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Evidence for a Wind-Driven Intensification of the Kuroshio Current Extension
from the 1970s to the 1980s
CLARA DESER
National Center for Atmospheric Research,* Boulder, Colorado

MICHAEL A. ALEXANDER AND MICHAEL S. TIMLINCIRES, Climate Diagnostics Center, University of Colorado, Boulder, Colorado(Manuscript received 31 December 1997, in final form 2 July 1998)
ABSTRACTThe spatial and temporal characteristics of oceanic thermal variations in the mixed layer and main thermoclineof the midlatitude North Pacific are distinctive, suggesting different physical origins. Within the main thermocline(400-m depth), the variability is dominated by a westward-intensified pattern of decadal scale, indicative ofenhanced eastward geostrophic flow along the southern flank of the Kuroshio Current extension during the 1980s
relative to the 1970s. The authors argue that the decadal-scale change in the strength of the Kuroshio extension
was a result of the dynamical adjustment of the oceanic circulation to a decadal variation in wind stress curl
according to Sverdrup theory. Four-times daily wind stress fields from the National Center for AtmosphericResearch–National Centers for Environmental Prediction reanalysis project are used to compute the decadalchange in Sverdrup transport associated with the 1976/77 climate transition. It is shown that the decadal changes
in Sverdrup transport inferred from the wind stress curl field and in observed geostrophic flow inferred fromthe upper-ocean thermal field are consistent both in terms of spatial pattern and magnitude. The decadal change
in depth-averaged geostrophic transport along the Kuroshio extension (referenced to 1 km) is 11.6 Sv, similar
to the Sverdrup transport change (11.5–13.9 Sv). The decadal-scale thermocline variation along the westernboundary between 308 and 408N exhibits a lag of approximately 4–5 yr relative to the decadal variation in thebasin-wide wind stress curl pattern. This delay may be indicative of the transient adjustment of the gyre-scale
circulation to a change in wind stress curl via long baroclinic Rossby waves.

1. Introduction
There is increasing evidence for coherent decadalscale fluctuations of the ocean–atmosphere system in
the North Pacific (see Mantua et al. 1997 and Nakamura
et al. 1997 for recent reviews). A particularly prominent
example occurred from approximately 1977 to 1988,
when sea surface temperatures (SSTs) were below normal across the central and western North Pacific andabove normal along the west coast of North America.
The SST anomalies were accompanied by a stronger
than normal Aleutian low pressure system and associated midlatitude westerly winds, particularly during thewinter half of the year (Fig. 1). Epipelagic marine eco-

* The National Center for Atmospheric Research is sponsored by
the National Science Foundation.
Corresponding author address: Dr. Clara Deser, Climate and Global Dynamics Division, NCAR, P.O. Box 3000, Boulder, CO 803073000.
E-mail: cdeser@ucar.edu

q 1999 American Meteorological Society

sytems, including salmon production, were affected by
the decadal climate change (Polovina et al. 1995; Brodeur et al. 1996; Mantua et al. 1997).
Several mechanisms have been proposed to explain
the occurence of decadal-scale climate variations in the
North Pacific. Decadal changes in tropical Pacific rainfall may force atmospheric teleconnections over the
North Pacific (cf. Nitta and Yamada 1989; Trenberth1990; Graham et al. 1994; Zhang et al. 1997), which inturn alter SSTs locally via turbulent energy fluxes, Ekman current advection, and vertical mixing (Alexander
1992; Luksch and von Storch 1992; Miller et al. 1994).
Local extratropical air–sea interaction may amplify thetropically induced decadal signal over the North Pacific,
as suggested by Trenberth (1990), Trenberth and Hurrell(1994), and Yukimoto et al. (1996). Alternatively, thedynamical adjustment of the extratropical ocean circulation via long baroclinic Rossby waves to stochastic
wind stress forcing may lead to SST variability on decadal timescales, independent of the Tropics (Frankignoul et al. 1997; Jin 1997). Latif and Barnett (1994; seealso Robertson 1996 and Jin 1997) argue that coupled
extratropical air–sea interaction involving ocean gyre

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FIG. 1. Time series of winter (Nov–Mar) SLP (mb) averaged overthe central North Pacific (308–658N, 1608E–1408W). The solid circles
denote the raw winter SLP anomalies, and the shaded curve the anomalies smoothed with a three-point binomial filter. This record is areliable indicator of the leading mode of wintertime atmospheric circulation variability over the North Pacific. After Trenberth and Hurrell(1994).

