The mean and time-varying meridional transport of heat at the
boundary of the North Pacific Ocean
Dean Roemmich, John Gilson, Bruce Cornuelle and Robert Weller
Ocean heat transport near the tropical/subtropical boundary of the North
Pacific during 1993-1998 is described, including its mean and time variability.
Twenty-five transpacific high resolution XBT/XCTD transects (Fig. 1) are
used together with directly measured and operational wind estimates to
the geostrophic and Ekman transports. The mean heat transport across the
XBT transect is 0.77 ± 0.12 pW. The large number of transects enables
a stable estimate of the mean field with realistic error bars based on the
known variability. The North Pacific heat engine is a shallow meridional
overturning circulation that includes warm Ekman and western boundary current
components flowing northward, balanced by southward flow of cool thermocline
waters (including Subtropical Mode Waters). A near-balance of geostrophic
and Ekman transports holds in an interannual sense as well as for the time
mean. The interannual range in heat transport was about 0.3 pW during
with maximum values of about 1 pW in early 1994 and early 1997. The repeating
nature of the XBT/XCTD transects, with direct wind measurements, allows
a substantial improvement over previous heat transport estimates based on
one-time transects. A global system is envisioned for observing the
ocean heat transport and its possible feedback in the coupled climate
Figure 1. The PX37/10/44 ship track (solid line from San Francisco to Honolulu
to Guam to Taiwan) is shown together with the ship track from the 24°N
hydrographic transect (black line), and ECMWF air-sea heat flux (w/m2) averaged
over the period 1993-1998. Positive numbers indicate heat loss by the ocean.
The red line shows the ship track from the WOCE 24 N CTD section.
The surplus of solar heating in the tropics and the corresponding deficit
in polar regions gives rise to time-varying atmospheric and oceanic
that carry enormous amounts of heat poleward from the low latitudes.
of the Earth's radiation budget (Stephens et al, 1981, Trenberth
and Solomon, 1994) indicate that about 5 pW is carried northward across
24°N by the combined atmosphere and oceans. Estimates of the oceanic
component, based on either oceanic measurements (Bryden et al, 1991)
or on the residual of atmospheric transport and top-of-atmosphere radiation
budgets (Trenberth and Solomon, 1994), have converged to about 2 pW at
However, the combined and individual estimates for ocean basins still have
very large errors - 0.3 pW or greater. An improved understanding of the
coupled climate system requires that the atmospheric and oceanic contributions
to the planetary heat balance be known with better accuracy than at present.
Further, it is necessary to investigate the low frequency variability of
the coupled system. If there is large interannual variability in ocean heat
transport, then the feedback effects on the atmosphere must be considered.
The present work focuses on the mean and time-varying heat transport of
the North Pacific Ocean, using repeating zonal transects near the
Estimates of ocean heat transport in the North Pacific (Fig 2) have been
made from hydrographic sections at 24°N (Roemmich and McCallister,
1989, Bryden et al, 1991, Macdonald and Wunsch, 1996) and 10°N
(Wijffels et al, 1996, Macdonald and Wunsch, 1996). These and additional
estimates based on basin-integrals of air-sea flux from climatological data
(DaSilva et al, 1995) and from operational analyses (National Center
for Environmental Prediction, NCEP, Kalnay et al, 1996, and European
Centre for Medium Range Weather Forecasts, ECMWF, 1993) are shown in Fig
1. The large error bars on the hydrographic estimates and the large spread
of values between the climatological and operational analyses of air-sea
fluxes illustrate the high uncertainty in either form of measurement. In
the case of the estimates based on hydrographic transects, most of the
derives from two sources:
(1) Upper ocean geostrophic variability. At 10°N and 24°N in the
Pacific, the meridional heat transport is dominated by shallow circulation.
