Clim Dyn
DOI 10.1007/s00382-010-0764-8

Coupled atmosphere–mixed layer ocean response to ocean heat
flux convergence along the Kuroshio Current Extension
Young-Oh Kwon • Clara Deser • Christophe Cassou

Received: 30 October 2009 / Accepted: 8 February 2010
Ó Springer-Verlag 2010

Abstract The winter response of the coupled atmosphere–ocean mixed layer system to anomalous geostrophic ocean heat flux convergence in the Kuroshio
Extension is investigated by means of experiments with anatmospheric general circulation model coupled to an
entraining ocean mixed layer model in the extra-tropics.
The direct response consists of positive SST anomalies
along the Kuroshio Extension and a baroclinic (low-level
trough and upper-level ridge) circulation anomaly over the
North Pacific. The low-level component of this atmospheric circulation response is weaker in the case without
coupling to an extratropical ocean mixed layer, especiallyin late winter. The inclusion of an interactive mixed layer
in the tropics modifies the direct coupled atmospheric
response due to a northward displacement of the PacificInter-Tropical Convergence Zone which drives an equivalent barotropic anomalous ridge over the North Pacific.
Although the tropically driven component of the North
Pacific atmospheric circulation response is comparable tothe direct response in terms of sea level pressure amplitude,

Electronic supplementary material The online version of thisarticle (doi:10.1007/s00382-010-0764-8) contains supplementarymaterial, which is available to authorized users.
Y.-O. Kwon (&)
Physical Oceanography Department,
Woods Hole Oceanographic Institution,
Clark 320A, MS#21, Woods Hole, MA 02543, USAe-mail: yokwon@whoi.edu
C. DeserNational Center for Atmospheric Research,
Boulder, CO, USA
C. Cassou
CNRS-CERFACS, Toulouse, France

it is less important in terms of wind stress curl amplitude
due to the mitigating effect of the relatively broad spatial
scale of the tropically forced atmospheric teleconnection.

1 Introduction
Various observational studies indicate that decadal-scalesea surface temperature (SST) variations along the Kuroshio Current Extension result from geostrophic ocean circulation changes in response to basin-scale wind stress curl
fluctuations 3–5 years earlier (Miller et al. 1998; Deseret al. 1999; Xie et al. 2000; Qiu 2000; Seager et al. 2001;
Schneider and Miller 2001; Vivier et al. 2002; Qiu 2003;
Kelly 2004). The circulation-induced SST anomalies inturn alter the local air–sea energy exchange, in particular
the surface turbulent energy fluxes in winter (Seager et al.
2001; Tanimoto et al. 2003; Kelly 2004). The nature of theatmospheric response to the surface energy flux anomalies
remains difficult to discern from observations alone, however, due to the high levels of internal and tropically forced
atmospheric variability (Qiu et al. 2007; Frankignoul andSennéchael 2007).
Some coupled general circulation (Pierce et al. 2001;
Kwon and Deser 2007; Zhong et al. 2008) and observationally based analytical (Qiu et al. 2007) modellingstudies suggest that ocean–atmosphere interaction within
the North Pacific can sustain decadal-scale oscillations. In
these studies, the response of the large-scale atmosphericcirculation to SST anomalies induced by ocean heattransport variations along the western boundary currents isa key component. In particular, the strength and pattern of
the atmospheric circulation response to the extra-tropical
SST anomalies determines whether such decadal-scalevariability is a result of mutual coupling between the ocean

123

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

and atmosphere (Latif and Barnett 1996; Neelin and Weng
1999; Pierce et al. 2001; Wu and Liu 2005; Kwon andDeser 2007; Zhong and Liu 2009) or an internal mode of
the ocean (Frankignoul et al. 1997; Saravanan and
McWilliams 1998; Cessi and Primeau 2001; Qiu 2003;
Berloff et al. 2007).
The tropics may also play a role in initiating decadalscale variability over the North Pacific (Newman et al.
2003; Vimont 2005; Schneider and Cornuelle 2005;
Alexander 2010). In particular, the El Nino/Southern
Oscillation (ENSO) phenomenon has been suggested tocontribute to the Pacific Decadal Oscillation (PDO),
defined as the leading mode of SST variability in the North
Pacific (Mantua et al. 1997; Zhang et al. 1997), via
atmospheric teleconnections and the thermal response of
the deep winter ocean mixed layer in the North Pacific(Newman et al. 2003; Vimont 2005). These ENSO-based
reconstructions of the PDO, however, do not fully accountfor the decadal-scale SST variations along the Kuroshio
Extension. Schneider and Cornuelle (2005) showed that
additional oceanic processes are needed to explain the
low-frequency component of SST variability along theKuroshio Extension. In particular, westward propagating
baroclinic Rossby waves driven by ENSO-induced windstress curl fluctuations in the central and eastern North
Pacific contribute to SST anomalies along the Kuroshio
Extension via changes in vertical mixing and heat transport. Although the tropics play an important role in generating decadal-scale SST variability over the North Pacific(i.e., the PDO), feedbacks from the ocean to the atmosphere within the North Pacific may also contribute.
It is clear from the issues outlined above that a better
understanding of the atmospheric response to ocean circulation-induced SST anomalies along the Kuroshio
Extension would advance understanding of decadal climate
variability over the North Pacific. A common approach tothis problem has been to use an atmospheric general circulation model (AGCM) forced with prescribed SST
anomalies; however, the results diverge greatly among
different studies (see the review by Kushnir et al. 2002).
Progress made in the past decade suggests that the diverse
equilibrium atmospheric responses among different SSTspecified AGCM studies can be attributed in part to differing sensitivity of transient eddy feedbacks (Peng et al.
1997; Peng and Whitaker 1999) and structure of themodels’ internal variability (Peng and Robinson 2001;
Deser et al. 2004). Linear model studies predict downstream low pressure anomalies at sea-level with a baroclinic vertical structure in response to shallow positive
diabatic heating anomalies (Hoskins and Karoly 1981;
Hendon and Hartmann 1982). This so-called ‘‘warm-low’’equilibrium response is associated with an anomalous
northerly low-level wind balancing the imposed anomalous

123

diabatic heating. However, nonlinear transient eddy feedbacks may substantially modify the linear response andproduce an equivalent barotropic downstream highresponse (e.g. Peng and Whitaker 1999; Deser et al. 2007).
Several recent studies have noted the limitation of
AGCM experiments with specified SST anomalies (e.g.
Kumar and Hoeling 1998; Bretherton and Battisti 2000;
Yulaeva et al. 2001; Sutton and Mathieu 2002). The
experiments in which SSTs are not allowed to respond tochanges in atmospheric circulation can produce misleading
results, since the SST anomalies should result from theimbalance between the net heat flux across the air–sea
interface and the lateral convergence of heat into the ocean
mixed layer.
One way to provide a more realistic assessment of theocean’s role in climate variability is to use a model configuration in which the atmosphere is coupled to an ocean
mixed layer with prescribed ocean heat transport variations
due to dynamical processes not included in the ocean
mixed layer model. In this configuration, SST and surface
heat flux anomalies reflect to a greater degree the true twoway coupling between the atmosphere and the ocean mixed
layer. Two pioneering studies, one for the North Pacific(Yulaeva et al. 2001) and the other for the North Atlantic(Sutton and Mathieu 2002), have applied this approach.
Both studies used idealized forms of ocean heat flux convergence (OHFC) anomalies, and both employed AGCMs
coupled to a non-entraining slab ocean mixed layer modelfor their 10-year long integrations. The winter atmosphericcirculation response in the study of Yulaeva et al. (2001)
consisted of an equivalent barotropic low downstream ofanomalously warm SST induced by lateral OHFC in the
Kuroshio Extension. Sutton and Mathieu (2002) also
reported a downstream equivalent barotropic low pressure
response to anomalous OHFC in the Gulf Stream region.
Liu and Wu (2004) compared four approaches forinvestigating the atmospheric response to SST anomalies in
the Kuroshio Extension region: (1) specified SST anomalies in an AGCM, (2) specified surface air-sea heat flux
anomalies in an AGCM, (3) specified ocean temperature
anomalies in the upper 200 m as initial conditions to a fully
coupled ocean–atmosphere GCM, and (4) as in (3) but
allowing only thermal coupling. The specified SST
anomaly experiment produced a downstream equivalentbarotropic high-pressure response to a positive SST
anomaly, while the air–sea heat flux anomaly experimentproduced a downstream equivalent barotropic low
response. The fully coupled experiment produced anequivalent barotropic high pressure response which they
concluded was a combination of the responses in cases (1)and (2) above. The thermally coupled experiment produced
a similar response to that of the fully coupled experiment
but with about half the magnitude. Like the specified SST

