesting that the Bering Strait throughflow may play aless important role in the other models.
In the open Bering Strait experiment, the surface
turbulent (sensible plus latent) heat flux anomaly is
downward over the region of maximum cooling in the
North Pacific (the coast of Kamchatka Peninsula, theSea of Okhotsk, and the oceanic frontal zone along
408N: Fig. 8), confirming that this cooling is largely
caused by oceanic processes. The maximum downward
heat flux anomalies over the oceanic frontal zone areslightly displaced to the south of the maximum SST
cooling owing to the large westerly wind anomalies thatcause upward turbulent heat fluxes between 508 and608N. To the north of the oceanic frontal zone, theseanomalous upward turbulent heat fluxes contribute to acooling of the sea surface. The increased sea ice along
the Kamchatka Peninsula and in the Sea of Okhotsk
induces strong surface cooling and enhances the upward
sensible heat flux to the east through atmospheric
thermal advection. The enhanced westerly winds between 508 and 608N also lower the SSTs by increasing
the southward Ekman heat transport although its impact
is much weaker than the turbulent heat fluxes (Fig. 9).
To the south of the oceanic frontal zone, upward turbulent heat flux anomalies are found in association withweak SST warming (Fig. 4), which is likely due to ananomalous northeastward oceanic current in this region(not shown). In the closed Bering Strait experiment, theheat flux anomaly is, in general, negative and very small
over the North Pacific, indicative of atmospheric forcing. Large negative values along the east coast of theKamchatka Peninsula and northern Japan in this experiment are mostly attributable to sensible heat flux
anomalies due to the outflow of anomalously cold
continental air, while wind-induced latent heat flux becomes more important in the interior basin (not shown).
The stronger Aleutian low response in the closed Bering
Strait experiment (Fig. 4) also increases the southward
Ekman transport, resulting in maximum cooling of 215
W m22 in the western basin (Fig. 9).
In both open and closed Bering Strait experiments, thetropical North Pacific cooling is associated with enhanced
trade winds (Fig. 6) and upward heat flux anomalies (Fig.
8); however, the wind and heat flux anomalies are twiceas large in the open Bering Strait experiment comparedto the closed one. The heat flux anomalies in this region
are mostly due to a wind-induced latent heat component
with longwave and shortwave radiative fluxes nearlycanceling each other (not shown). This suggests that theextratropical and subtropical cooling caused by oceanic

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JOURNAL OF CLIMATEVOLUME 22

FIG. 8. As in Fig. 4 but for the sum of latent and sensible heat flux
anomalies (W m22, downward positive).

FIG. 9. As in Fig. 4 but for the Ekman ocean heat transport
anomalies (W m22).

advection intensifies the trade winds, which in turn cool
the tropical North Pacific by triggering wind–evaporation–SST feedback (Chiang and Bitz 2005; Wu et al.
2008). Off the Pacific coast of the Central American
isthmus, negative heat flux anomalies are of similar
magnitude between the two experiments (Fig. 8), indicating that the North Atlantic cooling and resulting
cross-isthmus wind anomalies play a major role in cooling this region (Xie et al. 2008).

ing from a weakening of the AMOC may affect the atmospheric circulation over the North Pacific throughseveral pathways, including the excitation of preferredcirculation variability patterns and the propagation of
Rossby waves remotely forced by tropical precipitation
changes. We investigate these pathways by forcing
CAM2 with regional SST and sea ice anomalies simulated in the closed Bering Strait experiment in which the
Aleutian low response is less influenced by local SSTs.
Each experiment is integrated for 30 years.
When CAM2 is forced with the SST and sea ice
anomalies in the entire Atlantic (258S–758N) and tropical Pacific (158S–158N) combined, the Aleutian low

b. Atmospheric teleconnection from the tropics
What enhances the wintertime Aleutian low when the
AMOC slows down? SST and sea ice anomalies result-

15 MARCH 2009

OKUMURA ET AL.