Bryden et al (1991), referring to the top 700 m, note "The upper water
circulation carries essentially all of the heat transport across
However, the upper layers also have substantial temporal variability of
geostrophic circulation that results in an error of unknown magnitude in
heat transport estimates based on single hydrographic transects. In order
to quantify and reduce this error, we have collected a large number (presently
27) of eddy-resolving boundary-to-boundary transects using expendable
(XBT) and expendable conductivity-temperature-depth (XCTD) profiles to 800m
depth. Using the quarterly cruises spanning the Pacific at average latitude
of 22°N (Fig 2), the temporal mean and variability of upper ocean
transport is estimated.
(2) Ekman transport. Ekman transport makes a large contribution to heat
flux at these latitudes because volume transports are substantial (>
10 Sv) and the surface layer is very warm. However, previous estimates use
either climatological wind stress or operational analyses to estimate the
Ekman contribution, with unknown systematic errors in both cases. We addressed
this problem by installing a high quality anemometer on the same ship that
collects XBT/XCTD data. The anemometer is used here for calibration and
correction of systematic errors in ECMWF winds in order to estimate the
mean and variability of Ekman transport for the same time period as the
geostrophic transports from XBT/XCTD data.
Errors in the present analysis remain substantial, about 0.1 pW. However,
these errors will diminish further as the time-series is extended and the
profile measurements are supplemented with new and deeper XBTs. The study
illustrates a fundamental limitation of one-time hydrographic surveys. In
spite of the high value of such surveys, they cannot provide reliable estimates
of mean ocean heat transport. Moreover, the variability in the heat budget
is also of great intrinsic interest. It is now practical to estimate all
components of the oceanic heat budget with accuracy that was not achievable
a few years ago. These estimates are potentially of great value in
seasonal to interannual variability in the coupled climate system.
Figure 2. Estimates of northward heat transport from integrals of air-sea
flux and from zonal hydrographic transects in the North Pacific. The air-sea
flux estimates, integrated from 63°N to the latitude shown, include
an estimate from the COADS climatology (Da Silva et al, 1995), plus
estimates from NCEP (Kalnay et al, 1996) and ECMWF (ECMWF, 1993)
operational models for the years 1993-1998. The hydrographic estimates,
including error bars, are based on one-time sections at 10°N (W96 is
Wijffels et al, 1996) and 24°N (B91 is Bryden et al,
1991). The estimate from the present study is also shown, 0.77 ± 0.12
pW for 1993-1998, at the equivalent latitude (See Section 5, Eq. 1) appropriate
for comparison with ECMWF.
2. XBT/XCTD transects
Sampling was initiated in September 1991 on SS Sea-Land Enterprise, a container
ship operating along a track from San Francisco to Taiwan via Honolulu and
Guam (Fig 2). A scientist rides on the ship for the 17-day crossing
every three months, collecting an eddy-resolving temperature transect using
Sippican Deep Blue (800 m) XBTs. About 305 temperature profiles are obtained
on each voyage, with probe spacing ranging from about 10 km near the western
boundary to 50 km in mid-ocean. Sampling terminates near the 200 m isobath
at both endpoints, with closely spaced probes near topography. Salinity
sampling with XCTD probes was begun in 1994 and about 18 XCTDs are deployed
on each cruise to determine large-scale temperature/ salinity characteristics
and variability. Processing of the XBT and XCTD data, including fall-rate
correction and interpolation onto a uniform grid (with spacing 0.1°
of longitude by 10 m depth) was described by Gilson et al (1998). Geostrophic
velocity was calculated from gridded specific volume, using a reference
level at 800 m. The specific volume calculation uses salinity that is estimated
on a cruise-by-cruise basis from XCTD profiles when available (Gilson et
Through January 1999, 27 XBT/XCTD transects have been completed on the
Initial cruises were at 6-month intervals, changing to approximately 3-month
intervals at the end of 1992. Here, we will consider the 25 quarterly transects
from November 1992 to January 1999 as representing the period 1993-1998.