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

experiment, the specified surface air–sea heat flux experiment is limited in the sense that the forcing does not result
from an imbalance between the heat flux convergences in
the lower atmosphere and upper ocean. Thus, the surface
air-sea heat flux specified experiments are not necessarilyequivalent to the simulations in Sutton and Mathieu (2002)
and Yulaeva et al. (2001) who specified the lateral OHFC
anomalies and allowed both SSTs and the surface air-sea
heat fluxes to be determined prognostically. On the other
hand, the thermally coupled experiment with initial upper
ocean temperature anomalies is more in line with theexperimental design of Sutton and Mathieu (2002) and
Yulaeva et al. (2001).
A subsequent study by Liu et al. (2007) explored further
the fully coupled model approach, specifying an initial
distribution of upper ocean temperature anomalies in the
Kuroshio Extension and examining the subsequent transient coupled response over a 4-year period. They foundthat the atmospheric circulation response for early and late
winter differ considerably, as in Peng et al. (1997). In
November–January, the warm SST anomalies are accompanied by an equivalent barotropic ridge over the North
Pacific and downstream wave train resembling the PacificNorth American (PNA) pattern (Wallace and Grutzler
1981). On the other hand, the atmospheric response in
February–April is significantly weaker than that inNovember–January, with a baroclinic warm–low structure.
They briefly speculated that the distinctive responses inearly and late winter may be due to the different background atmospheric states as suggested by Peng et al.
(1997).
Kwon and Deser (2007) analyzed decadal variability in
the North Pacific from a 650-year control integration ofCommunity Climate System Model version 2 (CCSM2), a
coupled general circulation model. They found that thewind stress curl response to ocean circulation-induced SST
variations in the Kuroshio Extension played a key role insustaining quasi-decadal (15–20 year) coupled atmosphere–ocean oscillations. Specifically, positive wind stresscurl anomalies in the latitude band 30°–50°N occur in
response to positive SST anomalies along the Kuroshio
Extension. This response in turn implies subsequentdevelopment of negative SST anomalies in the Kuroshio
Extension due to oceanic Rossby wave adjustments, i.e. a
negative feedback. The wind stress curl response diagnosedin Kwon and Deser (2007) was determined empiricallyusing lag regression analysis, supplemented with preliminary results from additional atmospheric modelingexperiments. A summary of the various studies cited above
is given Table 1.
The purpose of this study is to provide a more comprehensive assessment of the atmospheric circulation
response to OHFC anomalies in the Kuroshio Extension

using a broader suite of modeling experiments than inKwon and Deser (2007). In particular, we systematicallycompare the responses using the standard ‘‘specified SST’’approach with the alternative ‘‘specified OHFC’’ approach.
In order to understand the atmospheric response in the fully
coupled system of CCSM2, we retain the component
models of CCSM2 as much as possible to ensure consistency between the CCSM2 and the simplified experiments.
Specifically, we employ the atmospheric model componentof CCSM2 for the ‘‘specified SST’’ experiments, and asimplified coupled model consisting of the atmosphericmodel component of CCSM2 coupled to an entraining
ocean mixed layer model for the ‘‘specified OHFC’’ integrations. The OHFC anomalies along the Kuroshio
Extension diagnosed from the CCSM2 are prescribed in thesimplified coupled model experiments. The resulting SST
anomalies are then specified in the atmospheric GCM
experiments. Interaction of the atmospheric response over
the North Pacific with the tropics is also examined in an
idealized setting by comparing the results from the simplified coupled model experiments with and without aninteractive tropical ocean mixed layer. Section 2 describes
the model and experiment design. Results are presented in
Sect. 3 and discussed in Sect. 4. Conclusions are presented
in Sect. 5.

2 Model experiments
2.1 Coupled model: CMLM
The coupled model used in this study consists of an
atmospheric general circulation model, a mixed layer
ocean model, a thermodynamic sea-ice model and a landmodel (Cassou et al. 2007). The coupled model has almost
identical atmospheric and land components as the CCSM2(Kiehl and Gent 2004), while the ocean component isgreatly simplified, and the sea-ice component retains only
the thermodynamic portion of that in CCSM2. The atmospheric component model is the Community Atmospheric
Model version 2 (CAM2) (Kiehl and Gent 2004), which
has T42 horizontal resolution (=about 2.8° latitude by 2.8°
longitude) and 26 vertical levels using a hybrid coordinate.
The land surface and the sea-ice components are theCommunity Land Model (CLM2) (Bonan et al. 2002) andthe thermodynamic component of the Community Sea-Ice
Model (CSIM4) (Briegleb et al. 2004), respectively. Ocean,sea-ice, and land models are all aligned with the atmospheric model grid.
The ocean component is an entraining mixed layer
model (MLM), which is an array of independent single
column multi-layer models with explicit mixed layer
physics. The mixed layer depth (MLD) is calculated

123

123
SST
HFLX
0–200 m T
0–200 m T

AGCM–OMLM

AGCM

AGCM

CGCM–TCPL

CGCM

Yulaeva et al. (2001)

Liu and Wu (2004)

JFM

JFM

JFM
0.8

4.0

4.0

4.0

1.2

FMA
JFM

1.2

1.6

1.6

0.4

2.0

0.75

2.5
2.5

Maximum SST
anomalies (°C)

NDJ

Dec

Dec

Dec

Dec

DJF

Perpetual Jan
Perpetual Feb

Month(s)

-5

-2

-1

-2

-0.5

?1.5

?5

?2.5

–

?2

-2

-2.5
(-0.6)SLP (hPa)

-35

(?0)

?10

-10

-3

?15

?40

?20

-20

?20

[-10]

-12
(?6)

Z850 (m)

-60

(?10)

?20

(?10)

(0)

?15

?65

?35

–

?35

-20

-5
?20

Z500 (m)

-60

?30

?40

?20

(?1)

?30

?80

?40

-40

?40

[-20]

(0)
?30

Z250 (m)Maximum responses in the North Pacific

BT Trough

BC

BC*

BC

BC

BT Ridge

BT Ridge

BT Ridge

BT TroughBT Ridge

BT TroughBC
BT Ridge

Overall

Yes

No

Yes

No

Yes

Yes

Yes

Yes

No

No

Yes

No
No

Activetropics

100

100

100

100

160

160

40

60

40

40

10

384
384

Independent
sample size

AGCM Atmospheric General Circulation Model, OMLM Ocean Mixed Layer Model, CGCM Coupled Atmophere-Ocean General Circulation Model, TCPL Thermodynamic Coupling Only, ML
Mixed Layer, OHFC Ocean Heat Flux Convergence, HFLX Surface Heat Flux, T Ocean Temperature, BC Baroclinic, BT Equivalent Barotropic. BC* the baroclinic total response is the sum of
a local baroclinic response and a remote barotropic response

Values in (not in) parentheses are not (are) statistically significant at the 95% level; values within brackets lack estimates of statistical significance

ML OHFC

SST

AGCM

AGCM–OMLM

0–1000 m OHFC

AGCM–OMLM

Sutton and Mathieu (2002)

0–1000 m OHFC

AGCM–OMLM

This Study

0–560 m T

CGCM

Liu et al. (2007)

ML OHFC

SST

AGCM

Peng et al. (1997)

Specified anomalies

Experiment type

Reference

Table 1 Characteristics of the winter atmospheric circulation response to upper ocean thermal anomalies in the Kuroshio Extension region in various studies, except Sutton and Mathieu (2002)
which is for the North Atlantic Gulf Stream region