dynamics may result in self-sustained decadal-scale oscillations of the North Pacific climate system.
Clearly, it is important to extend the observationaldescription of decadal-scale variability in the North Pacific from the sea surface to the ocean interior. Does thelarge-scale oceanic circulation exhibit decadal fluctuations, and if so, to what extent are the changes winddriven? Recent analyses of upper-ocean thermal structure have begun to address these questions. Deser et al.
(1996) and Miller et al. (1998) documented a westwardintensified pattern of decadal-scale thermal anomalies
in the main pycnocline (400-m depth) in midlatitudes.
Using thermal wind arguments, they inferred that theeastward geostrophic flow of the Kuroshio Current extension was enhanced during the 1980s relative to the1970s (see also Yasuda and Hanawa 1997). A similardecadal-scale increase in eastward transport of the Kuroshio Current south of Japan was reported by Qiu andJoyce (1992). The decadal transport variations of theKuroshio Current system were interpreted by Miller etal. (1998) as a response to large-scale wind forcing
based on experiments with an ocean general circulation
model, a result also anticipated by Trenberth (1990),
Qiu and Joyce (1992), and Yasuda and Hanawa (1997).
Wind forcing was also implicated in the observed decadal-scale change of the Alaska gyre circulation reported by Lagerloef (1995).
A more passive role for the oceanic circulation was
indicated by Watanabe and Mizuno (1994), Deser et al.
(1996), Lysne et al. (1997), and Schneider et al. (1999),
who showed that thermal anomalies originating in the
deep winter mixed layer in the central North Pacific canbecome subducted into the upper portion of the main
thermocline and subsequently transported downward
and equatorward along isopycnal surfaces by the mean
subtropical gyre circulation. A warm anomaly during
the 1970s and a cold anomaly during the 1980s wereadvected by the mean circulation from the central North
Pacific to approximately 188N in the western Pacificover a period of approximately 8 yr, following closely
the model of the ventilated thermocline (Schneider etal. 1999). Elements of the ‘‘ventilated thermocline’’ sig-

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nal can also be found in Zhang and Levitus (1997), whoshowed that at 250-m depth thermal anomalies appear
to rotate clockwise around the subtropical gyre over thecourse of approximately a decade.
The simplest theory of the steady-state response of
the large-scale oceanic circulation to wind stress forcing
is the Sverdrup relation (Sverdrup 1947). Simply stated,
the Sverdrup relation is a linear vorticity balance between wind stress curl and meridional advection of planetary vorticity by interior flows; zonal currents are inferred from mass continuity considerations. Thus, thelarge-scale distribution of wind stress curl determines
the gyre-scale patterns of the oceanic circulation. Notethat Sverdrup balance does not account for the strong
western boundary currents that complete the gyre circulations.
There is considerable debate over the extent to which
Sverdrup balance governs the observed large-scale oceanic circulation, particularly in the Atlantic sector, given
the neglect of important processes such as interactionof the flow with bottom topography, thermohaline effects, and inertial and recirculation components (cf.
Wunsch and Roemmich 1985). However, despite its simplicity, Sverdrup balance appears to account quantitatively for much of the large-scale circulation in the
North Pacific (cf. Niiler and Koblinsky 1985; Qiu andJoyce 1992; Hautala et al. 1994; Kagimoto and Yamagata 1997).
Miller et al. (1998) invoked Sverdrup theory to explain the decadal strengthening of the Kuroshio Current
system in an ocean general circulation model forced
with observed wind stress and heat flux anomalies from1970 to 1988. The purpose of this study is to provide
additional observational support for the hypothesis thatthe decadal-scale increase in the transport of the Kuroshio extension was a result of large-scale wind stressforcing according to simple Sverdrup dynamics. In particular, we shall demonstrate that the decadal-scalechanges in observed geostrophic flow inferred from theupper-ocean thermal field and Sverdrup transport inferred from the wind stress curl field are similar in boththeir spatial pattern and magnitude.
2. Data and methods
a. Ocean temperature dataset
The subsurface ocean temperature profiles used in thisstudy are from the World Ocean Atlas 1994 (WOA94;
Levitus and Boyer 1994). These data were recovered aspart of a recent and extensive search through international archives and then quality controlled in a consistent manner (Boyer and Levitus 1994). Here we examine
thermal variations in the upper 450 m of the North Pacific (208–558N) from 1968 to 1991 using data fromthree sources (mechanical bathythermographs, expendable bathythermographs, and station data consisting primarily of Nansen bottles); however, expendable bathy-