3. Wind, wind stress and Ekman transport
An anemometer (R.M. Young 5103, propeller/vane) was installed above the
bridge of the Enterprise in 1995 and has operated since that time. GPS
is used to remove ship motion from the relative wind measurements. A second
anemometer was installed near the bow on the ship's foremast in 1998, a
location thought to be less subject to disturbance of the airflow by the
ship's structure. The anemometer is sampled and recorded once per minute
and is vector averaged to 6-hourly intervals. The 6-hourly ECMWF data are
interpolated in space, using a weighted average of the adjacent nine grid
points, to the instantaneous location of the Enterprise. In this and all
other transects the anemometer wind speed and direction are highly correlated
with ECMWF analyses.
A comparison of all 6-hourly winds from the Enterprise, corrected as noted
above, with the corresponding 6-hourly ECMWF winds is shown in Fig 3. The
figure also shows the average ECMWF wind speed in each 0.5 m/s bin of
wind speed. A systematic difference at high wind speed is clear, with ECMWF
values being lower than anemometer wind by 2.6 m/s at 15 m/s wind speed.
There is no systematic difference in wind direction, which is highly correlated
(R2=0.89) in the two datasets. A correction based on the regression line
in Fig 3 was applied to the ECMWF wind data, uniformly along track.
Wind stress is calculated from corrected ECMWF 6-hourly winds using the
drag coefficient formulation of Yelland et al (1998). The details of this
formulation differ from that of Large and Pond (1981). However, in the present
case, the mean Ekman transport across the XBT/XCTD transect differs by only
a few percent depending on which is used. The particular choice of drag
coefficient between these two does not have a large impact on the
Ekman transport across the ship track is tx/rf, where tx is the component
of wind stress parallel to the ship track, r is the water density, and f
is the Coriolis parameter. Ekman transport estimates were computed using
corrected ECMWF wind, and averaged over 1-month intervals from January 1993
to December 1998. The mean over 72 months is 15.9 ± 0.8 Sv northward.
A 12-month running mean is shown in Fig 6 (red line) to illustrate the
variability in Ekman transport. The interannual oscillation ranges between
13 and 20 Sv with maxima in early 1994 and early 1997.
The temperature associated with Ekman transport is ambiguous since the
depth of the directly wind-driven flow is not known. For the present XBT
dataset, we take the temperature at 5 m to represent the sea surface
and the temperature of the Ekman transport. A second calculation, with the
Ekman layer assumed to decay linearly over the top 50 m, produced a decrease
in the average temperature of the Ekman layer by 0.2°C and a decrease
in the net heat transport by 0.01 pW. This difference is not significant.
Figure 3. Regression plot of bridge anemometer wind speed (6-hourly values
from all transects, corrected downward by 5% for height and location bias)
versus ECMWF as small green symbols. Large black symbols show average values,
in bins of 0.5 m/s width of the bridge anemometer wind speed. Also shown
are a line having unit slope and a line that is the best fit to the 6-hourly
4. Geostrophic transport
The 25-cruise mean geostrophic transport is 17.5 ± 0.8 Sv southward
in the upper 800 m. In this transect, the northward flow of the Kuroshio,
about 21.5 Sv, is over-balanced by 39.0 Sv of southward flow in the interior.
Of the latter, 10.1 Sv of southward transport occurs to the east of Hawaii.
The maximum net southward transport in a single cruise was 26.5 Sv in November
1998, and the minimum was 11.1 Sv, in November 1995. This large spread
the potential hazard of using one-time transects as representative of the
mean transport. Transport variability is described in Sections 6-7.