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

prognostically for each time step based on Gaspar (1988)’s
formulation as implemented by Alexander and Deser(1995). In the present study we use a version of the MLM
modified by Cassou et al. (2007) which has 36 verticallayers in the upper 1500 m with 15 layers in the upper
100 m. The MLM allows us to specify the OHFC anomalies into multiple vertical layers of the upper ocean andalso to examine the role of entrainment.
The climatological mean surface heat flux correction
and salt flux correction are applied to account for missing
physics in the MLM such as ocean heat transport by themean currents and diffusion, as well as the model biases inthe surface fluxes at the atmosphere–ocean boundary to alesser extent. Note that Ekman transport is not included in
the version of MLM used in this study. Hereafter, thecoupled model is referred as CMLM. The reader is referred
to Cassou et al. (2007) for further description of the
CMLM.
2.2 CMLM control integration
The 150-year control integration of CMLM (hereafter
CMLM-C) from Cassou et al. (2007) is used as a baseline
experiment for comparison with the perturbation integrations. The control integration, with the inclusion of flux
corrections, produces a reasonable climatology as assessedin detail by Cassou et al. (2007). The main deficiency of the
CMLM is the weaker year-to-year variability of SST over
the North Pacific compared to that in CCSM2 and observations (maximum rms values *0.5°C in CMLM compared to 0.8°C in CCSM and 0.9°C in observations,excluding the region directly east of Japan). This may be
due in part to the absence of lateral ocean dynamics in the
MLM. Atmospheric variability is well simulated in CMLM
compared to observations and CCSM2, including theintensity and location of the storm track over the North
Pacific. The reader is referred to the web supplement(http://www.cgd.ucar.edu/cas/cdeser/CMLM_supfig.html)
for comparisons among observations and the CMLM andCCSM2 control integrations.
2.3 CMLM perturbation integrations
To investigate the coupled atmosphere–ocean mixed layer
response to ocean heat transport variations, perturbation
experiments are designed with additional geostrophic
OHFC anomalies in the Kuroshio Extension (35°–45°N,
140°E–180°) specified in the upper 1000 m of the CMLM.
Two 100-year long perturbation integrations are performed
with identical design except for the treatment of thetropical oceans. The first perturbation integration, which is
our primary experiment, has the interactive MLM only inthe extra-tropics (north of 20°N and south of 20°S) to

isolate the direct response to the local anomalous OHFC
forcing in the North Pacific. The climatological mean SST
from the CMLM-C is specified in the tropics. The additional perturbation integration has the interactive MLM
over the entire global ocean including the tropics. Hereafter, the first and second perturbation integrations arereferred as CMLM-E and CMLM-G, respectively, where E
stands for extra-tropics and G stands for global.
The geostrophic OHFC anomalies are diagnosed fromthe years 350–999 of the CCSM2 control integration(Kwon and Deser 2007). The anomalies for each month arecalculated as the root-mean-square of 10-year low-passfiltered time series of the corresponding month’s geostrophic OHFC. For example, January values for the 650year period are first 10-year low-pass filtered and then the
root-mean-square of the filtered time series is assigned as
the January anomaly. The 10-year low-pass filter is applied
to obtain typical amplitudes of the decadal variability. The
anomalies are calculated independently for each (longitude, latitude, depth) point. Then the repeating annual cycleof the anomalies are specified in CMLM. The Ekman
OHFC has been removed from the total OHFC to retain
only the geostrophic component, where we have assumed
that the total Ekman OHFC is distributed uniformly
throughout the mixed layer.
The distribution of specified geostrophic OHFC anomalies in the upper 10 m in March is shown in Fig. 1a.
Values are about 4.2 W m-2 on average in the Kuroshio
Extension, with maximum amplitudes up to 15 W m-2
along 40°N near the coast of Japan. Note that we havespecified only geostrophic OHFC anomalies in the domain
confined to the Kuroshio Extension (35°–45°N, 140°E–180°) to ensure an appropriate oceanic forcing based on theanalysis of CCSM2 by Kwon and Deser (2007). Spatial
distributions in March are representative of thosethroughout the year with modulation of amplitude asshown in Fig. 1b. The average anomalies in the upper 10 m
are maximum in February (=4.6 W m-2) and minimum in
September (=2.2 W m-2). The geostrophic OHFC
anomalies have significant magnitude beneath the climatological mean MLD (shown as the dashed lines inFig. 1b–d). The vertical sections along 40°N and 150°E
show maximum amplitudes at the surface along 40°N near
the coast of Japan and a gradual decrease with depth anddistance from the maximum amplitude region (Fig. 1c, d).
A few observational estimates of geostrophic OHFC
anomalies have been reported against which our estimatesfrom CCSM2 may be compared. Qiu (2000) estimatedJanuary–March average non-seasonal geostrophic OHFC
anomalies in the Kuroshio Extension to be about 5 9
10-8°C s-1. Qiu (2000) estimate is based on the surface
geostrophic velocity from sea surface height measurementsof satellite altimetry during 1993–1999. Vivier et al. (2002)

123

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

(b)

(a)
o

0

6

6
100 6

40oN

Depth (m)

60 N

5

6

4

200
2
300
400

o

20 N o
120 E

150oE

180oW

150oW

120oW
500

J

F

M

A

M

J

J

A

S

O

N

D

Month

(c)

(d)

0

0

10

10
5

200

400

3

600

2

Depth (m)Depth (m)

200

5
400
3
600
2

800
1000
140

800

1

160

180

1
200

220

240

o

Longitude ( E)

1000
20

30

40

50

60

o

Latitude ( N)

Fig. 1 Geostrophic ocean heat flux convergence anomalies for the
CMLM perturbation integrations. a Map of March values integrated
over upper 10 m. Contour interval is 2.5 W m-2. b Values averaged
over the Kuroshio Extension region (35°–45°N, 140°E–180°) as afunction of month and depth, expressed as a temperature tendency(°C s-1). Contour interval is 2 9 10-8°C s-1. Dashed line indicatesthe climatological monthly mean mixed layer depth (MLD) from the
CCSM2 control integration. c Vertical section of March values along
40°N. Contour interval is 5 9 10-8°C s-1 and 1 9 10-8°C s-1, forthe thick and thin solid contours, respectively. Dashed line indicatesthe March climatological mean MLD from the CCSM2 control
integration. d Same as c, but along 150°E

estimated non-seasonal lateral (geostrophic ? Ekman)
OHFC anomalies in the Kuroshio Extension to be about
2 9 10-8°C s-1. They used a three-dimensional advection–diffusion ocean model with specified geostrophic andEkman velocities estimated from satellite altimetry measurements and wind stress from the NCEP-NCAR Reanalysis (Kalnay et al. 1996). By assimilating observed SST
and subsurface temperature data into a simple upper-ocean
mixed layer model for 1970–2000, Kelly (2004) estimatednon-seasonal root-mean-square of lateral OHFC anomalies
in the Kuroshio Extension and its recirculation gyre to be
9 W m-2 (*0.8 9 10-8°C s-1) for the upper 300 m. Both
Vivier et al. (2002) and Kelly (2004) estimates are not justfor the winter months, but for the entire year. Thus, thegeostrophic OHFC anomalies from CCSM2 are consistent
with, albeit somewhat larger than, previous estimates basedon the available observations.
The equilibrium responses of CMLM-E and CMLM-G
are examined by computing the differences between themean of the 100-year perturbation integrations and themean of the 150-year control integration CMLM-C. The
first 10 years (5 years) of the CMLM-E and CMLM-G(CMLM-C) integrations are excluded due to spin-up
effects. The 95% statistical significance level of the
response is assessed using a Student’s t test. Each year is
considered as independent, so that the perturbation integrations and the control integration are considered as

ensembles of 100 and 150 members, respectively. The
robustness of the response is also evaluated by comparing
sub-samples of the model output (e.g., first vs. second half,
odd vs. even years). Overall, consistent results are found
for the sub-sampled years (not shown). Note that the perturbation integrations are 100 years in length in order toenhance the signal to noise ratios of the responses, not tosimulate the decadal-scale variability found in the CCSM2
coupled general circulation model.