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DESER ET AL.

thermographs (XBTs) account for the majority of theobservations during this time period (see Levitus andBoyer 1994; Deser et al. 1996). Our choice of study
period was dictated by the density of subsurface temperature profiles available from the WOA94, and thelower limit for our analysis (450 m) reflects the deepest
level routinely sampled by the XBTs.
Characteristics of the spatial and temporal distributions of the data may be found in Levitus and Boyer(1994) and Deser et al. (1996). Here we note simply
that the subsurface archive contains an order of magnitude fewer observations than surface marine datasets
such as the Comprehensive Ocean–Atmosphere Data
Set (COADS; Woodruff et al. 1987). However, as discussed in Deser et al. (1996), the relatively sparse sampling of the subsurface network appears adequate nevertheless for the study of decadal-scale variations in thenorth Pacific.
As in Deser et al. (1996), we formed monthly anomalies on a 18 3 18 grid at each standard level from thesurface to 400 m as departures from the Levitus andBoyer (1994) objectively analyzed monthly 18 3 18 climatology. The 18 3 18 anomalies were then averaged
into 28 lat 3 48 long grid boxes. We have made no
attempt to fill in missing data. Further information concerning data distribution, gridding procedures, and validation techniques may be found in Timlin et al. (1995)
and Deser et al. (1996).
We shall also make use of thermal profiles extending
to 1-km depth in the far western Pacific, obtained from
Nansen bottle casts.
b. Wind stress dataset
The surface wind stress data are from the National
Center for Atmospheric Research–National Centers forEnvironmental Prediction (NCEP–NCAR) reanalysis
project (Kalnay et al. 1996) for the period 1968–91(1968 was the earliest year available at the time thisstudy was undertaken). The reanalysis project uses a
fixed state-of-the art global data assimilation system anda comprehensive observational database to derive dynamically consistent atmospheric fields. As such, it represents one of the most complete and quality controlled
archives of atmospheric data currently available.
Wind stress is computed four times daily on a T62
Gaussian grid (roughly equivalent to 28 lat 3 28 long)using a bulk aerodynamic formula in which the dragcoefficient depends on both wind speed and static stability according to Businger et al. (1971). For further
information regarding the boundary layer scheme in theassimilation model, the reader is referred to Kanamitsu(1989).
Other studies rely on purely observational archives
such as the COADS to derive surface wind stresses over
the ocean. However, spurious trends in COADS windspeeds due to changes in observing practices (cf. Cardone et al. 1990; Cayan 1992; Ward and Hoskins 1996),

1699

FIG. 2. Leading principal component of annual subsurface temperature anomalies in the North Pacific. Units are standard deviations.
(a) 50 m. (b) 400 m.

coupled with the relatively sparse spatial and temporal
sampling of the ship-of-opportunity network, led us toadopt the reanalysis wind stress product for our study.
c. EOF analysis
To define the dominant structures of variability in thesubsurface thermal field, we apply empirical orthogonal
function (EOF) analysis to the covariance matrix of annual anomalies on a 48 3 48 grid based on the period
1968–91. Annual anomalies are defined as January–December averages of the individual monthly departuresfrom the monthly long-term means. The domain for the
EOF analyses is 248–568N, 1508E–1208W. The principal
component time series are obtained by projecting the
spatial EOF patterns onto the original anomaly fields
for each year.
3. Resultsa. Upper-ocean thermal variability
Figure 2 shows the leading principal component timeseries of annual temperature anomalies at 50- and 400-m
depth, which account for 37% and 26% of the variance,
respectively. The 50-m record represents thermal variations within the winter mixed layer (summer mixed
layer depths are less than 20–30 m), while the 400-m
record is indicative of conditions in the permanent thermocline below the winter mixed layer. The leading principal component at 50 m (Fig. 2a) is similar to the wintersea level pressure (SLP) time series shown in Fig. 1:
the correlation between them is 0.82 (0.90 after smoothing with a three-point binomial filter). Note, however,1700

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FIG. 3. (a) Decadal-scale change in annual temperature at 50 m
(thick contours; negative values are dashed), formed by subtracting
the period 1968–76 from the period 1977–88. The contour interval
is 0.38C, and the zero contour has been omitted. The thin contours
denote the long-term mean temperature field. (b) As above but for
400 m, formed by subtracting the period 1970–80 from the period
1982–90.