Fig 4a illustrates how the mean of geostrophic plus Ekman transport is
in temperature classes. The southward geostrophic transport shows a broad
maximum centered in the 17-18° band. Waters in this temperature class
are the Subtropical Mode Waters (STMW) formed in both the western (15-19°C
waters, e.g. Masuzawa, 1969) and eastern (16-22oC, Hautala and Roemmich,
1998) Pacific. The broad temperature range of southward transport is consistent
with these STMWs plus the colder Central STMW (9-13°C) described by
Suga et al (1997). Cruise-to-cruise standard deviations (Fig 4a,
red bars) in the thermocline are substantial, but smaller than the mean
transports. For example, combined transport in the 15-22° range is
9.0 Sv with cruise-to-cruise standard deviation of 2.7 Sv. The higher
at warmer temperatures is primarily due to seasonally changing temperature
in the surface layer rather than to changing velocity. Interestingly, the
maximum southward transport in the 24°N hydrographic transect (Fig
4b) occurred at a warmer 20°C, presumably another artifact of selecting
a single realization. The contrast of Fig 4a and 7b again demonstrates the
necessity of multiple realizations in order to determine the mean and
of basin-wide geostrophic transport.
Transport in salinity classes (Fig 4c) shows northward flow concentrated
around a salinity of 34.75, between two bands of southward transport. There
is a strong southward maximum at 35.0, characteristic of the subtropical
gyre interior. A band of fresher transport, mostly between 33.0 and 34.0
is due to the California Current system carrying waters of subpolar origin.
Transport in density classes (Fig 4d) shows the same pattern as Fig 4a,
indicative of the strong control of density by temperature at these latitudes.
The thermal overturning circulation carries heat and buoyancy northward,
while the transport of freshwater appears small.
Figure 4. a. Geostrophic plus Ekman transport in 1°C temperature bins.
Black bars show the mean values from 25 cruises. The width of the red bar
beyond the end of the black bar is the standard deviation. b. Geostrophic
plus Ekman transport from the one-time hydrographic transect at 24°N
(Bryden et al, 1991). c. Same as a) but for salinity bins. d. Same
as a) but for density bins of sigma-theta.
The volume transport (Fig 4) includes both the gyre-scale parts of that
field and components due to boundary currents and eddies. The characteristic
structure and transport of eddies, including their interannual variability,
is the subject of a separate study (Roemmich and Gilson, 1999).
Fig 9 (from that work) shows eddy locations identified independently in
the T/P and XBT datasets. The coincidence of features in the two datasets
is remarkable. The T/P dataset allows individual eddies in the XBT cruises
to be clearly tracked for a year or longer. In that time, they move thousands
of kilometers to the west at about 10 cm/s. It is the diagonal nature of
the ship track (Fig 1) toward Guam and Taiwan that limits the ability to
follow individual features even farther westward. A separate map of T/P
height along constant latitude of 22°N (not shown) demonstrates that
individual features can be tracked continuously from near Hawaii to the
western boundary. In Fig 9, the decrease in eddies near 150° E is clearly
associated with the southward dip of the ship track toward Guam, away from
the eddy-rich latitudes near 20°N (Fig 1).
The eddy transport is significant, enhancing the thermal overturning
by about 4 Sv in the mean (due to correlation of velocity anomalies with
layer thickness anomalies) and accounting for a large fraction of the
variability in southward thermocline transport. For the present work, it
suffices to say that the eddy transports are included in these calculations.
They are embedded in the boundary-to-boundary integrals of the geostrophic
flow and are a part of the signals described here.
5. Closure of the mean mass and heat budgets
The 1.6 Sv difference between the net northward Ekman transport and the
net southward geostrophic flow could be due to random sampling errors or
systematic error in the estimate of Ekman transport. If it is not one of
these errors, then the additional 1.6 Sv of northward transport needed to
complete the mass balance is due to barotropic transport or to deep baroclinic
shear (below 800 m) not sampled by the XBTs. We will close the mass budget,
assuming that the imbalance is not due to sampling errors, and then consider
the additional uncertainty due to errors. With respect to estimating heat
transport, the extreme possibilities are as follows:
1a) The additional 1.6 Sv of northward transport is all in the upper 800
m in the relatively warm western boundary. In other words, suppose there
is mean shear in the Kuroshio just below the 800 m level. Then, the appropriate
temperature for the balance is the 0-800 m average temperature, 13.9°C
(120.6°E to 123.8°E)
2a) The 1.6 Sv is all in the barotropic component (3.6°C, Bryden et
al, 1991). Closure of the mass balance in these two scenarios results in
northward heat transport estimates of 0.80 and 0.73 pW respectively.