123

2.4 Specified SST experiments
In addition to the coupled model experiments, we haveconducted specified SST experiments with CAM2 using
SSTs from the CMLM experiments. This SST-specified
configuration is the traditional method for examining theatmospheric response to SST anomalies as discussed
above. To ensure an appropriate comparison with the
response in CMLM-E to the anomalous OHFC, a 150-year
stand-alone CAM2 control integration (hereafter AGCMC) is first performed with climatological mean annual cycleof SST and sea-ice concentration from the CMLM-C
specified as the boundary condition. The annual cycle ofthe equilibrium SST responses of CMLM-E in the Kuroshio Extension (35°–45°N, 140°E–180°) is then added asanomalies to the SST boundary condition of AGCM-C
in another 100-year stand-alone CAM2 perturbation

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response
Table 2 Summary of the simulations used in this study
Experiment nameSpecified perturbation

Mixed layer ocean

Length of integration (years)

CMLM-C

None

Global150

CMLM-EOHFCA within KE from CCSM2

Poleward of 20°N,S

100

CMLM-GOHFCA within KE from CCSM2

Global100

AGCM-C

None

None

150

AGCM-ESSTA within KE from CMLM-E

None

100

KE and OHFCA refer to the Kuroshio Extension (35°–45°N, 140°E–180°) and the ocean heat flux convergence anomalies, respectively

We will begin with the results from the CMLM-E experiment to describe the responses most directly driven by
North Pacific OHFC anomalies. The annual cycles of the
responses within the forcing region (35°–45°N, 140°E–180°) are presented in Fig. 2. Ocean temperatures increase
throughout the water column as a direct response to thespecified anomalous OHFC. Maximum temperature
anomalies are found around 300 m, although the specified
anomalous OHFC is largest near the surface (Fig. 1b). The
maximum warming is located below the maximum winterMLD (*220 m) due to damping of the temperature
anomalies within the mixed layer by the surface heat flux.
Maximum SST anomalies are found in February and Marchin accordance with the maximum anomalous OHFC forcing as well as the deepest mixed layer which entrains
warmer anomalies from the thermocline. The mixed layer
is deeper in the perturbation experiment (thick solid line in
Fig. 2c) than the control integration (dashed line in thesame panel) especially in winter by up to 100 m. Deepening of the winter mixed layer is primarily due to a
weakening of the stratification in response to the specified
anomalous OHFC between 100 and 200 m where the top of
the thermocline is located in the control integration.
Net surface heat flux anomalies are from the ocean toatmosphere year-round, with maximum values in January–
March (Fig. 2b). The net surface heat flux feedback in
December–March ranges from 26 to 33 W m-2°C-1,
somewhat smaller than the corresponding values estimatedfrom the observations (40–50 W m-2°C-1) (Frankignoul
and Kestenare 2002; Park et al. 2005) and the CCSM2(39 W m-2°C-1 for the December–March average) (Kwon

Pressure (hPa)

3.1 CMLM-E experiment(a) Qdia
200
400
600

10

1000
JAN

2

2

10

800

4

4

FEB MAR APR MAY JUNJULAUG SEP OCT NOV DECJUL

AUG SEP OCT NOV DEC(b) Qnet
W m−2

3 Results

and Deser 2007). The annual cycle of the atmosphericdiabatic heating response closely follows that of the net
surface heat flux (Fig. 2a). The significant responses reachup to *400 hPa from September to March. Diabatic
heating responses near the surface below *900 hPa aredominated by the vertical diffusion of sensible surface heat
flux anomalies, while condensational heating anomalies areprimarily responsible for the responses above *900 hPa.
Hereafter, we focus on the January–March (winter)average response when the ocean-to-atmosphere heat flux
and SST anomalies are largest. The individual monthly
responses during winter are given at http://www.cgd.ucar.
edu/cas/cdeser/CMLM_supfig.html. A cross-section of thewinter response along 40°N exhibits maximum MLD,

100
80
60
40
20
0
JAN

FEB MAR APR MAY JUN(c) T & MLD
0

Depth (m)integration (hereafter AGCM-E). The equilibrium response
for the stand-alone CAM2 perturbation integration isassessed following the same method as the CMLM perturbation integrations. All five integrations used in thisstudy are summarized in Table 2.

100

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4

300
400
500
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4
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FEB MAR APR MAY JUNJULAUG SEP OCT NOV DEC

Fig. 2 Annual cycle of CMLM-E responses averaged over the
Kuroshio Extension (35°–45°N, 140°E–180°). a Diabatic heating.
Contour interval is 2 9 10-6 K s-1. Zero contour is suppressed.
Shading indicates significance at 95% level. b Net surface heat flux.
Positive is from ocean to atmosphere. Note that the values for all
12 months are significant at 95% level. c Ocean temperature. Contour
interval is 0.5°C. Shading covers the whole section because all values
are significant at 95% level. Dashed thick contour is the MLD fromthe control integration and the solid thick contour is the MLD fromthe perturbation experiment. MLD anomalies for all the months
except July are significant at 95% level

123

Pressure (hPa)

Y.-O. Kwon et al.: Coupled atmosphere: mixed layer ocean response(a) Qdia(a) SST

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Fig. 3 Same as Fig. 2 but for winter (JFM) CMLM-E responses
along a zonal section averaged over 35°–45°N. The contour interval is
5 9 10-6 K s-1 in (a) and 1°C starting from 0.5°C in (c). Dashedcontours in (a) indicate the negative anomalies. The net surface heat
flux responses in (b) are significant at 95% level except for thelongitude band 195°–205°E, and the MLD responses in (c) aresignificant at 95% level except east of 215°E

surface heat flux, and diabatic heating anomalies near thewestern end of the section (Fig. 3), collocated with the
maximum specified OHFC (Fig. 1). The cross-section
exhibits a marked contrast between the directly forced
region to the west of the dateline and the downstream
region to the east of the dateline where the surface heat flux
response is weak and negative, i.e. from atmosphere toocean. The anomalous heat is transported from the forcing
region by the mean westerlies and deposited into the ocean
via surface heat flux anomalies to cause a warming
downstream. In contrast to the subsurface temperature
anomaly maximum in the directly forced region west of thedateline, positive ocean temperature anomalies east of thedateline are concentrated within the mixed layer, althoughvertical diffusion extends them downward. Similarly, thediabatic heating response exhibits a maximum in the
middle troposphere near the forcing region, while tothe east of the dateline the negative anomalies are confinednear the surface and are due to vertical diffusion associatedwith anomalous sensible surface heat flux from atmosphereto ocean.
The spatial pattern of the winter SST response in theforcing region directly reflects that of the OHFC forcing(Fig. 4a). However, the amplitude of the SST response isabout four times larger than the root-mean-square of the
10-year low-pass filtered SST in the same region

123

120oE

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Fig. 4 Winter (JFM) CMLM-E responses. a SST. Contour interval is
1°C, with the 0.25°C and 0.5°C contours indicated by the dashedlines. Shading indicates significance at 95% level. b Net surface heat
flux. Positive values denote a flux from ocean to atmosphere. Contour
interval is 50 W m-2 for values greater than or equal to zero (solid
contours); dashed contours show the -8 W m-2 values. c 850 hPa
diabatic heating. Contour interval is 10 9 10-6 K s-1. Dashedcontours indicate -2 9 10-6 K s-1. d Surface wind stress (vectors)
superimposed upon zonal wind stress from the control integration(CMLM-C; contours). Contour interval is 0.05 N m-2. Positive andzero (negative) values are plotted with solid (dashed) contours(35°–45°N, 140°E–180°) from the CCSM2, i.e. 3.0°C for
the CMLM-E compared to 0.7°C for the CCSM2. Lack ofdynamical damping due to neglected lateral processes in
the MLM (i.e. horizontal advection and diffusion) appears
to be responsible for the discrepancy.
Weak (\1°C) positive SST anomalies (indicated by thedashed contours) outside of the forcing region are driven
by atmosphere-to-ocean (i.e. negative) surface heat flux
anomalies (Fig. 4a, b). The spatial pattern of the surface
heat flux and the 850 hPa diabatic heating anomalies
exhibit a clear distinction between the directly forced
region in the Kuroshio Extension and the opposite-signed
anomalies outside of this region. The spatial pattern ofpositive diabatic heating anomalies does not change muchup to 400 hPa except for a slight eastward displacementwith height (not shown). The surface wind stress response
exhibits an anomalous cyclonic circulation basin-wide inY.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