that the decadal signal is more prominent in the oceanic
thermal field than in SLP (relative to interannual variations). Decadal-scale changes are even more pronounced (relative to interannual changes) at 400 m (Fig.
2b) than at 50 m. The decadal variation at 400 m isdelayed relative to that at 50 m, with the transition frompositive to negative anomalies occuring in 1981/82 rather than 1976/77.
On the basis of the principal component time series,
we identified a warm period (1968–76 at 50 m and
1970–80 at 400 m) and a cold period (1977–88 at 50
m and 1982–90 at 400 m). Figure 3 shows the spatial
patterns of the difference between the cold and warm
periods (these maps resemble closely the EOF patterns
upon which they are based). The decadal-scale change
at 50 m (Fig. 3a) exhibits a basin-wide pattern, withnegative anomalies in excess of 20.68C in the central
and western Pacific, and positive anomalies of similar
magnitude along the west coast of North America. Thispattern closely resembles the decadal change of SST in
winter (not shown, but see comparable figures in Trenberth and Hurrell 1994; Deser et al. 1996; Zhang et al.
1997; Nakamura et al. 1997, among others). In contrast,
the decadal-scale change at 400 m (Fig. 3b) exhibits awestward-intensified pattern with the largest cooling (inexcess of 20.68C) centered along approximately 358N
and weaker warming directly south of Japan. The strongest cooling lies along the axis of maximum meridionaltemperature gradient that marks the path of the Kuroshio

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Current extension. The regional warming south of Japan
may reflect a weaker Kuroshio meander after 1982 asshown in Qiu and Joyce (1992). A similar westwardintensified pattern of 400-m thermal variability was
shown by Miller et al. (1998) based upon extended EOF
analysis, a technique that accomodates propagating aswell as stationary signals. The negative anomaly centered near 258N and 1508W may be associated with
subduction as documented by Deser et al. (1996) and
Schneider et al. (1999). Note the absence of positive
temperature anomalies in the eastern third of the basin
at 400 m.
The distinctive spatial and temporal signatures of thedecadal-scale temperature changes in the mixed layer
and the main thermocline suggests that they are controlled by different physical processes. Ocean modelingstudies have shown that the basin-scale pattern of thermal variability within the winter mixed layer is largelya result of anomalous fluxes of latent and sensible energy at the sea surface due to changes in atmosphericcirculation (cf. Alexander 1992; Luksch and von Storch
1992; Miller et al. 1994), although wind-induced anomalous advection by Ekman currents and entrainment may
also contribute in certain locations. The large simultaneous correlation between the 50-m temperature principal component and the winter Aleutian low SLP index
is consistent with the rapid (order 1 month) thermal
response of the winter mixed layer to anomalous turbulent energy fluxes (Frankignoul 1985; Deser and Timlin 1997). The main thermocline, on the other hand,isolated from direct thermodynamic contact with theatmosphere, adjusts to changes in atmospheric circulation via dynamical processes. We proceed next to test
the hypothesis that the decadal-scale change of temperature at 400 m resulted from the dynamical adjustment of the oceanic circulation to changes in wind stresscurl according to Sverdrup’s (1947) theory.
b. Sverdrup balance
Sverdrup balance may be derived from the vertical
integral of the linear vorticity equation,by 5 f]w,]zwhere y is the meridional velocity, f is the Coriolis
parameter, b is the latitudinal gradient of the Coriolis
parameter, and w is the vertical velocity. Integrating thisequation from a depth, z o , where the vertical velocity
is assumed to vanish, to the base of the surface Ekman
layer where the vertical velocity is assumed to result
from wind-driven convergence or divergence in the surface Ekman layer, one obtains the familiar form of Sverdrup balance:

bVg 5

12

t
f
k·= 3
,
ro
f

(1)

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1701

where V g is the vertically integrated meridional geostrophic transport, r o (51025 kg m23 ) is the density ofseawater, k is the unit vector in the vertical direction,t 5 (t l , t f ) is the wind stress vector, and = is thegradient operator in spherical coordinates. Equation (1)states that the input of relative vorticity to the ocean by
the wind stress curl will result in water parcels moving
northward in order to conserve their total (relative 1
planetary) vorticity. Note that the right-hand side of Eq.
(1) is the Sverdrup transport [(1/r o )k · = 3 t ] minus
the Ekman transport [2(b /r o f ) · t l ]. For the decadalscale changes considered here, the Ekman term is lessthan 3% of the Sverdrup term.
If one defines a two-dimensional streamfunction (c g )for the depth-integrated geostrophic velocity,
Vg 5 2

