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Subduction of Decadal North Pacific Temperature Anomalies: Observations andDynamics
NIKLAS SCHNEIDER AND ARTHUR J. MILLER
Scripps Institution of Oceanography, La Jolla, California

MICHAEL A. ALEXANDERCIRES, University of Colorado, Boulder, Colorado

CLARA DESER
National Center for Atmospheric Research, Boulder, Colorado(Manuscript received 9 January 1998, in final form 6 July 1998)
ABSTRACT
Observations of oceanic temperature in the upper 400 m reveal decadal signals that propagate in the thermocline
along lines of constant potential vorticity from the ventilation region in the central North Pacific to approximately
188N in the western Pacific. The propagation path and speed are well described by the geostrophic mean
circulation and by a model of the ventilated thermocline. The approximate southward speed of the thermal signalof 7 mm s21 yields a transit time of approximately eight years. The thermal anomalies appear to be forced byperturbations of the mixed layer heat budget in the subduction region of the central North Pacific east of thedate line. A warm pulse was generated in the central North Pacific by a series of mild winters from 1973 to1976 and reached 188N around 1982. After 1978 a succession of colder winters initiated a cold anomaly in thecentral North Pacific that propagated along a similar path and with a similar speed as the warm anomaly, thenarrived in the western tropical Pacific at 188N around 1991. Tropical Ekman pumping, rather than further
propagation of the midlatitude signal, caused the subsequent spread into the equatorial western Pacific and anincrease in amplitude. Historical data show that anomalous sea surface temperature in the equatorial central
Pacific is correlated with tropical Ekman pumping while the correlation with thermal anomalies in the North
Pacific eight years earlier is not significant. These results indicate no significant coupling in the Pacific ofNorthern Hemisphere midlatitudes and the equatorial region via advection of thermal anomalies along the oceanic
thermocline.

1. Introduction
Several theories have recently been advanced to attempt to explain decadal changes in the North Pacific
Ocean. Most of these rely on the long timescale set bybaroclinic Rossby wave propagation and its effect on
the baroclinic adjustment time of the subtropical oceanic
circulation (e.g., Latif and Barnett 1994, 1996; Robertson 1996; Jin 1997; Frankignoul et al. 1997). An intriguing alternative hypothesis is that the tropical andmidlatitude ocean are linked by subducted thermal
anomalies propagating equatorward along the oceanic
thermocline (Gu and Philander 1997; Inui 1997) or inwave guides along the western boundary (Lysne et al.

Corresponding author address: Dr. Niklas Schneider, Climate Research Division, Scripps Institution of Oceanography 0224, University of California, San Diego, La Jolla, CA 92093-0224.
E-mail: nschneider@ucsd.edu

q 1999 American Meteorological Society

1997; Venzke 1999). If the ocean can slowly send asignal to the tropical thermocline, which eventually alters tropical sea surface temperature (SST), the atmosphere could rapidly send the signal back to the midlatitudes via teleconnections (Bjerknes 1966; Alexander
1992; Philander 1990) and establish a decadal variation.
Because of the short observational record, however, it
is difficult to verify such a theory.
Do subsurface ocean observations support the mechanism of subduction as a viable candidate for tropical–
extratropical interaction on decadal timescales? Deseret al. (1996, DAT96 hereinafter) presented the first evidence for a subducting temperature anomaly in the
North Pacific. With origins north of Hawaii, the peakanomaly appears to move downward (approximately to
200 m) and southward (roughly 88 lat) toward 258N overa time period 10–15 yr. A great deal of study of theNorth Pacific thermocline followed that work (e.g.,
Zhang and Levitus 1997; Inui and Hanawa 1997; Yasuda

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and Hanawa 1997; Miller et al. 1998; Tourre et al. 1999)
and supports the idea of downward and southward moving temperature anomalies in the central North Pacific.
But it has not been clearly established if these anomalies
reach the equatorial thermocline and what physics controls the excitation, migration, and decay of these anomalies.
Here we address several questions associated with thesubduction of temperature anomalies in the North Pacific by studying the available observations (White
1995) in a new way and using simple models to help
interpret them. What is the horizontal, vertical, and temporal character of the subducting cold anomalies foundby DAT96 in the 1980s? Is that event unique or are
there other subducting anomalies in the historical record? Do anomalies move far enough south to interact
with the equatorial thermocline? Our novel approach is
to analyze the 3D evolution of temperature anomalies
in the North Pacific in terms of isotherm depths andpotential vorticity (PV). This leads to a clear delineationof decadal anomalies in regions associated with recirculation, ventilation, and shadow zones of the North
Pacific (e.g., Luyten et al. 1983, hereinafter LPS83; Talley 1985).
Is the downward and southward migration of thesetemperature anomalies consistent with the physics of
subduction? Anomalous subduction has often been considered as salinity-compensated temperature anomalies
being advected by the mean velocity field along mean
isopycnals. Liu (1999a,b) has proposed that highermode baroclinic waves may also be advected along isopycnals. These anomalies would not be salinity compensated but would have an anomalous pressure andvelocity signature. Local forcing by Ekman pumping
could also explain temperature anomalies in the midocean if wind stress curl has decadal variations (e.g.,Trenberth 1991; Miller et al. 1998). Baroclinic wave
propagation on interannual timescales may obscure thesubducting temperature anomalies (Miller et al. 1997),
and may explain why the DAT96 anomalies were foundto increase with time and depth. We study the motionof the subducting anomalies by invoking the ventilated
thermocline theory of LPS83 to clarify the roles of subduction and wind forcing. Do the anomalies follow apath along constant PV according to this theoreticalframework?
What processes drive these subducting anomalies?
The past several decades have seen evidence for regime
shifts in both the midlatitudes (Trenberth 1990; Graham
1994; Miller et al. 1994a; Mantua et al. 1997) and theTropics (Kleeman et al. 1996; Latif et al. 1997; Goddard
and Graham 1997). Are these regime changes associatedwith the surface forcing in the North Pacific that in turndrives the subducting anomalies? Is the surface forcing
primarily diabatic, driving the heat equation, or adiabatic, affecting the vorticity equation (cf. Liu 1998,
1999; Inui and Hanawa 1997; Venzke 1999)? We first
identify the source region for the subducting anomalies