extent and centered over the maximum positive SST
anomalies along the Kuroshio Extension (Fig. 4d). The
comparison with the climatological mean zonal wind stress(contours in the Fig. 4d) suggests that the anomalies represent a southward shift of the surface westerly jet.
The CMLM-E sea-level pressure (SLP) and 250 hPa
geopotential height (Z250) responses for January–March in
the Northern Hemisphere north of 20°N are presented in
the first column of Fig. 5 along with the SST and net
surface heat flux responses. The basin-wide negative SLP
anomalies in the North Pacific correspond to the cyclonicwind stress anomalies in Fig. 4d. The maximum low
pressure anomaly is located slightly downstream of the
maximum SST anomaly with amplitude *2 hPa in
response to a *3°C SST warming averaged over theforcing region. A significant high pressure response isfound over the North Atlantic and downstream over western Siberia, with maximum amplitude *2 hPa. This
atmospheric circulation response over the North Atlantic
locally drives positive anomalies in the atmosphere-toocean heat flux which in turn leads to positive SST
anomalies of modest amplitude. The Z250 responses
exhibit generally positive anomalies over high latitudes,with a maximum over the North Atlantic (*40 m), andnegative anomalies to the south. The zonally asymmetricFig. 5 Winter (JFM) responses.
(Top row) 250 hPa geopotential
height, (second row) sea-levelpressure, (third row) net surface
heat flux, and (bottom row) SST
from (first column) CMLM-E,(second column) CMLM-G,(third column) CMLM-G minus
CMLM-E, and (fourth column)
AGCM-E. Positive (negative)
responses are plotted with solid(dashed) contours. Zero
contours are suppressed.
Contour intervals are 10 m,
1 hPa, 50 W m-2, and 1°C,
respectively. Note that theadditional thin solid and dashedcontours for the net surface heat
flux and SST indicate±10 W m-2 and ±0.2°C,
respectively. Shading indicatessignificance at 95% level. The
projection is for the northernhemisphere poleward of 20°N
with the latitude circles at 40, 60
and 80°N

CMLM−E

component of the Z250 response exhibits a circumglobal
zonal wavenumber five pattern along the midlatitude jet(Branstator 2002). The vertical structure of the circulation
response is baroclinic over the North Pacific and equivalentbarotropic over the North Atlantic and western Siberia. The
North Atlantic component of the response is robust andstable throughout the winter months of November–March,while the negative SLP anomaly response in the North
Pacific begins to emerge in December and exhibits full
amplitude from January through March, following thedevelopment of the SST anomalies (Fig. 6). The Z250
anomalies in the North Pacific remain positive throughout
the winter although their location and magnitude vary
somewhat from month to month (not shown).
3.2 CMLM-G experiment
The CMLM-G exhibits very similar winter SST and surface heat flux responses to those of CMLM-E (Fig. 5).
However, the SLP and the Z250 anomalies from the two
experiments bear some interesting differences. Over the
North Pacific, the negative (positive) SLP (Z250) response
in CMLM-G is approximately half (double) the magnitudeof that in CMLM-E. In addition, the response over the
North Atlantic is diminished considerably and shifted
CMLM−G

CMLM−G minus CMLM−E

AGCM−E

Z250

SLP

Qnet

SST

123

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response
Fig. 6 Monthly equilibrium
responses of (top) CMLM-E
sea-level pressure (middle)
AGCM-E sea-level pressure,and (bottom) AGCM-E SST.
Positive (negative) responses
are plotted with solid (dashed)
contours. Zero contours aresuppressed. Contour intervals
are 1 hPa. Shading indicatessignificance at 95% level

Dec

Jan

Feb

Mar

CMLM−ESLP

AGCM−ESLP

AGCM−E
SST

eastward. To highlight the differences between the two
experiments, the CMLM-G responses minus the corresponding CMLM-E responses are presented in the thirdcolumn of Fig. 5. The differences in the atmospheric circulation responses consist of an equivalent barotropic ridgeover the North Pacific and a more spatially confined
equivalent barotropic low over the North Atlantic, withmaximum anomalies *2 hPa in SLP and *30 m in Z250.
The Z250 difference pattern resembles the Pacific-North
American (PNA) teleconnection pattern (Wallace andGrutzler 1981), although the downstream center over the
subpolar North Atlantic is stronger than typical. The SST
differences exhibit positive anomalies *0.2°C directly
south of the forcing region and a smaller patch of negative
anomalies to the north. These weak SST anomalies appear
to be driven by anomalous atmosphere-to-ocean heat flux
associated with the positive SLP anomaly center. Similar
atmospheric circulation-induced anomalous heat fluxes
also appear to drive the SST differences in the North
Atlantic.
What drives the differences in the circulation responses
in CMLM-E and CMLM-G, in particular the equivalentbarotropic high over the North Pacific and associated PNA
pattern? We argue that the presence of an interactive ocean
mixed layer in the tropics in CMLM-G leads to a tropicalrainfall response which in turn provides an additional
forcing mechanism for the extratropical atmospheric circulation. Figure 7 shows that CMLM-E does not exhibit
any significant tropical rainfall response while CMLM-Gshows negative (positive) rainfall anomalies along theequator (*5°N) in the central and western tropical Pacific,

123o(a) CMLM−E

30 N
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0o
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Fig. 7 DJF responses of precipitation in the tropics for (a) CMLM-E(b) CMLM-G, and (c) CMLM-G minus CMLM-E. Positive (negative)
responses are plotted with solid (dashed) contours. Zero contours aresuppressed. Contour intervals are 0.5 mm day-1. Gray shading
indicates significance at 95% levelwith maximum amplitudes *±1 mm day-1, which corresponds to about 40% of the root-mean-square of the
December–February average tropical precipitation in
CCSM2. The CMLM-G rainfall anomalies are accompanied by weak but significant tropical SST anomalies(*±0.2°C; not shown). Note that the tropical rainfall
anomalies are averaged over December–February to
account for the approximately 1-month delay between
tropical forcing and extra-tropical atmospheric circulation
response (Trenberth et al. 1998; Alexander et al. 2002).