]c g
,
]x

Ug 5

]c g
,
]y

then Eq. (1) can be rewritten as

cg (x) 5 2

f
bro

E 1
x

k· = 3

xE

2

t
dx9,
f

(2)

where x E represents the eastern boundary of the ocean
basin where we have assumed c g (x E ) 5 0. Thus according to Sverdrup theory, the geostrophic transport isrelated to the zonal integral of the wind stress curl fromthe eastern boundary. Note that Eq. (2) is evaluated ateach latitude separately.
c. Wind stress curl forcing
Next we derive the decadal change in c g from theobserved wind stress curl field and compare it to thedecadal change in the observed geostrophic flow inferred from the depth-averaged thermal field. Since annual mean wind stresses drive the large-scale oceanic
circulation on decadal timescales, we first examine theleading mode of variability of annual wind stress anomalies. The leading principal component (not shown),
which accounts for 48% of the variance, is very similar
to the record of winter SLP in the Aleutian low (Fig.
1): the correlation between them is 0.92 (0.95 aftersmoothing with a three-point binomial filter).
Figure 4 shows the decadal-scale change in annual
wind stress and its curl, formed by subtracting the period
1968–76 from the period 1977–88. The mean annual
wind stress and curl fields are also shown for reference.
Decadal westerly wind stress anomalies cover much of
the basin, with maximum values (approximately 0.02 N
m22 or 30%–60% of the mean) between 308 and 408N,
at the southern flank of the mean westerly wind stressbelt (see Fig. 4, top). The strongest decadal westerly
wind stress anomalies are straddled by a north–southcurl dipole between approximately 258 and 408N. Thisanomalous curl dipole extends across nearly two-thirdsof the basin, from 1508E to 1408W, with maximum values of 2 2 4 3 1028 N m23 . Because of its zonal elon-

FIG. 4. (a) Long-term mean annual wind stress (vectors; N m22 )and its curl (contours; 1028 N m23 ) from the NCEP–NCAR reanalyses
for the period 1968–91. The contour interval is 5 3 1028 N m23 , thezero curl line is darkened, and negative values are dashed. The scale
for the vectors is indicated at the lower right. (b) Decadal-scale change
in annual wind stress (vectors) and its curl (contours), formed bysubtracting the period 1968–76 from the period 1977–88, based on
the NCEP–NCAR reanalyses. The contour interval is 2 3 1028 N
m23 and values greater than 2 3 1028 N m23 (less than 22 3 1028
N m23 ) are indicated with light (dark) shading. The 21 3 1028 N
m23 contour is also shown for clarity. Note that the wind vector scaleand contour interval differ from those shown in (a).

gation, the curl dipole has important consequences for
the Sverdrup transport (see below). Note that the anomalous curl dipole is located south of the climatologicalzero wind stress curl line (see Fig. 4a) that delineates
the northern boundary of the mean oceanic subtropicalgyre.
Integrating the curl field westward from the coast of
North America according to (2) yields the geostrophic(Sverdrup minus Ekman) transport streamfunction (Fig.
5). Note that the direction of the transport is parallel tocontours of the streamfunction field, with the strength
inversely proportional to the contour spacing. The climatological mean Sverdrup (minus Ekman) transportshows the familiar pattern of anticyclonic flow aroundthe subtropical gyre and cyclonic flow around the subpolar gyre (Fig. 5a). The maximum strength of the subtropical (subpolar) gyre along the western boundary is
43 Sv (1 Sv 5 10 9 kg s21 ) at 308N (44 Sv at 568N).
Similar mean Sverdrup transport fields were shown by

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FIG. 6. Decadal-scale change in annual geostrophic (Sverdrup minus Ekman) transport streamfunction (thick black contours, takenfrom Fig. 5) superimposed upon the decadal-scale change in verticallyintegrated temperature from 0 to 450 m (color shading). The decadalscale thermal anomalies were formed by subtracting the period 1970–
80 from the period 1982–90.

FIG. 5. (a) Long-term mean annual geostrophic (Sverdrup minus
Ekman) transport streamfunction. The contour interval is 10 Sv, withnegative values dashed. (b) Decadal-scale change in annual geostrophic (Sverdrup minus Ekman) transport streamfunction, formedby subtracting the period 1968–76 from the period 1977–88. The
contour interval is 2 Sv, with negative values dashed. Values greaterthan 4 Sv (less than 24 Sv) are indicated with dark (light) shading.