1437

FIG. 10. Atmospheric anomalies simulated by CAM2 forced with SST and sea ice anomalies from the CCSM2 closed Bering Strait
experiment. SST and sea ice anomalies are prescribed over (top left) the Atlantic (258S–758N) and tropical Pacific (158S–158N), (top right)
Atlantic (258S–758N), (bottom left) tropical Atlantic (258S–308N), and (bottom right) extratropical North Atlantic (408–758N). Surface
temperature (shading, 8C), sea level pressure (purple contours at intervals of 1 hPa; negative contours dashed), and precipitation (green
contours . 1 mm day21 and orange contours , 21 mm day21 at intervals of 1 mm day21) anomalies are averaged over October–March.

deepens by 3 hPa during October–March, approximately 75% of the amplitude in the closed Bering Strait
experiment and significant at the 98% level (Fig. 10, topleft). The hemispheric pattern and vertical structure ofgeopotential height anomalies bear a close resemblance
to those in the closed Bering Strait experiment (cf. thefirst two columns in Fig. 11), indicating that the Atlantic
and tropical Pacific together force a large part of theatmospheric circulation changes in the Northern Hemisphere. It is noted that the equivalent barotropic vertical structure of the geopotential height response isevident only over the North Pacific.
Which part of the remote forcing is important for the
Aleutian low response? SST and sea ice anomalies inthe extratropical North Atlantic have been shown todrive a hemispheric annular mode response that resembles the leading internal mode of atmospheric circulation variability in an atmospheric GCM (Deseret al. 2004a; Magnusdottir et al. 2004), and Wu et al.
(2008) explain the North Pacific wind anomalies associated with an AMOC shutdown by this mechanism. In
CAM2 the SST cooling and sea ice increase in the ex-

tratropical North Atlantic (408–758N) raises sea levelpressure locally and downstream over Eurasia and theNorth Pacific (Fig. 10, lower right). Thus, the extratropical North Atlantic cooling is not a cause for the
Aleutian low response in the CCSM2 water-hosing experiments.
A southward shift of the Atlantic ITCZ is the mostrobust remote response to an AMOC weakening inboth paleoclimate records and coupled GCMs (Chiang
and Koutavas 2004; Stouffer et al. 2006). Large-scale
shifts in tropical atmospheric convection can dramatically affect the global circulation as revealed by atmospheric teleconnection studies related to ENSO (Trenberth
et al. 1998). When CAM2 is forced with the dipole
pattern of SST anomalies in the tropical Atlantic (258S–
308N), the Atlantic ITCZ is displaced to the south andsea level pressure decreases by 2 hPa in the North Pacific with a barotropic structure in the vertical (significant at the 98% level: Fig. 10, lower left; Fig. 11, fourthcolumn). Although the tropical and extratropical Atlantic SST anomalies have opposing impacts on North
Pacific sea level pressure, the CAM2 response to the1438

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FIG. 11. October–March geopotential height anomalies at (left) 1000, (middle) 500, and (right) 200hPa (contours at intervals of 5, 10, and 10 m, respectively) simulated in the CCSM2 closed Bering
Strait experiment and the CAM2 experiments forced with CCSM2’s SST/sea ice anomalies over the
Atlantic and tropical Pacific, Atlantic, tropical Atlantic, and extratropical North Atlantic. Orange
stipples indicate the region of significant anomalies at the 98% level.