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

The anomalous precipitation pattern in CMLM-G, which
implies a northward migration of the Inter-Tropical Convergence Zone (ITCZ), resembles that of ENSO in its
negative phase (Deser et al. 2006).
3.3 AGCM-E
Next, we compare the CMLM approach with the traditional
specified-SST approach for assessing the atmospheric
response to anomalous SSTs. In AGCM-E, SST anomalies
identical to those in CMLM-E are specified over theforcing region (35°–45°N, 140°E–180°; Fig. 5). The general features of the circulation responses in the two
experiments are similar, with a baroclinic vertical structure
over the North Pacific and an equivalent barotropic structure over the North Atlantic and Siberia. However, theNorth Pacific SLP response is more spatially confined over
the western half of the basin in AGCM-E compared
to CMLM-E. Thus, the SLP anomalies averaged over
the North Pacific basin (20°–60°N, 140°E–120°W) are-0.83 hPa for CMLM-E compared to only -0.26 hPa for
AGCM-E. The weaker January–March average response in
AGCM-E is due to the rapid decay of the response in late
winter, a feature not seen in CMLM-E (Fig. 6). By March,
the amplitude of the SLP anomalies averaged over the
North Pacific basin (20°–60°N, 140°E–120°W) in AGCME is \10% of that in CMLM-E and insignificant except fora small region east of Japan (Fig. 6). In addition to differences in the SLP responses, the Z250 response over the
North Pacific (Atlantic) is larger (smaller) in AGCM-E
compared to CMLM-E. Unlike the SLP response, the Z250
response in AGCM-E does not exhibit a weakening in late
winter over the North Pacific (not shown).
We propose that most of the differences between the
AGCM-E and CMLM-E responses can be attributed to therole of the ocean mixed layer outside of the forcing region.
In particular, AGCM-E exhibits a larger atmosphere-toocean heat flux to the east of the dateline compared to
CMLM-E (and CMLM-G) along 40°N (Fig. 8), because
the fixed SST in AGCM-E acts as an infinite heat reservoir(Barsugli and Battisti 1998). The maximum atmosphere-toocean heat flux anomaly is -37 W m-2 in AGCM-E,
while it is -12 W m-2 and -10 W m-2 in CMLM-E and
CMLM-G, respectively. As a result, the associated negative diabatic heating anomalies east of the dateline in
AGCM-E exhibit greater amplitude than either CMLM
experiment. We conjecture that this anomalous diabatic
cooling to the east of the dateline drives a local positive
SLP response that acts to partially offset the negative SLP
response due to the positive SST/diabatic heating anomalies in the Kuroshio Extension. Consequently, the atmospheric temperature and low-level geopotential height
responses are more confined to the western portion of thebasin in AGCM-E than CMLM-E (Fig. 8). In AGCM-E,
the growth of the SST anomalies in the forcing region from
December through March leads to a corresponding growthin the negative diabatic heating response to the east which
in turn explains the progressive weakening of the negative
SLP response as documented in Fig. 6. It is worth notingthat AGCM-E exhibits insignificant tropical precipitation
responses (not shown).
3.4 Wind stress curl response
Part of the motivation for this study is to better understandthe atmospheric circulation response to SST anomalies
in the Kuroshio Extension region due to its crucial role insustaining the decadal coupled atmosphere–ocean mode in
CCSM2 (Kwon and Deser 2007). Of particular importanceis the wind stress curl component of the response which
drives geostrophic ocean circulation changes that in turn
alter SSTs in the Kuroshio Extension region. The JFM
wind stress curl responses in CMLM-E, CMLM-G and
AGCM-E are shown in Fig. 9; their annual mean counterparts are very similar (not shown). The patterns of windstress curl response are similar in all three experiments,
with comparable amplitudes for CMLM-E and AGCM-E
and weaker magnitudes for CMLM-G, in keeping with the
SLP responses (recall Fig. 5). One may infer from thedifference between the CMLM-G and CMLM-E windstress curl responses that the tropically forced teleconnection partially opposes the direct response (Figs. 9d). Thus,
the thermally coupled tropics act as a weak damping to theNorth Pacific decadal oscillation in CCSM2. The implication of these response patterns for the decadal variability
in CCSM2 will be discussed further in Sect. 4.5.

4 Discussion
4.1 Role of the tropical ocean mixed layer
Previous modeling studies have repeatedly shown that tropics are sensitive to specified extratropical thermal anomalies either introduced directly as temperature or heat flux
anomalies or indirectly as sea-ice or fresh water anomalies(Chiang and Bitz 2005; Broccoli et al. 2006; Kang et al.
2008; Wu et al. 2008; Chiang et al. 2008; Okumura et al.
2009). Our winter atmospheric circulation response toanomalous OHFC in the Kuroshio Extension is substantially modified by the presence of a thermodynamically
active ocean mixed layer in the tropics. When the tropicalocean mixed layer is active, the Pacific ITCZ shifts northward toward the warmed hemisphere, consistent with previous studies (e.g. Chiang and Bitz 2005; Kang et al. 2008).
The ITCZ displacement induces an extratropical PNA-like123

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

Qnet

Pressure (hPa)
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CMLM−E

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Longitude (oE)with solid (dashed) contours. Zero contours are suppressed. Contour
intervals are 5 m, 0.5 K, and 5 9 10-6 K s-1, respectively. Shading
indicates significance at 95% level. (Bottom row) Positive values
indicate the anomalous net surface heat flux is from ocean toatmosphere. Dotted values are significant at the 95% confidence level(a) CMLM−Eo(b) CMLM−G

o

60 N

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atmospheric response, with an equivalent barotropic ridgeover the North Pacific that competes with the locally forced
baroclinic response. The mechanism for the northward
displacement of the Pacific ITCZ in our experiments withan active tropical ocean mixed layer requires further
investigation. A possible explanation is the framework
proposed by Kang et al. (2008) in which atmospheric
meridional energy fluxes change in such a way as to partially compensate the implied oceanic fluxes associatedwith the imposed OHFC anomaly. In addition, the wind-

123

180

Longitude (oE)Fig. 8 JFM equilibrium responses along a zonal section averaged for
30°–50°N of (top row) geopotential height (second row) atmospherictemperature (third row) diabatic heating, and (bottom row) net surface
heat flux from (left) CMLM-E (center) CMLM-G, and (right)
AGCM-E. (Top three rows) Positive (negative) responses are plotted
Fig. 9 JFM wind stress curl
responses from (a) CMLM-E(b) CMLM-G (c) AGCM-E, and(d) CMLM-G minus CMLM-E.
Contour intervals are 2 9 10-8
N m-3. Negative values aredashed and shading indicatessignificance at 95% level

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evaporation-SST feedback mechanism may also be important (Chiang and Bitz 2005). It remains to be seen whether
the inclusion of a dynamically active tropical upper ocean
produces a similar ITCZ response as our mixed layer
tropical ocean.
4.2 Role of the extratropical ocean mixed layer
In late winter, the presence of an active ocean mixed layer
in the extra-tropics leads to a stronger surface low response

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean responseover the North Pacific compared to the case without an
ocean mixed layer. We propose that this is due not to any
influence from the tropics, but because the ‘specified-SST’experiment (AGCM-E) lacks the so-called ‘‘reduced
damping’’ effect due to thermal adjustment between theocean and atmosphere (Barsugli and Battisti 1998). That is,excessive heat loss from the atmosphere to the ocean in the‘specified SST’ experiment acts as a negative feedback to
the direct response not just to the east of the dateline but
also within the forcing region. This result suggests a
potential deficiency inherent in specified-SST experiments.
We note that in the case of perpetual experiments for agiven month (e.g. Peng et al. 1997), the negative feedback
could have ample time to grow and completely offset the
primary response.
Liu and Wu (2004) also considered the lack of ‘‘reduced
damping’’ when comparing their specified SST experimentwith their fully coupled model and specified surface heat
flux experiments. However, in contrast to our results, their
specified SST and coupled model experiments yielded
similar atmospheric circulation responses. We note, however, that their findings are based on ensemble integrations
for early winter (November–December) with constant
amplitude SST forcing, in contrast to our results for late
winter and amplifying SST anomalies.
4.3 Comparison with previous studies
As discussed in the introduction, Yulaeva et al. (2001)
conducted an experiment very similar to CMLM-G.
However, unlike the baroclinic (low-level trough, upperlevel ridge) response found in CMLM-G, Yulaeva et al.
(2001) obtained an equivalent barotropic response consisting of a trough over and downstream of the forcing
region directly east of Japan and a ridge over Alaska andwestern Canada. The possible role of tropical precipitation
changes in modifying the atmospheric circulation responseover the north Pacific was not explored in Yulaeva et al.
(2001), although their experimental design included aninteractive ocean mixed layer in the tropics. This aspect,
coupled with the relatively small sample size (10 ensemble
members compared to *100 here) and different definitionof winter (December–February compared to January–
March here), makes it difficult to ascertain the causes forthe differing responses in Yulaeva et al. (2001) and
CMLM-G. In terms of wind stress curl, the North Pacificresponse in Yulaeva et al. (2001) resembles that in CMLMG (and CMLM-E), with positive anomalies in the latitude
band 30°–50°N which in turn provides a negative feedback
to the prescribed OHFC. It provides a negative feedback
because the positive anomalous wind stress curl response
implies the southward migration of the zero wind stresscurl line accompanied by the southward migration of theKuroshio Extension front 3–5 years later, which causes thenegative OHFC and cold SST anomalies in the Kuroshio
Extension. Thus, their wind stress curl response would
support the existence of a coupled atmosphere–ocean
decadal oscillation as in our experiments, despite differences in the vertical structures of the atmospheric circulation response.
The fully coupled model experiments by Liu and Wu(2004) and Liu et al. (2007) may also have involved
tropical precipitation anomalies but these were notexplicitly examined. Liu et al. (2007) compared the
responses for November–January and February–April andfound an equivalent barotropic ridge response in early
winter and a baroclinic (low-level trough and upper-level
ridge) response in late winter. The monthly evolution of theatmospheric response in our coupled model experiments(CMLM-E and CMLM-G) is largely consistent with that of
Liu et al. (2007) although the equivalent barotropic ridge
response is confined to November with a gradualstrengthening of the baroclinic response from December toMarch.
4.4 Comparison with the CCSM2 responses
Kwon and Deser (2007) empirically diagnosed the windstress curl response to anomalous geostrophic OHFC along
the Kuroshio Extension in CCSM2 by linearly regressing
the anomalous wind stress curl field lagged by 1-year
upon the Kuroshio Extension SST anomaly index. The
1-year lag was used to distinguish the atmospheric response
to SST from the atmospheric forcing of SST which isotherwise confounded in the simultaneous regression. Thisapproach relies on the fact that the wind stress curl
anomalies contain practically no memory from 1 year to
the next, while SST anomalies exhibit high persistencefrom one winter to the next. The CCSM2 JFM SLP and
Z250 responses obtained using this empirical technique areshown in Fig. 10, scaled by the ratio of the standard
deviation of the CCSM2 and CMLM-E SST anomalies
averaged over the Kuroshio Extension (0.7°C/3.0°C; recall
Sect. 3.1). Note that the ENSO influence is removed basedon subtracting linear regressions of SLP and Z250 on thefirst and second principal component time series of
the tropical SST (10°S–10°N, 130°E–90°W) to obtain the
direct responses most relevant to the CMLM-E response.
The pattern and amplitude of the SLP response in CCSM2
is similar to that in CMLM-E, with negative values over
the North Pacific centered along 40°N and positive values
north of 60°N. The Z250 response pattern in CCSM2 alsobears some similarity with that in CMLM-E, with positive
values near the Arctic north of 50°N. However, the negative values over the North Pacific centered along 40°N aremuch stronger and broader. Thus, the extra-tropical123