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and then study the local atmospheric forcing fields[Comprehensive Ocean–Atmosphere Data Set(COADS)] to identify the dominant processes that initiate the subducting anomalies.
The next section introduces the datasets and describes
the methods used to determine the depth of isotherms.
Section 3 discusses the observed mean state and its
comparison with LPS83; section 4 describes decadal
anomalies of isotherm depths and compares propagation
paths with the distribution of potential vorticity. The
relative roles of propagation and local wind stress forcing are discussed in section 5 followed by the conclusionand discussion of results in terms of proposed dynamics
of decadal climate signals.
2. Dataa. Upper-ocean temperature
All available BT, XBT, and CTD stations were collected and optimally interpolated by White (1995) to athree-dimensional grid of 28 lat by 58 long, 11 standard
depth levels between the surface and 400 m, and monthly anomalies from 1955 to 1996. Decorrelation scalesof the optimal interpolation are 108 long, 58 lat, and 90
days. In the Pacific data density is sufficient to describe
the three-dimensional temperature structure north of
208S and we use data from 1970 to 1996.
Using the same ocean observations but a different
interpolation scheme, DAT96 identified southward
propagating cold anomalies in the thermocline during
the 1980s. We computed annually averaged anomalies
for the same vertical sections as DAT96 and applied a5-yr filter with linearly varying weights. The White dataset captures the same cold anomaly starting at thesurface after 1977 and propagating southward in later
years (Fig. 1). As found by DAT96, the cold anomalies
appear to increase in amplitude with time. The primaryreason for this is that cold (upwelling) thermocline
anomalies associated with the 1984–85 and 1987–88 La
Niña events (Miller et al. 1997) propagate from the coastof North America westward into the averaging area near
158–258N roughly 2 years after their excitation in thetropical Pacific. As described by Miller et al. (1997) andMiller et al. (1998), extended empirical orthogonal function analysis can separate these interannual propagating
signals from the decadal-scale subducting signal. Thus,
the low-pass filtered anomalies described below and by
DAT96 are real and not aliased higher-frequency signals.
The data also show warm anomalies that first appear
close to the surface in the early 1970s (Fig. 1) andmigrate southward somewhat faster than the cold anomaly. The southward travel of both the cold and warm
anomalies follows the mean isotherms 128–188C that
also straddle the change of SST associated with the
1976–77 climate shift (Fig. 2). This suggests a description of vertical and horizontal propagation of thermal

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FIG. 1. Filtered decadal temperature anomalies from the White (1995) data analysis in the central
North Pacific between 1708 and 1458W as a function of latitude and depth and centered on theindicated year. A five-point triangular filter was applied to annual averaged temperature data. The
temperature scales in kelvin is given by the gray scale on the left and contour interval is 0.2 K.
Superimposed are the 128 and 188C isotherms of the time-averaged temperature field (dashedlines).

anomalies by mapping undulations of the depth of the
128–188C isotherms. Such a description has the added
benefit that it lends itself to comparison with the dynamical framework of layers that have been used to
explain and explore the thermocline and its variations(Luyten et al. 1983; Pedlosky 1996).
Isotherm depths were calculated using the following
procedure. The three-dimensional temperature was reconstructed from the monthly climatology and anomalies.
At every horizontal location vertical temperature profiles
were sorted to remove temperature inversions that occur
quite frequently and were of much smaller magnitude
than differences of and shallower than the depth of isotherms of interest. The profiles were then made strictly

monotonic with depth by the introduction of an insignificantly small (1024 K m21 ) stratification to homogenous levels. Isotherm depths were then obtained by linear
interpolation of these vertical temperature profiles. Overall, this procedure is equivalent to determining the mean
depth of a particular isotherm in a multivalued profile.
Layer thickness is then determined as the difference in
depth of the 128C and 188C isotherms.
A monthly mean climatology of the isotherm depths
was then computed and anomalies were formed withrespect to it. To extract low-frequency signals, annual
averaged anomalies were smoothed with a five-point
filter with linearly decreasing weights (a normalized 1–
2–3–2–1 filter) and a half power point at 6.7 yr.

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FIG. 2. Surface temperature change associated with the 1976–77
climate change in the North Pacific. Shown as shaded contours is thedifference between 1970 to 1974 and 1978 to 1982 based on a reconstruction using the leading EOF of surface temperature anomalies.
Shading is explained by the bar on the right and contour intervals is
0.3 K. Superimposed as black lines with white borders are mean
wintertime (Jan, Feb, Mar) surface isotherms of 98, 128, 158, 188, and218C. For later reference the 128 and 188C isotherm are emphasized
by a heavy contour.