VOLUME 22

15 MARCH 2009

OKUMURA ET AL.

entire Atlantic (258S–758N) SST anomaly pattern isdominated by the tropical Atlantic forcing (Fig. 10,
upper right; Fig. 11, third column).
The tropical Atlantic SST anomalies explain about
two-thirds of the October–March Aleutian low responseto the full Atlantic plus tropical Pacific forcing (cf. thetwo left panels in Fig. 10). The remaining one-third is
attributable to the tropical Pacific SST anomalies andthe associated changes in atmospheric convection.
These tropical Pacific SST anomalies are, in turn, forced
by atmospheric teleconnection from the tropical Atlantic, implying a pivotal role of the tropical Atlantic in
linking the AMOC and North Pacific climate. Previous
studies of interannual and interdecadal Pacific variability show that a southward shift of the Pacific ITCZ andweakening of the Walker circulation act to strengthen
the Aleutian low by forcing barotropic Rossby waves
and triggering the Pacific–North American pattern(Trenberth et al. 1998; Alexander et al. 2002; Deser andPhillips 2006), consistent with our results.
The relative importance of the tropical Atlantic andPacific SST anomalies to the Aleutian low response
undergoes substantial changes from late fall throughearly spring (Fig. 12). Forced by the Atlantic and tropical Pacific SST anomalies, CAM2 simulates the monthly
cycle of the Aleutian low response in the closed Bering
Strait experiment fairly well, confirming that these SST
anomalies are responsible for the Aleutian low response(cf. gray and black lines in Fig. 12). The tropical Atlantic
SSTs play a dominant role in late fall (October–
November) and early spring (February–March, red line)
while the tropical Pacific SSTs become more importantin midwinter (December–January, green line). This seasonal sensitivity of the Aleutian low to tropical Pacific SSTs
is in agreement with studies of ENSO teleconnections(Alexander et al. 2002). The extratropical North Atlantic
cooling, on the other hand, damps the Aleutian low response during winter when the annular mode variability
becomes most vigorous (Thompson and Wallace 2000).
The consistency of the Aleutian low response among
the four coupled GCMs suggests that similar teleconnection mechanisms are at work in the other models.
The relative importance of each remote forcing mechanism, however, may be model-dependent because ofdifferences in the models’ mean climate. For example, inthe tropical Atlantic, most coupled GCMs fail to simulate the eastern equatorial cold tongue, resulting in an
ITCZ that is displaced too far south compared to observations (Davey et al. 2002; Richter and Xie 2008).
This tropical Atlantic bias is prominent in CCSM2 (Fig. 13).
The Caribbean warm pool is also absent and the mean
precipitation is less than 2 mm day21 in the western
tropical North Atlantic, where a strong cooling occurs in

1439

FIG. 12. Sea level pressure anomalies (hPa) averaged over the
Aleutian low region (308–708N, 1608E–1408W) as a function of
calendar month: CCSM2 closed Bering Strait experiment (thickgray) and CAM2 experiments forced with the CCSM2 SST/sea ice
anomalies over the Atlantic and tropical Pacific (black), tropicalAtlantic (red), extratropical North Atlantic (blue), and tropical
Pacific (green, estimated by taking the difference between the
Atlantic 1 tropical Pacific and Atlantic runs).

response to an AMOC weakening. With the mean SST
too low to sustain deep atmospheric convection, further
SST cooling does not significantly impact the atmospheric convection and hence the resultant forcing ofbarotropic Rossby waves. To test this hypothesis, werepeated the CAM2 tropical Atlantic experiment byreplacing the mean SST and sea ice fields with observed
climatologies from the Hadley Centre Sea Ice and Sea
Surface Temperature (HadISST) dataset for the period
1949–2001 (Rayner et al. 2003). In response to the same
tropical Atlantic SST forcing, replacing the climatology
with observations results in a doubling of the precipitation anomalies in the tropical North Atlantic andtripling of the Aleutian low response (cf. Fig. 14 to
the lower left panel of Fig. 10). The precipitation
anomalies in the tropical South Atlantic, on the other
hand, weaken significantly and the southward shift of theAtlantic ITCZ becomes inconspicuous. CCSM2 is oneof the models with the most severe tropical Atlantic bias
(see Fig. 1 in Timmermann et al. 2007) and its tropicalNorth Atlantic precipitation anomalies in the waterhosing experiment are smallest among the four coupled
GCMs (Fig. 1). Therefore, it is plausible that the atmospheric bridge from the tropical North Atlantic to
the North Pacific is stronger in the other three models

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JOURNAL OF CLIMATEVOLUME 22

FIG. 14. As in bottom left of Fig. 10 but the climatological SST fieldhas been replaced with observational data (HadISST).

importance of tropical and extratropical Atlantic SST
anomalies.