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response(a) Z250

teleconnections may have a smaller influence on windstress curl than on SLP over the North Pacific, a result that
has implications for decadal variability of the Pacificcoupled atmosphere–ocean system.
4.5 Linear Rossby wave model(b) SLP

Fig. 10 JFM CCSM2 responses to the Kuroshio Extension SST
anomalies for (a) 250 hPa geopotential height, and (b) sea-levelpressure. The responses are calculated after the ENSO influence isremoved based on the linear regression and scaled by the ratio
between the standard deviation of CCSM2 and CMLM-E SST
anomalies averaged in the Kuroshio Extension. Positive (negative)
responses are plotted with solid (dashed) contours. Zero contours aresuppressed. Contour intervals are 10 m and 1 hPa, respectively.
Shading indicates significance at 95% level

response is equivalent barotropic, even over the North
Pacific, unlike that in CMLM-E, CMLM-G or AGCM-E.
Despite the fact that the North Pacific SLP responses aretwice as large in CMLM-E and CCSM2 as in CMLM-G,
the North Pacific wind stress curl response is only 1.2 times
as large (Fig. 11). This is due to the fact that the anomalous
wind stress curl is proportional to the amplitude of theanomalous SLP and inversely proportional to the spatial
scale of the anomalous SLP via the Laplacian operator. The
larger spatial scale of the SLP response in CMLM-G
compared to CMLM-E and CCSM2 partially offsets thedifference in amplitude, yielding a more comparable windstress curl response. This mismatch occurs because theremote component of the North Pacific SLP response
forced by tropical precipitation anomalies has a broader
spatial scale than the local component of the SLP response
forced directly by the Kuroshio Extension SST anomalies.
Thus, due to their larger spatial scale, tropical atmospheric

123

Kwon and Deser (2007) suggested that North Pacific quasidecadal (15–20 year) variability in CCSM2 results fromtwo-way ocean–atmosphere coupling involving the
dynamical adjustment of the ocean gyre circulation via
first-mode baroclinic Rossby waves to basin-scale windstress curl anomalies *3–5 years earlier and a response of
the large-scale wind stress curl field to SST anomalies in
the Kuroshio Extension induced by the ocean gyre circulation changes. Their argument was based on calculations
with a linear Rossby wave model driven by the CCSM2
wind stress curl anomaly pattern that precedes Kuroshio
Extension SST anomalies by 3–5 years (Fig. 12a). The
temporal evolution of the wind stress curl anomaly patterncontains a weak quasi-decadal spectral peak which,although not statistically significant, is crucial for sustaining a robust quasi-decadal spectral peak in SST along theKuroshio Extension. Kwon and Deser (2007) argued thatthe annual mean wind stress curl response pattern to SST
anomalies in the Kuroshio Extension (Fig. 12b), diagnosed
empirically from CCSM2 using lag-regression analysis as
described in the previous section, was sufficiently similar
to that of the forcing pattern (but of opposite sign andweaker amplitude) so as to support the existence of theocean–atmosphere coupled mode.
However, closer inspection reveals that the wind stresscurl response patterns (Fig. 12b, c) exhibit largest amplitudes over the western half of the North Pacific, unlike theforcing pattern which is strongest in the central basin. This
may diminish the amplitude and period of the oscillation
due to the shorter fetch for the westward-propagating
Rossby waves. Therefore, we assess to what extent thequasi-decadal variability can be recovered from the linear
first-mode baroclinic Rossby wave model of Kwon andDeser (2007) when explicitly using the annual-mean windstress curl response patterns identified empirically fromCCSM2 and obtained directly from the CMLM-E experiment. Note both response patterns in Fig. 12 have beenscaled by the ratio of the standard deviations of the CCSM2
and CMLM-E SST anomalies averaged over the Kuroshio
Extension for ease of comparison with the CCSM2 forcing
pattern. The time series associated with the response pattern is obtained by projecting the pattern upon the CCSM2
wind stress curl anomaly in each month and year of the
650-year simulation over the latitude band 30°–50°N. The
response pattern and time series are then added to
the CCSM2 wind stress curl forcing pattern whose timeY.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response
CMLM−E

CMLM−G

CCSM2

SLP

60oN

∇×τ

60oN

o

o

40 N

40 N

o

o

o

20 N ooooo120 E 150 E 180 W 150 W 120 W

o

40 N

40 N
o

(a) CCSM2 Forcing

60 N
o

−0.8

40 N

0.4

0.2

o

o

o

180 W

o

150 W

o

120 W

(b) CCSM2 Response

o

60 N
40oN

0.2

o

20 N o
120 E

150oE

180oW

150oW

120oW

(c) CMLM−E Response

o

60 N
0.2

o

40 N
o

20 N o
120 E

o

150 E

o

o

180 W

o

150 W

40 N

20 N ooooo120 E 150 E 180 W 150 W 120 W

Fig. 11 JFM responses for (top row) sea-level pressure (middle row)
wind stress curl, and (bottom row) Laplacian of sea-level pressure
from (left column) CMLM-E (center column) CMLM-G, and (rightcolumn) CCSM2. The CCSM2 responses are calculated after the
ENSO influence is removed based on the linear regression and scaledby the ratio between the standard deviation of CCSM2 and CMLM-E

150 E

60 N

o

o

20 N o
120 E

o

60 N

o

20 N ooooo120 E 150 E 180 W 150 W 120 Wo

20 N ooooo120 E 150 E 180 W 150 W 120 W

o

60 N

∇ SLP

o

40 N
20 N ooooo120 E 150 E 180 W 150 W 120 W

2

60oN

o

120 W

Fig. 12 Annual mean wind stress curl anomaly pattern of (a) CCSM2
forcing (b) CCSM2 response, and (c) CMLM-E response. Contour
intervals are 0.2 9 10-8 N m-3. Negative values are dashed andshading indicates significance at 95% level. The CCSM2 response iscalculated after the ENSO influence is removed based on the linear
regression. Note that the response patterns are scaled by the ratio
between the standard deviation of CCSM2 and CMLM-E SST
anomalies averaged in the Kuroshio Extensiono