b. SST and atmospheric forcing
In addition to subsurface temperatures, anomalies of
surface temperature compiled by the U.K. Hadley Centre were used. This dataset covers the time period from1930 to the present (Rayner et al. 1996). Also the windstress product of Da Silva et al. (1994) based on COADS
data and covering 1945–92 and surface heat turbulent
fluxes computed by Cayan (1992a) from COADS are
employed. The latter fluxes were used by Miller et al.
(1994a,b) to model the oceanic response to the 1976–
77 climate shift and are discussed in detail by them.
3. Mean state depths and potential vorticity
a. Layer depth
Mean depths of the 128C and 188C isotherms display
a bowl shape in the subtropical gyre and in the transition
to the equatorial current system near the edge of theNorth Equatorial Countercurrent (Fig. 3). The 128C isotherms outcrop along 388N with a slight northward slope
toward the eastern half of the basin (Fig. 2). The 188C
isotherm outcrops between 358 and 308N with a southward slope toward the east. South of these outcrop lines
the depth of isotherms increases sharply in the west, as
expected from the strong zonal currents in the Kuroshio
and Kuroshio extension region (Fig. 3), and reach thegreatest depths of 400 m at 308N for the 128C isothermand 150 to 200 m at about 208N for the 188C isotherm.
The mean depth of the layer defined by the 128 and
188C isotherms exceeds 300 m in the western subtropics.
In the eastern half of the basin the layer shoals andforms a region of small depth changes that connect waters off North America with the Tropics. Farther south

FIG. 3. (Top panel) Average depth of the layer defined by the 128
and 188C isotherm derived from the White (1995) analysis of temperature in the upper ocean. Depth is given in meters and depictedby gray shading explained by the bar on the right. Depth varies fromless than 50 m in the white areas to more than 300 m in darkest areas
with a contour interval of 50 m. Superimposed white lines are isolinesof potential vorticity of this layer in 10210 m21 s21 . (Bottom panel)
The mean depth of layer 2 of a 3-layer model (LPS) of the ventilated
thermocline shown in the same gray scale as the top panel. Superimposed are the model’s potential vorticity in layer 2. The model(Luyten et al. 1983; Talley 1985) was forced by time averaged Ekman
pumping derived from the Da Silva et al. (1994) wind stress. Notethat the model captures depth and potential vorticity very well in theventilated region, but fails to capture the structure of shadow region
in the east.

a ridge in the depth of the layer crosses the Pacific andseparates the North Equatorial Current on its northernslopes from the North Equatorial Countercurrent on its
equatorward slopes. On the ridge, the 128C isotherms
are shallower than 150 m at 1408W and reach close to
200 m in the western part of the basin. The 188C isotherm has a similar east–west gradient and varies between 150 m in the west and 50 m in the east. On theequator, the depth of the isotherms show the east–west
slope with a deep thermocline in the western Pacific anda shallow thermocline in the upwelling regions of theeastern equatorial Pacific.
b. Potential vorticity
Potential vorticity (PV) in the limit of small Rossby
number, in a layered fluid of small aspect ratio, is obtained from
PV 5

Dr f
,
r 0 DH(1)

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where f is the Coriolis frequency, DH is the layer thickness estimated as the difference of depth of the 128C
and 188C isotherms, r o is the mean density, and Dr isthe density difference between these two isotherms (e.g.,
Talley 1988; Pedlosky 1996). Salinity has not been included in the estimation of potential vorticity. This isvalid if either temperature dominates density over salinity or if the temperature–salinity relationship is a constant so that horizontal gradients of temperature anddensity are parallel. Experiments with data of Levitus(1982) show that, in the North Pacific away from theeastern boundary, potential vorticity based on isotherms
and isopycnals correspond closely.
Lines of constant potential vorticity (Fig. 3) emanatefrom the outcrop region of the layer in the north Pacificand form a closed region in the west, reminiscent of theRhines and Young (1982) homogeneous pool of potential vorticity. From the outcrop regions east of the dateline, lines of constant PV extend southward, swing toward the west, and identify the ventilated region(LPS83). Potential vorticity lines that emanate from theeastern boundary swing toward the west and delineate
the shadow zone to the east of a potential vorticity maximum. South of 108N the potential vorticity barrier described by McCreary and Lu (1994) associated with theNorth Equatorial Countercurrent is seen, and toward theequator potential vorticity decreases due to the decreaseof f. However, for the very low latitudes, planetaryvorticity is no longer dominant over relative vorticityof the mean flow (the Rossby number is no longer small
compared to one), and Eq. (1) is no longer valid. These
results are in broad agreement with results with Talley(1988) even though a detailed comparison is difficult
because of the coarse resolution of isotherms.
c. Application of the ventilated thermocline
To establish our dynamical framework and for later
application to observations of the anomalous layer
depths, we compare observed mean layer depth and potential vorticity with the predictions of the ventilated
thermocline (LPS83; Talley 1985). This model predicts
the time-independent thickness of three active layers
overlying a deep motionless ocean based on conservation of potential vorticity and the Sverdrup balance.
It has previously been successfully applied to the timemean circulation of the North Pacific (Talley 1985). Parameters required are average densities of the layers,
reduced gravities between layers, latitudes of outcrop
lines, and the depth of the lowest active layer at theeastern boundary. Densities are estimated from the Levitus (1982) climatology for waters above 1500 m. The
density of layer 1 is defined as the average density ofwaters warmer than 188C, layer 2 by temperatures between 128 and 188C, and layer 3 by temperatures between 128 and 68C. Reduced gravities are determinedfrom the density differences between the layers. The
depth of the lowest active layer at the eastern boundary

VOLUME 29

TABLE 1. Parameters for the three-layer thermocline model (Luyten
et al. 1983; Talley 1985). The outcrop latitudes were chosen fromthe (Jan–Mar) outcrops of the respective isotherms in the North Pacific from the White (1995) temperature dataset. Reduced gravities
were obtained from Levitus by averaging densities of the North Pacific above 1500-m depth that correspond to temperatures within thelayer bins. The undisturbed depth of layer three of 600 m was determined from the Levitus data as the depth of the 68C isotherm in
the eastern subtropical Pacific.