The above range of heat transports must be expanded to reflect the random
sampling uncertainty. Again, we select the possibilities that produce the
1b) Change scenario 1a) above by increasing the northward Ekman transport
by the standard error, 0.8 Sv at 27.1°C, balanced barotropically at
2b) Change scenario 2a) by decreasing Ekman transport by 0.8 Sv, again with
a barotropic mass balance.
This expands the range of values to 0.88 and 0.65 pW respectively. We choose
the central value as the best estimate, hence, Mean meridional heat transport
= 0.77 ± 0.12 pW. This uncertainty does not include the possibility
of systematic errors in the wind field, which are discussed in Section 8.
The estimate given above, 0.77pW ± 0.12, is for ocean heat transport
across the XBT transect. For comparison with other techniques, it is desirable
to give an estimate of heat transport across a constant latitude. To do
this, we define an "equivalent latitude", YE, across which the
heat transport is the same as it is across the XBT track. YE is about
for the ECMWF (1993-1998) values of QNET, 20°N for NCEP (1993-1998),
and 22°N for the COADS climatology of DaSilva et al, 1995. In Fig 1,
the present estimate is plotted at the ECMWF equivalent latitude, where
it is in very good agreement with the ECMWF air-sea heat flux integral.
By shifting the estimate to 20°N or 22°N, it is seen to be slightly
inconsistent with the NCEP and COADS values.
6. Seasonal variability
The annual cycles of geostrophic and Ekman transport are shown in Fig 5.
For the geostrophic component (Fig 5a), accurate averages for every month
cannot be obtained because of the small number of cruises. With 25 cruises,
there are only about two in a given month. We therefore averaged cruises
in a given month with those of the month before and after. Thus, Fig 5a
shows 3-month running averages based on about six cruises each. It is apparent
that none of the estimates is significantly different from the mean and
that there is not a large annual cycle in geostrophic transport.
Ekman transport (Fig 5b) shows a semi-annual behavior, with maxima in April
and November and minima in January and September. Here the means and standard
errors for each month are based on the six monthly average values for a
given month during 1993-1998. Wind variability is large, so the annual cycle
is not accurately determined from the 6-year dataset. However, climatological
data (Hellerman and Rosenstein, 1983) examined along this track show a similar
semi-annual pattern, amplitude, and phase, so the 6-year interval is thought
to be representative.
The absence of an annual cycle in baroclinic transport implies that the
mass balance for the annually varying Ekman transport is accomplished through
barotropic waves. This is not surprising, and is consistent with evidence
from TOPEX/Poseidon altimetric data. Baroclinic waves are seen to propagate
westward from Hawaii to Taiwan in a period of about 2 years (Chelton and
Schlax, 1996, Roemmich and Gilson, 1999). Adjustment of baroclinic transport
to changes in basin-wide wind forcing is expected on interannual but not
Figure 5. Annual cycle of geostrophic and Ekman transports. a. The mean
and standard error of geostrophic transport using all cruises in a sliding
3-month window (e.g. all cruises in Dec-Jan-Feb from any year, Dec 1992
- Jan 1999 for the Jan estimate) b. The mean and standard error of Ekman
transport, from corrected (see text) ECMWF values. The Jan value is the
mean of all Jan data 1993-1998, etc.
7. Interannual variability
The interannual pattern of Ekman transport (red line in Fig 6, the 12-month
running mean) is plotted together with the corresponding time-series of
geostrophic transport. The interannual geostrophic transport time-series
is obtained by applying a 4-cruise running mean to the transport from
cruises. The geostrophic and Ekman transports have interannual fluctuations
that are similar to one another but of opposite sign. There is an approximate
balance between northward Ekman transport and southward geostrophic transport,
with an oscillation period of about 3 years. There is also a suggestion
of a trend in geostrophic transport toward greater southward values and
in Ekman transport toward weaker northward ones.