5. Summary and discussion

FIG. 13. (top) SST climatology simulated in the control run of
CCSM2 closed Bering Strait experiment (shading, 8C) and mean
precipitation simulated in CAM2 forced with this SST climatology(contours at intervals of 2 mm day21; stippled . 2 mm day21).
(bottom) As in the top panel but for observed SST climatology(HadISST, 1942–2001) and mean precipitation simulated in CAM2
forced with this SST. All fields are shown for October–March.

and also in the real world during the last glacial period
compared to CCSM2.
Using the atmospheric GCM component of HadCM3,
Sutton and Hodson (2007) examined the relative importance of tropical and extratropical Atlantic SST
anomalies for the atmospheric circulation changes associated with the Atlantic multidecadal oscillation,
which exhibits anomaly patterns similar to those in thewater-hosing experiment. The tropical Atlantic SST
anomalies explain a large part of the Aleutian low
change in this model (cf. their Figs. 3 and 4), consistent
with our result. On the other hand, Zhang et al. (2007)
find that extratropical North Atlantic SST anomalies
are more important than tropical Atlantic ones for the
North Pacific climate response associated with the Atlantic multidecadal oscillation using the atmospheric
GCM of GFDL CM2.1. In their experiment, however,
the atmospheric GCM is coupled to a slab ocean model
globally and the heating anomaly prescribed in the extratropical North Atlantic induces significant SSTchanges in the tropical North Atlantic via atmosphericteleconnections, making it difficult to isolate the relative

The analysis of water-hosing experiments with four
coupled GCMs revealed a strong linkage between the
AMOC and North Pacific climate, consistent withpaleoclimate reconstructions. When the AMOC is forced
to slow down in the models, strong cooling occurs along
the North Pacific oceanic frontal zone and the wintertime Aleutian low intensifies. This remote response
establishes quickly within a few decades. Through additional experiments with CCSM2 and CAM2, mechanisms for both oceanic and atmospheric teleconnections
are proposed. For the oceanic teleconnection, part of
the freshwater applied to the North Atlantic flows
through the Bering Strait into the North Pacific where
the anomalously cold and freshwater is advected bymean ocean currents. The resulting extratropical oceanic cooling strengthens the northeasterly trade winds,
reducing the tropical North Pacific SSTs and shifting the
Pacific ITCZ southward through the wind–evaporation–
SST feedback. For the atmospheric teleconnection,
precipitation changes associated with the tropical Atlantic and Pacific SST anomalies force barotropic
Rossby waves, inducing a cyclonic circulation anomaly
over the North Pacific from late fall through early
spring. The enhanced Aleutian low, in turn, cools theoceanic frontal zone by increasing surface heat losses
and southward Ekman transport, although this coolingis considerably weaker than that due to the Bering Strait
throughflow change. The tropical North Atlantic is aparticularly important region for this atmosphericbridge since it influences the atmospheric circulation in
the North Pacific not only directly but also indirectly bytriggering SST and precipitation changes in the eastern15 MARCH 2009

OKUMURA ET AL.