Layer

Outcroplatitude

Reducedgravity(m s22)

Minimum
temperature
(8C)

1
2
3

338N
418N
—

15·10239·1023
7·102318
12
6

was determined from Levitus (1982) as the depth of the
68C isotherm off California. In the solution of the ventilated thermocline of LPS83 and Talley (1985) the outcrop lines are zonal. Their latitudes are determined here
from the winter surface conditions of the North Pacificof the climatological temperature data based on White(1995). The resulting parameters are listed in Table 1.
The solution of the mean ventilated thermocline simulates the mean depth of the layer defined by the 128
and 188C isotherms very well in the ventilation region
and western North Pacific, including the position, southwestward swing, and magnitude of the potential vorticity isopleths (Fig. 3). Observed average layer depth andpotential vorticity isopleths to the south of the ventilation region are not well captured by the model. Inclusion of the eastern boundary ventilation of layer 2
would generate a ridge of high potential vorticity at theeastern edge of the ventilation region and deepen theequatorial layers (Pedlosky 1983) in agreement with observations. However, this arduous extension of the analytic theory suggested by Pedlosky (1983) was not attempted since the region of ventilation from the outcrops
in the north Pacific, the main focus of this work, is well
reproduced.
4. Decadal anomalies
Anomalies in the depth of the layer defined by the
128 to 188C isotherms are estimated as an average ofthe depth anomalies of the 128, 158, and 188C isotherms.
The signal described below is consistently found in eachof the isotherms and the average is chosen to increase
the clarity of the presentation.
Root-mean-square values of low-frequency depth
anomalies (derived by application of the five-point filter
to annually averaged anomalies) from 1970 to 1996
(Fig. 4, top panel) show that variability occurs in three
regions delineated by the mean potential vorticity structure. Rms values are large in the Kuroshio Extension
where potential vorticity has small gradients, in the ventilation region of the thermocline defined by values ofmean potential vorticity between 3.5 and 6.5 (310210

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FIG. 4. Rms values of decadal depth anomalies of the 128–188C
layer (in m) with a contour interval of 2 m (upper panel) and of
anomalies of layer thickness normalized by the mean thickness inpercent with a contour interval of 2% (lower panel). The latter quantity is a measure for the anomalous perturbation vorticity. Values of
decadal depth anomalies larger than 3 m (top panel) and values of
the relative thickness change larger than 2% (bottom pane) are shaded.
Black lines with white borders show the mean potential vorticity of
the layer, and the white and black dashed line in the upper panel
surrounds a region over which depth anomalies are zonally averaged
in Fig. 6. It is designed to capture the large variance of the depth
anomalies. Poleward of 158N it corresponds to the ventilation region
defined by values of the mean potential vorticity between 3.5 and
6.5 (310210 m21 s21 ).

m21 s21 ) extending from the central North Pacific to thetropical western Pacific, and in the shadow zones of theeastern North Pacific where potential vorticity lines
originate at the coast of North America. The elevatedrms values in the shadow zone are predicted by Liu(1996) as a Rossby wave response to changes of the
Ekman pumping, while Liu and Pedlosky (1994) suggest that variability in the ventilation region results primarily from changes of the surface temperatures in the
subduction region. The variability in the Kuroshio Extension is consistent with the spinup of the gyre in response to decadal changes of the wind stress curl (Miller
et al. 1998; Deser et al. 1999).
The depth anomalies in the ventilation region propagate southwestward from the outcrops in the central
north Pacific to the western tropical Pacific (Fig. 5). In
1972 a positive depth anomaly, corresponding to a warm
signal, existed north of Hawaii in the outcrop region of1061

the 3.5 to 6.5 (310210 m21 s21 ) potential vorticity contours. This anomaly strengthened in 1974 to achieve
values in excess of 15 m in 1976, moved from 1974 to1980 in a southwestward direction along lines of constant potential vorticity, and disappeared at 208N in
1982. In the western Pacific south of 208N a warm
anomaly with depth anomalies of up to 9 m existed from1974 to 1978 that cannot be traced to the subduction
region in the central North Pacific.
Starting in 1980, a negative depth anomaly, corresponding to cold conditions, appeared in the subduction
region in the central North Pacific. This anomaly
strengthened in 1982 and possibly 1986 to values of 9m and moved along lines of constant potential vorticity.
From 1984 to 1990 the signal propagated toward thewestern Pacific in a southwestward direction along lines
of constant potential vorticity and reached the western
Pacific south of 208N in 1990. From 1990 to 1994 it
appears that the signal spreads to the equatorial western
Pacific and increases in magnitude to as large as 15 m
while the zonal tilt of the equatorial thermocline is reduced by a deepening in the eastern equatorial Pacific.
However, we will show in the following that the large
anomalies in the western Pacific are a response to tropical wind forcing and not due to propagation of themidlatitude signal to the equator.
The southward propagation and travel time can beseen most clearly by averaging the depth anomalies zonally in the ventilation region (shown by the dashed line
in Fig. 4) that, poleward of 158N, correspond to the area
between the mean potential vorticity lines of 3.5 and
6.5 (310210 m21 s21 ). The southward propagation has
a meridional speed of approximately 7 3 1023 m s21
and a travel time from the subduction region north of
Hawaii at 358N to the western Pacific of approximately
8 yr (Fig. 6). The propagation compares very well withadvection by geostrophic currents between the 128 and
188C isotherms and by the circulation predicted by themodel of the ventilated thermocline. Meridional geostrophic velocities from the geostrophic current relative
to 1500 m from Levitus (1982) data and from the solution of the ventilated thermocline (Fig. 3) were zonallyaveraged over the ventilation region (Fig. 4) and integrated in time to yield the paths in Fig. 6.
In contrast to anomalies of layer depth, anomalies ofpotential vorticity do not show any coherent signals.
Rms values of the ratio of layer thickness anomaly andthe time-averaged layer thickness, a measure for the
perturbation potential vorticity [see Eq. (1)], are lessthan 5% (Fig. 4) and do not show any coherent propagation (not shown). Despite the small values, the rms
of perturbation PV are elevated along the westernboundary, ventilation region, and in the eastern shadowzone (Fig. 4). The small values of the perturbation PV
suggest the possibility of linearization of the dynamics
around the mean state and indicates that the perturbation
forcing is small compared to the mean forcing. Also,
the usual ventilated thermocline analysis approach of