Trenberth et al. (1990) and Yasuda and Hanawa (1997),
although our values for the subtropical gyre are slightly
lower than those in Trenberth et al.
Like the mean Sverdrup transport distribution, thedecadal difference is westward intensified, with maximum positive anomalies of 6 Sv at 308N and maximum
negative anomalies of 28 Sv at 388N along the westernboundary (Fig. 5b). Enhanced eastward transport is indicated between 308 and 388N, along the southern flankof the region of climatological mean eastward flow (Fig.
5a). The decadal strengthening of the subtropical gyre,as measured by the maximum streamfunction value at
308N, is approximately 15%. Similar decadal-scale differences in Sverdrup transport were shown by Trenberth(1991), Yasuda and Hanawa (1997), and Miller et al.
(1998).
d. Comparison of Sverdrup and observed geostrophic
flow
How well do the geostrophic (Sverdrup minus Ekman) transport streamfunction (c g ) differences matchthe observed upper-ocean thermal changes? Note that a
low in c g should correspond to an upward displacementof the thermocline (low heat content) resulting frompositive wind stress curl, and a high to a depressed
thermocline (high heat content) resulting from negative

wind stress curl. Figure 6 shows the decadal-scalechange in c g superimposed upon the decadal change invertically integrated temperature from 0 to 450 m. Notethat the thermal difference is lagged relative to thestreamfunction change in accord with the results shownin Figs. 1–3. The two fields are in good spatial agreement west of 1708W, with below-normal heat content
corresponding to below-normal streamfunction centeredalong 388N, and above-normal heat content corresponding to above-normal streamfunction centered along
308N east of 1508E. (The peak thermal anomaly westof 1458E is displaced slightly north of the corresponding
high in c g .) Both fields indicate enhanced eastward flow
between 308 and 388N, using thermal wind arguments
to infer geostrophic flow from the thermal field. The
negative c g anomaly center near 478N, 1508E has nocounterpart in the thermal field. Conversely, the negative thermal anomalies extending from 258N, 1808
northeastward to 348N, 1508W do not correspond withthe c g field and, as mentioned above, may result fromanomalous subduction.
The spatial structures of the decadal anomalies in heat
content and c g are compared further in Fig. 7, whichshows latitudinal profiles of the two fields averaged
across the western Pacific (1428–1758E). The meridional
structure of the two fields is similar, with negative values
between approximately 368 and 428N, and positive values between approximately 288 and 328N. However, in
addition to a small constant offset between the two quantities, the positive streamfunction anomalies between
approximately 258 and 338N are enhanced relative tothe thermal anomalies. As shown below, the reduced
amplitude of the heat content anomalies relative to thestreamfunction changes south of 338N reflects that thethermal anomalies extend well below 450 m in this region.
To estimate quantitatively the geostrophic tranportfrom the heat content observations, the vertical structureof the thermal changes down to a ‘‘level of no motion’’is required. The studies of Qiu and Joyce (1992) andJUNE 1999

DESER ET AL.

FIG. 7. Latitudinal structure of the decadal-scale change in geostrophic (Sverdrup minus Ekman) transport streamfunction (solid
curve) and vertically integrated temperature from 0 to 450 m (dashedcurve) averaged across the western half of the Pacific (1428–1768E).

Huang and Russell (1994) suggest that this level is approximately 1–1.5 km for the subtropical gyre. To extend the decadal-scale thermal changes down to 1 km,
we used the station data (Nansen bottle) archives (on a
18 3 18 grid) from the WOA94. The sparsity of Nansen
bottle observations precludes basin-wide maps of thedecadal changes below 500 m. Fortunately, the data
density is highest in the far western Pacific (258–438N,
1428–1508E), where the decadal-scale thermal variations
are largest. However, the results should be consideredpreliminary until more observations below 500 m are
obtained.
Figure 8 shows the vertical structure of the decadalscale thermal changes in the upper kilometer for the far
western Pacific as a function of latitude, superimposed
upon the climatological mean temperature distribution.
The thermal anomalies exhibit little tilt with depth andreach a maximum in the main thermocline. Thus, thelargest positive anomaly (1.28C), located at 328N, iscentered around 500–600-m depth, well below the maximum level sampled by the XBTs, whereas the the largest negative anomaly (21.88C), located at 378N, is situated considerably shallower, near 300-m depth.
Using the profiles shown in Fig. 8 and climatologicalsalinity profiles from Levitus and Boyer (1994), we
computed the decadal change in the depth-averaged geo-