tropical Pacific. In the tropical North Atlantic, the excitation of atmospheric Rossby waves is quite sensitive
to the background SST and precipitation distributions,which are underestimated in most coupled GCMs.
The response of the Aleutian low to tropical Atlantic
SSTs cannot be readily explained and will require further investigation. A preliminary analysis, however,suggests that the upper-tropospheric convergence associated with the suppressed precipitation in the tropical North Atlantic forces barotropic Rossby waves that
propagate along the South Asian westerly jet and theninto the North Pacific (not shown). Near the entrance of
the South Asian waveguide, where the energy dispersion is small and the group velocity is exceptionallylarge (Hoskins and Ambrizzi 1993; Branstator 2002),
atmospheric convection changes in the tropical North
Atlantic can exert a significant impact on the North
Pacific atmospheric circulation. Haarsma and Hazeleger(2007) show that atmospheric Rossby waves forced byequatorial Atlantic SST anomalies propagate circumglobally through the South Asian waveguide in a matterof two weeks in their model. A Japanese community
atmospheric GCM forced with an idealized tropicalAtlantic SST dipole (Okumura et al. 2001) also obtains
a deepening of the wintertime Aleutian low when atmospheric convection is suppressed in the tropicalNorth Atlantic. In addition to Rossby wave propagation, the structure of intrinsic circulation variability may
also play an important role in shaping the anomaly
pattern over the North Pacific. The Aleutian low response in both open and closed Bering Strait experiments projects onto the second leading mode of North
Pacific variability in the control simulations, which resembles the Pacific–North American pattern (Wallace
and Gutzler 1981; not shown).
During the last glacial period when Bering Strait wasclosed, the atmospheric teleconnection from the tropicalAtlantic and Pacific is likely to have played a fundamental role in the AMOC–North Pacific teleconnection.
In the closed Bering Strait experiment, however, theNorth Pacific cools only by 18C, much smaller than themagnitude of 38–68C inferred from paleoclimate records(Thunell and Mortyn 1995; Hendy and Kennett 1999;
Kienast and McKay 2001). The weaker atmosphericbridge from the tropical North Atlantic due to model
errors may partly contribute to this gap. With the mean
Atlantic ITCZ displaced north of the equator, the tropical North Atlantic cooling induces a stronger Aleutian
low response that enhances the SST cooling by increasing surface heat fluxes and southward Ekman transport,
and shifting the subarctic gyre to the south. Despite theabsence of freshwater exchange with the Arctic Ocean
through Bering Strait, Kotilainen and Shackleton (1995)

1441

report rapid variability in ice-rafted detritus in the subarctic North Pacific Ocean sediment cores associatedwith the Greenland ice core records. During cold phases
in Greenland, increased discharge of icebergs might havefreshened the upper ocean and contributed to the surface cooling by inhibiting oceanic convection.
Another important issue not explored in this study is
the role of sea ice and its interaction with the ocean
and atmosphere. Cheng et al. (2007) compare two setsof water-hosing experiments under different backgroundstates for the present day and Last Glacial Maximum—thetime of maximum extent of the ice sheets during the lastglacial (;20 kyr ago). They find that the fundamental
mechanisms for global adjustment to an AMOC weakening are similar between the two experiments. However, under the Last Glacial Maximum condition, theclimatic impact is much more severe than that under thepresent-day condition primarily because of the more
extensive distribution of sea ice and stronger ice–albedo
feedback. In our CCSM2 experiments, the closure ofBering Strait alone only modestly enhances the meansea ice distribution in the Sea of Okhotsk, and sea ice
concentration anomalies in the water-hosing experiment are less than 10% over the subarctic North Pacific(not shown). The presence of continental ice sheets andtheir interaction with the overlying atmosphere may
have had a more profound impact on the North Pacificsea ice distribution during the Last Glacial Maximum
than the closure of Bering Strait. Sea ice, if present, acts
as an insulator between the ocean and atmosphere, reducing the thermal damping of anomalously cold outflow
from the Eurasian continent and thus amplifying theNorth Pacific cooling. The results of Cheng et al. (2007)
and our analysis of the oceanic teleconnection via Bering
Strait point to the importance of background states indetermining the climate response to freshwater perturbations in the North Atlantic.
Our analysis of coupled and atmospheric GCM experiments offers a framework for interpreting abrupt
climate shifts in the North Pacific and their relation tochanges in the AMOC in paleoclimate records beyond
the thermal advection mechanism. Furthermore, it has
significant implications for present-day and future climate. Concern about future climate change induced
by increasing greenhouse gases partly stems from thepossibility that Arctic warming and rapid melting ofcontinental/sea ice may halt the AMOC, leading to a
drastic transition of global climate (Manabe andStouffer 1993; Stocker and Schmittner 1997). Although
the current generation of coupled GCMs suggests thatsuch an extreme scenario is unlikely, most models do
predict a gradual weakening of the AMOC in responseto an increase in greenhouse gases (Gregory et al. 2005;