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FIG. 5. Decadal anomalies in the North Pacific of the depth of the 128–188C layer for years 1972 to 1994. Annual
averaged anomalies were smoothed with a filter with linearly varying weights over a 5-yr interval. Dark shaded areas
mark positive anomalies larger than 3 m, corresponding to anomalous large depth of the layer and warm conditions,light gray shading denotes negative depth anomalies in excess of 23 m and correspond to cold conditions (see grayscale on the left). Contour interval is 3 m, and the zero contours are not shown for clarity. Overlaid as dashed lines are
the 4 and 6 3 10210 m21 s21 contours of the mean potential vorticity (e.g., Fig. 3).

estimating PV anomalies on constant density surfaces
is difficult due to the small amplitude of PV anomalies.
In summary, the data show a clear propagation ofthermal signals along isopleths of mean potential vorticity from the subduction region in the central North
Pacific north of 338N to the tropical western Pacific to
188N, while the perturbations of potential vorticity aresmall. The meridional speed of the signals (;7 mm s21 )closely matches advection by geostrophic currents andthe solution of the ventilated thermocline driven by thetime-averaged winds. These findings strongly suggestthat poleward of 188N the perturbations are carried by
the mean circulation, which is governed by dynamics
of the ventilated thermocline. Equatorward of 188N thecorrespondence between the propagation of thermal
anomalies and the predictions by the ventilated ther-

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1063

a. Surface fluxes in outcrop region
Consider the equation for temperature T in oceanic
surface mixed layer of depth H,]Qw
T 5 2u · =T 1
2 e DT,]t
Hr 0 c p
H

FIG. 6. Zonal average of decadal anomalies of the depth of the 128
to 188C layer in meters as a function of year and latitude. Average
was performed in the region denoted by a black and white dashedline in Fig. 4. The black line on white background denotes the pathof a particle advected by the zonally averaged meridional speed in
the same region but predicted from the model of the ventilated thermocline forced by average of observed mean winds from 1969 to
1993. The dashed line on white background shows the path of sucha particle advected with the geostrophic flow relative to 1500 m basedon the Levitus (1982) data.

mocline fails and suggests that local forcing is important.
5. Surface forcing
The propagation of thermal anomalies in the thermocline begs the question of how these are forced in
the outcrop region and whether their propagation andgrowth in the tropical Pacific could be influenced bytropical Ekman pumping. To investigate these issues,
surface forcing in the source region in the central North
Pacific and the Ekman pumping affecting the areas visited by the propagating anomalies are investigated using
COADS-based surface forcing from Cayan (1992a) andfrom Da Silva et al. (1994). In the following, we define
the subduction source region as 288–388N, 1558–1258W.

(2)

where the left-hand side describes the time (t) rate ofchange of temperature. The right-hand side represents
advection by mixed layer velocity u, turbulent fluxes
into the mixed layer due to the surface heat flux Q,
normalized by the ocean density r 0 and ocean’s specificheat c p , and due to entrainment velocity w e importingwater into the mixed layer across its base with temperature jump DT. In midlatitudes, variations of the surface
heat flux are generally the largest contributor to SST
anomalies in winter (e.g., Frankignoul 1985; Cayan
1992b), the time when properties of subducted waters
are set (Cushman-Roisin 1987). Other important processes affecting SST anomalies are variations in mixing
and advection by ocean currents. These two processes
are predominantly controlled on seasonal timescales byanomalous wind speed and Ekman currents, respectively(e.g., Miller et al. 1994b). We evaluate the magnitudeof these three forcing effects on (2) in the following
way. Assuming the mixed layer to be of constant depth(H 5 100 m) in the winter months of January–March,
the heat flux term was approximated as Q9(r 0c pH)21 ,the mixing term as w9e DTH 21 , and the Ekman advection
term as u9e · =T M . Here, Q9 is the January to March(JFM) seasonal mean total surface heat flux anomaly(Cayan 1990), r 0 5 1025 kg m23 , c p 5 4180 J kg21
K21 , u9e is the Ekman velocity anomaly averaged overthe mixed layer and T M is the mean SST computed fromCOADS. Ekman pumping is evaluated as the divergenceof the Ekman currents (wind stress curl plus the beta
term: Gill 1982) from a blend of the COADS wind stresswith FSU wind stress (Miller et al. 1994b). Anomalous
entrainment heat flux was determined from the turbulent
kinetic energy balance between wind stirring, proportional to wind friction velocity cubed (u e3 , estimated according to Miller et al. 1994b) and the increase of potential energy due to the entrainment heat flux.
The time series of these anomalous winter SST heatbudget terms reveal intervals of strong or persistentanomalous forcing in the subduction source region (Fig.
7). All three terms in the heat budget act to warm thesource region in the years 1973–76 when the warm
anomaly of layer depth was formed (Fig. 7). After 1976,
a series of anomalous severe winters, interrupted by
years of near normal or warm conditions, coincide withthe appearance of a cold depth anomaly in the source
region. The total cooling during this period is dominatedby the anomalous severe winter associated with the
1982–83 El Niño such that the total forcing during thewarm periods and the winter of 1982/83, normalized toaccount for the different duration, are of equal magnitude but opposite sign. The anomalous wintertime cool-