1703

FIG. 8. Latitude–depth cross section of the decadal-scale temperature change in the far western Pacific (1428–1508E; thick contours)
superimposed upon the long-term mean thermal field (thin contours).
The contour interval for the decadal differences is 0.38C, with negative values dashed. The contour interval for the mean field is 28C.

strophic current (referenced to 1 km) between 328 and
378N (the extrema in the vertically integrated thermal
anomaly field). The resulting decadal change in the geostrophic transport is 11.6 Sv. For comparison, the decadal change in geostrophic transport inferred from thewind stress curl field is 11.5 Sv between 328 and 378N
(13.9 Sv between 31.58 and 398N, the extrema in the
Sverdrup transport field). We note that different windstress products may yield different estimates for the
Sverdrup-minus-Ekman transport; however, decadalchanges in the transport should be relatively insensitive
to the wind stress formulation, barring any systematic
trends.
Our estimate of the decadal change in geostrophic
zonal transport east of Japan may be compared withflow variations upstream in the Kuroshio south of Japan
given in Qiu and Joyce (1992). Qiu and Joyce usedbiannual hydrographic surveys along 1378E conductedby the Japanese Meteorological Agency from 1967 to
1988 to compute the geostrophic transport of the Kuroshio, defined as the total eastward flow north of 308N
referenced to a depth of 1.25 km. Their Kuroshio transport record, shown in Fig. 9, displays a gradual increase
from the mid-1970s to the early 1980s. The decadalchange from 1970–80 to 1982–88 is 10.8 Sv, similar to
our estimate for the Kuroshio Extension.
It is instructive to view the decadal changes not only

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FIG. 9. Time series of geostrophic eastward transport Sverdrups(Sv) of the Kuroshio Current south of Japan referenced to 1.25 km,
taken from Qiu and Joyce (1992). The thin horizontal line denotes
the mean transport during 1969–88 (51.7 Sv).

in terms of anomalies from the climatological mean, butalso in terms of the total flow. Figure 10a shows thelatitudinal structure of the total geostrophic zonal flow
averaged between the surface and 1 km in the far western Pacific (1428–1508E) for the periods 1970–80 and
1982–90. Analogous profiles for the total zonal windstress averaged across the Pacific (1508E–1408W) forthe periods 1968–76 and 1977–88 are shown in Fig.
10b. The midlatitude eastward flows in both the ocean
and atmosphere intensified along their southern flanks
from the 1970s to the 1980s, with the largest increases(38% of the mean) at 358N in both fields.
4. Discussion
In agreement with Miller et al. (1998), our results
suggest that the recent decadal-scale change in annual
wind stress curl over the North Pacific altered thestrength of the Kuroshio Current extension in a mannerconsistent with Sverdrup dynamics. Compared to the1970s, the 1980s were characterized by an equatorward
intensification of the mean midlatitude eastward flows
in both the atmosphere and ocean.
The dynamical adjustment of the midlatitude oceanic
gyre circulation to a decadal-scale change in wind stresscurl is accomplished largely by westward-propagating
first-mode baroclinic Rossby waves (cf. Anderson and
Gill 1975). At low frequencies and long wavelengths,
these Rossby waves are zonally nondispersive. If theduration of the anomalous wind stress curl forcing isgreater than the time it takes the forced Rossby wave
to cross the basin, then the gyre circulation will be inquasi-equilibrium with the winds, a state described by
Sverdrup balance. For the case examined here, the winds
are predominantly stronger than normal for about 11 yr(1977–88; recall Fig. 1) and predominantly weaker thannormal for at least 15 yr (1962–76; not shown, but see
the expanded record shown in Trenberth and Hurrell1994). The Rossby wave transit time is approximately
3.5–6 yr assuming a group velocity of 3–5 cm s21 appropriate for the latitude band 308–408N (based on recent satellite altimetry measurements; see Chelton and
Schlax 1996) and a zonal extent for the wind forcingof 608 longitude from the western boundary (recall Fig.

FIG. 10. (a) Latitudinal structure of the total geostrophic zonal flow
averaged between the surface and 1 km (m s21 ) in the far western
Pacific (1428–1508E) for 1970–80 (solid curve) and 1982–90 (dashedcurve). (b) As in (a) but for the total zonal wind stress (N m22 )averaged across the Pacific (1508E–1408W) for 1968–76 (solid curve)
and 1977–88 (dashed curve).