20 N ooooo120 E 150 E 180 W 150 W 120 W

SST anomalies averaged in the Kuroshio Extension. Positive (negative) responses are plotted with solid (dashed) contours. Zero
contours are suppressed in the sea-level pressure plots. Contour
intervals are 1 hPa, 2 9 10-8 N m-3, and 1.5 9 10-12 hPa m-2,
respectively

series has been artificially whitened so that it contains nospectral peaks. The amplitude of the whitened forcing is set
to correspond to the minimum spectral power in the original forcing time series at periods longer than 10 years. It
should be noted that the forcing pattern resembles theleading mode of atmospheric circulation variability over
the North Pacific (see Kwon and Deser 2007).
The calculated sea surface height (SSH) responses(equivalent to SST in the Rossby wave model) in the
Kuroshio Extension reproduces the quasi-decadal spectral
peak in CCSM2 SSH, albeit 20% weaker (Fig. 13). The
white-noise wind stress curl forcing produces a backgroundred-noise response (gray line in Fig. 13b), while the windstress curl response to SST anomalies in the Kuroshio
Extension generates the additional quasi-decadal spectral
peak. The multi-decadal (*40 years) peak is considerably
damped in the Rossby wave model hindcast, suggestingthat this peak is generated by processes other than firstmode baroclinic Rossby wave adjustment along the latitude
band of the Kuroshio Extension, for example higher-mode
Rossby waves or first-mode baroclinic Rossby wave
adjustment at higher latitudes (Zhong and Liu 2009). The
weaker amplitude of the quasi-decadal spectral peak is dueto the presence of negative wind stress curl anomalies tothe north and south of the Kuroshio Extension in the
CCSM2 and CMLM-E response patterns: when these are
artificially damped, the Rossby wave model hindcast produces a spectral peak that is comparable to that in CCSM2
SSH (not shown). Therefore, the wind stress response
patterns, albeit weaker and centered farther west comparedto the forcing pattern, do maintain a quasi-decadal coupledoscillation in the North Pacific.

123

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response
Period (year)
Inf

50

Power (m2/cpy)

0.02

25

16.6

12.5

10

8.3

7.1

6.2

5.5

(a)

5
AR1
95%

0.015
0.01
0.005
0
0

0.02

0.04

0.06

0.08

0.1

0.12

0.14

0.16

0.18

0.2

Frequency (cycles/year)
Period (year)
Inf

50

25

16.6

12.5

10

8.3

7.1

6.2

5.5

5

0.02

0.04

0.06

0.08

0.1

0.12

0.14

0.16

0.18

0.2

Power (m2/cpy)

0.02

(b)
0.015
0.01
0.005
0
0

Frequency (cycles/year)
Period (year)
Inf

50

25

16.6

12.5

10

8.3

7.1

6.2

5.5

5

0.02

0.04

0.06

0.08

0.1

0.12

0.14

0.16

0.18

0.2

Power (m2/cpy)

0.02

(c)
0.015
0.01
0.005
0
0

Frequency (cycles/year)

Fig. 13 Power spectrum of winter (DJFM) SSH anomalies averaged
over 34°–43°N, 140°E–180° from (a) the CCSM2, and the Rossby
wave model hindcast driven by (b) combination of CCSM2 windstress curl response and whitened forcing patterns, and (c) combination of CMLM-E wind stress curl response and whitened forcing
patterns. The gray lines in (b) and (c) are the Rossby wave model
hindcast driven by the whitened forcing pattern alone. Note that (a) isreproduced from the Fig. 13b of Kwon and Deser (2007). The best-fit
AR1 spectrum and 95% significant level with respect to the AR1
spectrum are plotted with dashed lines in (a)5 Conclusions
The wintertime (JFM) response of the coupled atmosphere–ocean mixed layer system to geostrophic OHFC
anomalies in the Kuroshio Extension has been examined bymeans of experiments with an AGCM coupled to an
entraining ocean mixed layer model. In addition to thedirect response, interaction with the tropics and the role ofthe extra-tropical ocean mixed layer were addressed
through modifications of the experimental design.
The direct coupled ocean mixed layer–atmospheric circulation response to a positive OHFC anomaly in the
Kuroshio Extension (experiment CMLM-E) consists ofpositive SST anomalies along the Kuroshio Extension anda baroclinic atmospheric response characterized by a lowlevel trough and upper-level ridge over the North Pacific.
For an average SST anomaly of 3°C within the Kuroshio

123

Extension region, the response amplitudes over the North
Pacific are approximately 2 hPa for SLP and 10–20 m for
Z250. In addition to the North Pacific response, an equivalent barotropic ridge response is found over the North
Atlantic centered at 50°N, with amplitudes *2 hPa for
SLP and 40 m for Z250. These atmospheric circulation
anomalies drive a weak (0.2–0.7°C) SST warming of theNorth Atlantic via perturbations to the surface turbulent
energy flux.
When the ocean mixed layer is thermodynamicallycoupled to the atmosphere in the tropics as well as theextra-tropics (experiment CMLM-G), the tropical Pacific
ITCZ shifts northward in response to oceanic warming
along the Kuroshio Extension. The meridional displacement of the ITCZ generates a PNA-like atmospheric
response that includes an equivalent barotropic ridge over
the North Pacific. This tropically induced teleconnection
modifies the direct coupled atmospheric circulation
response, reducing the surface trough and enhancing theupper level ridge over the North Pacific.
The uncoupled atmospheric circulation response to a
positive SST anomaly along the Kuroshio Extension(experiment AGCM-E) is generally similar to the coupled
response, except that the anomalous SLP trough over the
North Pacific is confined to the western half of the basin in
AGCM-E. This sensitivity of the North Pacific SLP
response to the presence of an interactive ocean mixed
layer was suggested to result from the artificial enhancement of the anomalous atmosphere-to-ocean net surface
heat flux downstream of the Kuroshio Extension in
AGCM-E (e.g., lack of reduced thermal damping) compared to CMLM-E. The atmospheric cooling that resultsfrom the net surface heat flux response drives a positive
SLP anomaly that partially offsets the negative SLP
anomaly that is a direct response to the warming along theKuroshio Extension. This effect increases over time, suchthat by March the SLP trough response over the North
Pacific is nearly absent in AGCM-E. This result may have
implications for SST-forced AGCM experiments conducted in ‘‘perpetual mode’’ for which the month or season
is kept fixed.
The AGCM, CMLM, and CCSM2 experiments are
similar in terms of their wind stress curl responses despite
differences in their SLP responses. This is due to that thetropically-induced portion of the SLP responses in CMLMG has a larger spatial scale than the locally-induced SLP
response, a factor which affects the wind stress curl via the
Laplacian operator. This result has implications for decadal
variability of the Pacific coupled ocean–atmosphere system
to the extent that the North Pacific ocean gyre circulation is
primarily driven by wind stress curl fluctuations.
The linear Rossby wave model driven by the combination of a wind stress curl response taken from CMLM-E

Y.-O. Kwon et al.: Coupled atmosphere–mixed layer ocean response

and a white noise wind stress curl forcing pattern takenfrom CCSM2 (which resembles the leading EOF over the
North Pacific) reproduces the quasi-decadal (15–20 year)
oscillation evident in CCSM2. This result supports theconclusion of Kwon and Deser (2007) that the quasi-decadal variability in CCSM2 owes its existence primarily to
two-way atmosphere–ocean coupling in the North Pacific.
There is some suggestion that this mechanism may also be
operating in nature (Qiu et al. 2007), and work is currently
underway to investigate this issue further.
Acknowledgments We wish to acknowledge helpful discussions
and technical support from many colleagues including D. Ferreira,J. Goodman, A. Phillips, and J. Yin. We gratefully acknowledge
financial support from NOAA’s Office of Global Programs (grant to
C. Deser and Y.-O. Kwon). Y.-O. Kwon is also supported through the
Claudia Heyman Fellowship of the WHOI Ocean Climate Change
Institute.

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