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FIG. 7. (Top panel) Time series of anomalous heat budget terms in 1027 K s21 . Solid line denotes
anomalies of surface heat flux, dashed line denotes entrainment, and dotted lines denotes Ekman
advection. (Bottom panel) Time series of divergence of Ekman pumping in 10 27 m s21 in theventilation region of the 128 to 188C layer. Solid line is winter (Jan–Mar) and dash-dotted isannual average.

ings are dominated by the entrainment heat flux, thatis, by changes of the wind friction velocity, consistent
with changes of the zonal wind speed reported by
DAT96 and the changes of Ekman advection. In addition, the cooling episodes in the winters of 1980/81,
1982/83, and 1986/87 are associated with the occurrenceof El Niño and suggest that the anomalous atmosphericconditions are communicated to the North Pacific via
atmospheric teleconnections (Alexander 1992; Lau
1997). The anomalous Ekman pumping on the other
hand has large year-to-year changes, but no signals (foreither annual or winter anomalies) that are coincident

with the occurrence of anomalies of the 128–188C layer
depth (Fig. 7).
Comparisons of the spatial patterns of these heating
terms averaged over 1974–76 and for the winter of
1982/83 show that the subduction source region is close
to extrema of heat flux, of flux of turbulent kinetic energy, and of zonal wind stress. In contrast, the source
region straddles the nodal line of anomalous Ekman
pumping (Fig. 8). Moreover, south of the source region
the Ekman pumping acts in the opposite sense of theobserved anomalies, for example, downwelling during
the cold subducting phase. Thus it appears that oceanic

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FIG. 8. Anomalies of surface forcing for winters of 1973–76 (left column) and winter 1982/83 (right column) in the
North Pacific derived from the COADS and FSU databases. Shown are anomalies of fluxes of heat in W m 22 (top row),
flux of turbulent kinetic energy measured by wind speed cubed in m 3 s23 (second row), heat advection by Ekman currents
in K s21 (third row), and Ekman pumping in 1027 m s21 (bottom row). The solid box shows the region where the 128
to 188C layer ventilates.

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1978 the winter position of these isotherms was locatedhalf a degree farther northward than average, and thenshifted southward during the 1980s. In the early 1990s
the isotherms moved again northward by over one degree of latitude from its position in the 1980s.
c. Sverdrup transport

FIG. 9. Time series of outcrop latitude of 128, 158, and 188C winterisotherms (Jan, Feb, Mar) averaged from 1708 to 1408W. Data is basedon surface XBT data. Legends denote average latitudes of isotherms
at the surface during winter.

thermal anomalies that originate in the source region
are generated by diabatic surface forcing, that is, surface
heat flux and entrainment and Ekman advection.
It is noteworthy that nature provides two distinct scenarios for the generation of the subducting anomalies:
the warm anomaly is a response to persistent anomalous
heating over several successive winters, while the cool
anomaly results from a series of anomalously severe
winters interrupted by near normal or warm conditions.
It is interesting to speculate that the action of Stommel’s
demon (Pedlosky 1996; Williams et al. 1995; Stommel1979) requires that forcing of a warm anomaly has tolast several years to allow anomalies to subduct andescape the influence of a subsequent harsher winter. In
contrast, an anomalously cold winter with deep mixed
layers can generate subducted anomalies that are insulated from subsequent milder winters and are therefore
likely to persist and propagate into the main thermocline. Lastly, the unique winter of 1976/77 associatedwith a climate shift of the midlatitude North Pacific does
not stand out as a key forcing period of the observed
subducting anomalies. It serves, however, as a transition
point for modulating the effects of interannual variations
in forcing, allowing a preponderance of warming eventsin antecedent years versus cooling events in subsequent
years.
b. Variation of outcrop latitude
Diabatic forcing and Ekman advection in the source
region affects the SST and leads to displacements of theoutcrop latitudes of the 128 to 188C layer. Since themean meridional SST gradient is large in the source
region (Fig. 2), surface temperature anomalies of 18C
are equivalent to variations of the outcrop latitudes ofless than a degree of latitude (Fig. 9). From 1970 to

The observations presented suggest a propagation of
decadal signals from the mid-latitude North Pacific tothe subtropical western Pacific. However, the appearance of a warm anomaly at 108N in the western Pacificfrom 1974 to 1978 was apparently unconnected to propagation from the midlatitudes. Likewise, the strong
growth of the cold anomaly in the western Pacific after
1990 (Fig. 5) does not favor the advection mechanism
but instead suggests that vertical displacements of thelayer caused by local or remote wind forcing may be
important. To test this hypothesis the interface displacement associated with Sverdrup transport (Gill 1982) was
estimated from Da Silva et al. (1994) winds, annually
averaged from 1969 to 1993 and subjected to a 5-yr
filter with linearly varying weights.
Deviations of the thermocline depth expected fromthe Sverdrup transport alone do not show any relationship with the observed undulations of the 128 to 188C
layer depth north of 188N (cf. Fig. 6 with Fig. 10) andsuggest that observed thermal signals are indeed propagating and not locally forced. South of 208N the steadyoceanic response to wind stress forcing is similar to andprecedes oceanic observations by approximately one
year (Fig. 10). This lag is expected from the adjustment
time of the low-latitude ocean via Rossby waves andsuggests that the observed thermal signals are a response
to zonally integrated Ekman pumping.
d. Ocean response to Ekman pumping and variationof outcrop latitude
The results so far indicate that the anomalies in thethermocline are forced by the wind stress to the southof 188N, while the anomalies in the latitude of the outcrop line due to diabatic forcing and Ekman advection
might generate subsurface temperature anomalies in theventilation region between the source region and 188N.
To test this hypothesis, the ventilated thermocline modelof LPS83 and Talley (1985) was applied under the assumption of instantaneous adjustment to the time varying forcing. The model was forced by filtered Da Silva
et al. (1994) winds and outcrop line variations from1970 to 1992 and the subsurface depth anomalies fromthe time-average state were calculated. In all experiments the potential vorticity of the western pool in layer
3 (Talley 1985) was held constant to focus on the dynamics in the ventilation region.
Zonal averages over the ventilated region (Fig. 4)
show that the large anomalies south of 188N can belargely explained by anomalous Ekman pumping at