4). Thus, the duration of the wind stress curl forcing is
at least twice the transit time of the long baroclinic
Rossby wave, indicating that Sverdrup balance may be
expected to be achieved over the decadal time averages
considered here. We note that the Rossby wave speeds
derived from recent satellite altimetry measurements (cf.
Chelton and Schlax 1996) are nearly twice as fast as

JUNE 1999

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DESER ET AL.

predicted by linear theory, but can be reconciled withmore modern theoretical treatments (Killworth et al.
1997; Qiu et al. 1997).
The results indicate a 4–5-yr delay between the decadal-scale variation in thermocline depth near the western boundary and the basin-wide wind stress curl pattern(recall Figs. 1 and 2). The observed delay is roughly
consistent with that expected from the Rossby wave
adjustment process as discussed above. It is also worth
noting that the observed vertical structure of the decadal-scale thermal change (specifically the amplitude
maximum within the main thermocline seen in Fig. 8)
is consistent with the signature of a first-mode baroclinic
Rossby wave that has its maximum vertical velocity inthe main thermocline. However, ocean modeling experiments are needed to substantiate whether the observed delay and vertical structure are dynamically consistent with the Rossby wave adjustment process.
Other thermal signals extending down into the main
pycnocline have been reported in association with themid-1970s climate transition. For example, Watanabe
and Mizuno (1994), Deser et al. (1996), and Schneider
et al. (1999) identified an apparent subducted thermal
anomaly in the main thermocline of the central North
Pacific during the 1980s and traced its movement southwestward and downward following the mean subtropicalgyre circulation, in accord with the theory of the ventilated thermocline. We view the diabatically induced
subduction anomaly as distinct from the thermal signal
associated with the dynamical spinup of the gyre circulation documented here, due to differences in both
location and timing.
Zhang and Levitus (1997) present some evidence for
decadal-scale anticyclonic rotation of subsurface thermal anomalies around the subtropical gyre. This advective signal is most apparent at 250-m depth, withlittle systematic movement indicated at the sea surface
or at 400 m (Zhang and Levitus). The lack of movement
at 400-m depth is consistent with the findings of Miller
et al. (1998) and our own analyses (not shown), whichindicate that the westward-intensified pattern of thermal
anomalies at 400 m (e.g., Fig. 3) is, to first order, astationary feature that grows and decays in place on thedecadal timescale. The clockwise movement at 250 m
may be associated with anomalous subduction as described above, particularly since this depth lies within
the ventilated thermocline zone. Further work is needed
to understand the different spatial and temporal characteristics of decadal-scale thermal variations at 250and 400-m depth.
It is interesting to speculate about the possible future
effects of the southward-intensified subtropical gyre circulation during the 1980s. According to the paradigm
put forth by Latif and Barnett (1994), increased poleward heat transport by a stronger subtropical gyre maylead to a reduction of the meridional sea surface temperature gradient, which in turn may diminish thestrength of the zonal wind stress, resulting in a weak-ening of the gyre over the next decade or two. Further
observational and modeling experiments are needed toevaluate the realism of this scenario.
5. Conclusions
The purpose of this study was to test the hypothesisthat the westward-intensified pattern of decadal-scaletemperature change at 400-m depth in the North Pacificresulted from the dynamical adjustment of the oceanic
circulation to changes in wind stress curl according to
Sverdrup theory. Using 6-hourly wind stress curl datafrom the NCEP–NCAR reanalyses and upper-ocean
thermal profiles from the WOA94, we demonstrated thatthe spatial patterns of the decadal-scale change in Sverdrup transport streamfunction and heat content in theupper 450 m are very similar. We extended the thermal
data down to 1-km depth to estimate quantitatively the
decadal-scale changes in geostrophic transport along theKuroshio extension and found good agreement with thewind stress curl induced changes in transport (11.6 vs11.5–13.9 Sverdrups). It remains to be seen whether theobserved 4–5-yr delay of the thermocline anomalies in
the western portion of the subtropical gyre relative tothe decadal-scale wind stress curl forcing is physicallyconsistent with the Rossby wave adjustment process.
Acknowledgments. We wish to thank Drs. Arthur
Miller, Ping Chang, William Large, and Peter Gent for
helpful discussions during the course of this work. The
graphics were produced with the GrADS software package developed at the University of Maryland.
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