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Meehl et al. 2007). Superimposed on an overall greenhouse warming, the weakening AMOC may influence
climate over the North Pacific and other distant regions
according to the pathways described in this study. Indeed, the North Pacific displays a cooling trend over the
recent 50 yr associated with a cooling (warming) trend
in the North (South) Atlantic that is possibly caused bya weakening of the AMOC (Trenberth et al. 2006). This
SST anomaly pattern may also reflect Atlantic multidecadal oscillation, which involves modulation of the
AMOC intensity (Timmermann et al. 1998; Enfield andMestas-Nunez 1999; Delworth and Mann 2000; Knight
et al. 2005; Sutton and Hodson 2005). Based on analysesof reconstructed surface temperature data and coupled
GCM simulations, these studies show that the Atlantic
multidecadal oscillation yields a hemispheric anomaly
pattern with a significant SST change in the extratropical North Pacific that is nearly in phase with the
SST change in the North Atlantic. Atmospheric teleconnections from the Atlantic can be responsible for the
Pacific portion of this interbasin multidecadal variability (Timmermann et al. 1998; Enfield and Mestas-Nunez
1999; Sutton and Hodson 2007; Zhang and Delworth2007). Zhang and Delworth (2007) show that the
Aleutian low change forced from the Atlantic induces
SST anomalies in the oceanic frontal region with a lag ofseveral years in GFDL CM2.1, suggesting the transientadjustment of the gyre-scale circulation to a change inwind stress curl via long baroclinic Rossby waves (e.g.,
Miller et al. 1998; Deser et al. 1999). Under present-day
boundary conditions, the oceanic teleconnection via
Bering Strait may also come into play. This interbasin
oceanic adjustment is suggested to stabilize the AMOC,
hence present-day, interglacial climate (De Boer andNof 2004; Hu et al. 2007). While it remains to be seen
whether the AMOC–North Pacific teleconnection
mechanisms described in this study operate in association with the anthropogenic change and multidecadal
variability of the AMOC, they potentially offer wideranging applications to Atlantic and Pacific climate
variability on long time scales.
Acknowledgments. The authors are grateful to J.
Jungclaus, J. Yin, and B. Dong for providing the CMIP/
PMIP water-hosing experiment data from ECHAM5OM1, GFDL CM2.1, and HadCM3. Discussions with
Y.-O. Kwon and J. M. Wallace helped refine the analysis and its physical interpretation. D. S. Kaufman
shared his unpublished analysis of lake sediments from
Alaska, which motivated us to study the mechanismsof the Aleutian low response in the models. We thankthe anonymous reviewers for their useful suggestions.
Y. Okumura is supported by the NOAA Climate and

VOLUME 22

Global Change Postdoctoral Fellowship. A. Timmermann
and S.-P. Xie acknowledge support from the Japan
Agency for Marine-Earth Science and technology(JAMSTEC), NASA through Grant NNX07AG53G,
and NOAA through Grant NA17RJ1230. A. Timmermann
is also supported by NSF Grant ATM06-28393. This
study was in part supported by the Office of Science(BER), U.S. Department of Energy, Cooperative
Agreement DE-FC02-97ER62402. The CAM2 simulations were conducted using computational resourcesfrom the NCAR Director’s Reserve.

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