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therms of the simple model closely matches the observations.
6. Discussion and conclusions
We have presented an analysis of observed long-term
anomalies in the thermocline of the North Pacific. In
contrast to earlier studies, the description of thermal
anomalies by undulations of the depth of isotherms allowed a succinct tracing of the propagation of the thermal anomalies and lends itself to a comparison withpredictions of the ventilated thermocline model. Specifically, we focused on the layer bounded by the 128
and 188C isotherms since it outcrops in the central North
Pacific in a key region noted for large decadal changesof sea surface temperature. Results of the study are:

FIG. 10. Anomalies of thermocline displacement in meters due to
Sverdrup transport. Sverdrup transports were estimated from the Da
Silva et al (1994) wind stress data for each year from 1969 to 1993
and filtered with a 5-yr running mean with linearly varying weights.
Transports were converted to displacements by assuming a reduced
gravity of 1.6 3 1022 m 2 s21 . Anomalies from the time average were
zonally averaged in the ventilation region (Fig. 4) and results are
displayed as a function of time and latitude.

these latitudes (Fig. 11) as seen previously in the calculation of the Sverdrup transport (Fig. 10). However,wind-stress-induced signals north of the 188N are smaller and cannot explain the observations (Fig. 6). On theother hand, variations of the outcrop latitude (Fig. 9)
affect areas to the north of 188N. Changes there are inphase with the observed depth anomalies, suggesting
the importance of forcing by displacements of the outcrop lines. The preponderance of forcing of decadal
anomalies in the ventilated region of the ocean by diabatic forcing and Ekman advection as compared to Ekman pumping is consistent with theoretical predictionsof Liu and Pedlosky (1994) and Liu (1996).
Note that propagation of subducted anomalies cannot
be seen in Fig. 11 because it represents a final steadystate to outcrop line changes and not the time-dependent
adjustment process observed in nature. It is neverthelessremarkable and attests to the relevance and skill of themodel of the ventilated thermocline that the southward
penetration limit of the instantaneously adjusted iso-R Decadal variance of the anomalies of the layer depth
is large in three regions that are distinguished by theirpotential vorticity (PV) contours: The Kuroshio Extension region in the western North Pacific north of
a pool of nearly homogenous PV, the ventilation region from the central North Pacific east of the dateline extending to the tropical western Pacific, and the
shadow zone where isopleths of PV originate at theeastern boundary.
R In the early 1970s a warm decadal anomaly originatedin the central North Pacific and propagated toward thewestern subtropical Pacific. In the early 1980s a cold
anomaly was generated in the central North Pacificand propagated along a similar path to the western
subtropical Pacific. Upon arrival in the subtropics, thecold signal strengthened and appeared to spreadthroughout the equatorial western Pacific.
R The propagation occurred along isopleths of averagepotential vorticity. Its speed of 7 mm s21 is well described by the mean advection predicted by the ventilated thermocline forced by observed wind stress.
The travel time from the central North Pacific to thesubtropical western Pacific is approximately eight
years.
R Anomalies of potential vorticity are at most a few
percent of the mean potential vorticity and suggestthat dynamics are governed by a linear perturbationof the background state.
R North of 188N thermal anomalies are forced by diabatic processes and Ekman advection that shift theoutcrop lines of a layer by typically one degree of
latitude.
R Equatorward of 188N thermal anomalies are primarily
forced by the wind stress curl and do not originate
from the North Pacific.
In summary, decadal anomalies are forced by diabatic
processes and Ekman advection in the subduction regionof the central North Pacific are swept toward the subtropical western Pacific with the mean circulation andarrive there approximately eight years later. Anomalies
equatorward of 188N are forced by low-latitude wind

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FIG. 11. Depth of layer two of the ventilated thermocline (Luyten et al. 1983; Talley 1985)
forced by annually averaged Ekman pumping from 1970 to 1992 and changes of the outcrop
latitudes as given in Fig. 9 (left panel). Shown is the deviation from the mean state during theforcing period. Positive anomalies are shaded and the contours are drawn for 64, 8, 16, and 32
m. Center panel shows the results if only changes of the surface Ekman pumping are consideredwhile the outcrop latitudes remain at their climatological position. The left panel shows changesof depth of layer 2 if the outcrop latitudes are changed only, while the Ekman pumping remains
constant. In all of the calculations, the potential vorticity of layer 3 in the western ‘‘homogenous
pool’’ was held constant.

stress curl. This implies that propagation of thermal
anomalies from the midlatitudes cannot explain anomalous warm conditions at the surface and reduced zonal
slope of the equatorial thermocline (Fig. 5) observed in
the equatorial Pacific in the early 1990s. These results
are largely consistent with observational analyses reported by Tourre et al. (1999) and Zhang and Levitus(1997), but we have been able to demonstrate the physics of the subducted anomalies and their confinement
to the midlatitudes and subtropics.
These results are inconsistent with the hypothesis thatthe decadal timescale results from propagation of anomalies from the midlatitudes to the equatorial region (Guand Philander 1997) and the observational analysis of
Zhang et al. (1998). However, the development of Ekman pumping south of 188N just at the time when thecold anomaly from midlatitudes arrives could indicate
a possible, albeit tenuous, linkage. The equatorial response could be due to the midlatitude signal if the
upward shift of isotherms in the thermocline associatedwith the cold anomaly affects sea surface temperaturein the western tropical Pacific through perturbations of
the surface layer heat budget and if the resulting atmospheric response induces the Ekman pumping thatstrengthens the signal and communicates it to the equatorial region. Alternatively, the arrival of the subductedsignal in the western tropical Pacific could be com-municated to the equatorial region via coastal Kelvinwaves (Lysne et al. 1997) or the western boundary currents (McCreary and Lu 1994) and initiate positive feedbacks in the atmosphere there. In either case, decadal
anomalies of sea surface temperature in the midlatitude
source region should be correlated with decadal anomalies in the Tropics at a lag of 8 to 10 years. The correlation of decadal anomalies in the source region andin the central equatorial Pacific eight years later is 0.23,
not significant at the 95% confidence level (0.55 for acharitable estimate of ten degrees of freedom). The correlation is significantly smaller than the correlation of
SST anomalies in the central equatorial Pacific with the
Sverdrup streamfunction in the western Pacific of 20.73(Fig. 12). This suggests that local forcing in the equatorial region is the main driving force and largely independent of the arrival of thermal signals from thecentral north Pacific. Thus, the data do not support significant coupling of the midlatitudes and Tropics via
anomalous subduction from the central north Pacificalong the oceanic thermocline.
The theoretical study of McCreary and Lu (1994) andthe simulation of Rothstein et al. (1998) suggest thatthe location of origin of subducted particles in the North
Pacific determines if these are either recirculated in thesubtropical gyre or penetrate to the equatorial region via
low-latitude western boundary currents or via a mid-

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FIG. 12. Time series of winter surface temperature in the ventilation
region of the 128–188C layer (top panel), sea surface temperature
anomalies in the central equatorial Pacific (center panel), and anomalies of the transport stream function in the western tropical Pacific(bottom panel). The decadal signals are extracted by application of
a 5-yr filter with linearly varying weights after removal of the mean
trend. Correlations are highest between SST in the equatorial Pacificand stream function (20.73). In contrast, the correlation of 0.23 between SST anomalies in the central North Pacific and in the equatorial
Pacific at a lag of eight years (north Pacific leading) is not significant
at the 95% confidence level. This weak lag correlation between SST
in the north and the equatorial Pacific is inconsistent with the notionthat strong coupling between the midlatitude North Pacific and equatorial regions by way of the advection in the oceanic thermocline.

ocean pathway. Thus thermal anomalies that originate
to the east or south of the central north Pacific could
communicate the decadal signal between the midlatitudes and the equator. However, decadal anomalies of
the 128–188C layer that originate east of 1358W do notshow any coherent southward propagation (Fig. 5). Similarly, depth anomalies of isotherms that are warmer andoutcrop on the southern flanks of the decadal midlatitudechange (Fig. 2) do not display coherent propagation thatis independent of those discussed (not shown). The dataalso provide no evidence that the thermal anomalies
return to the midlatitudes via the western boundary current. This, however, might be due to the optimal interpolation used in preparation of the dateset (White 1995)
that discriminates against small spatial scales such as
might be expected from changes of temperature andtransport of the western boundary current. Finally, thepossible coupling of the equatorial region with the datapoor midlatitudes of the Southern Hemisphere remains
to be investigated.
The subducted thermal anomalies originating in thecentral North Pacific appear to be distinct from the thermal signal associated with the acceleration of the North
Pacific gyre circulation. Miller et al. (1998) and Deseret al. (1999) document that the increase in wind stressover the North Pacific from the 1970s to 1980s led toa spinup of the subpolar and subtropical gyres in accordance with Sverdrup theory. Associated with thisspinup are western-intensified thermal anomalies ex-

tending down to at least 400 m in the Kuroshio Extension that can be seen in Fig. 1 between 408 and 508N.
The importance of this gyre spinup and the thermal
anomalies within the thermocline for a mechanism of
decadal oscillations ‘‘à la Latif and Barnett (1994,
1996)’’ remains to be determined.
The propagation and forcing of the thermal anomalies
in the thermocline bear striking resemblance to the second-mode motions in the thermocline discussed by Liu(1999a). This is yet another piece of evidence that dynamics of the ventilated thermocline are at the heart ofthe dynamics of decadal anomalies in the midlatitude
North Pacific thermocline. However, due to the lack ofsalinity observations we cannot say whether the subducted temperature anomalies are salinity compensated(with no anomalous velocity field) or are high-mode
baroclinic waves (which have anomalous velocity)
whose propagation path and speed are close, but not
identical, to advection by the mean circulation. Efforts
by T. Suga (1997, personal communication) to obtainadequate historical salinity observations may resolve
this issue. The statistical significance of the analysis
presented here is obviously limited by the small numberof observed realizations. Analysis of targeted experiments with numerical models (Venzke 1999; Inui 1997;
Liu 1999; Miller et al. 1994) are therefore needed toinvestigate fully the dynamics, forcing, and possible
feedbacks with the atmosphere of these thermal anomalies.
Acknowledgments. Discussions with Bruce Cornuelle,David Pierce, Tim Barnett, Paola Cessi, and Zhengyu
Liu are gratefully acknowledged. The authors thankW. B. White for generously making available the subsurface temperature analysis. This work was supportedby the National Science Foundation (NSF OCE9711265), by the National Oceanic and Atmospheric Administration as part of the Experimental Climate Prediction Center (NA77RJ0453), and the Consortium for
the Ocean’s Role in Climate (NA47GP0188).
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