Figure 6. Interannual variability of geostrophic and Ekman transports.
Ekman transport is the 12-month running mean (red line). Southward geostrophic
transport is the 4-cruise running mean (solid line and symbols). Southward
geostrophic transport from TOPEX/Poseidon is shown as the green line.
The time-series of geostrophic transport is based on quarterly XBT/XCTD
cruises (Fig 6), so an important question is whether this time-series is
badly aliased by unresolved high frequency variability. As a test, geostrophic
transport variability at the sea surface from the XBT/XCTD cruises is compared
to the same quantity derived from 10-day cycles of TOPEX/Poseidon altimetric
data interpolated onto the ship track (Gilson et al, 1998). Fig 7
compares XBT/XCTD-derived surface geostrophic transport (black line), with
TOPEX/Poseidon surface geostrophic transport (blue and red lines), The mean
transport cannot be determined from TOPEX/Poseidon, so that series is adjusted
to have the same mean as the XBT/XCTD series. The two measures of interannual
variability in surface geostrophic transport in Fig 7 show a similar 3-year
oscillation to the total geostrophic transport (Fig 6). At the times when
the two surface transport estimates are most different (e.g. mid-1994),
inspection of the datasets shows that the differences are due to Kuroshio
transport variability close to the western boundary rather than to temporal
aliasing. This transport variability is not adequately observed by
because there are no good altimetric data close to the inshore side of the
current at this location. In addition, if the TOPEX/Poseidon surface transport
time-series is smoothed with a 360-day running mean (not shown), it is very
similar to the sub-sampled version (red line in Fig 7). It is therefore
concluded that the interannual time-series of total geostrophic transport
(Fig 6) is not badly distorted by aliasing. The similarity of geostrophic
transport to Ekman transport is not simply a coincidence between heavily
filtered and noisy datasets.
Fig 7 also indicates that there is an annual cycle in the surface geostrophic
transport. The TOPEX/Poseidon series, smoothed by a running mean of 120
days, shows 6 peaks, one around January of each year. This is in contrast
to Fig 5, where an annual cycle was not detected in total (0-800 m) geostrophic
transport from XBT/XCTD data. The explanation of this disparity is that
the annual cycle is confined to depths above 100 m. It is detectable in
surface geostrophic transport from the XBT/XCTD data but not in the 0-800
m integrated transport.
Figure 7. Comparison of geostrophic surface transport from XBT/XCTD transects
(solid line, 4-cruise running mean) with TOPEX/Poseidon geostrophic surface
transport. (The blue line is the 120-day running mean from the complete
T/P dataset. The red line with symbols is sampled at the time of the XBT/XCTD
cruises, then subject to a 4-cruise running mean). The symbols mark the
The interannual time-series of Ekman and geostrophic transports (Fig 6)
are used to estimate heat transport (Fig 8a). At the time of each cruise,
the residual mass field (difference between northward Ekman and southward
geostrophic transport, filtered as in Fig 8a) is balanced in two different
ways, similarly to the time mean calculation in Section 5. One balance assumes
barotropic compensation at 3.6°C (red line), and the other assumes
that the balancing transport occurs in the upper 800 m in the western boundary
current, at an average temperature of 13.9°C (solid line). These two
possibilities for the unmeasured deep fields produce the greatest range
in heat transport. It is clear that interannual heat transport variability
is substantial, with changes on the order of 0.3 pW. The heat transport
pattern (Fig 8a) roughly follows that of Ekman and geostrophic transport
(Fig 6). It is modulated by variability in the temperature of the Ekman
layer and in the vertical structure of the geostrophic transport profile,
and importantly, by changes in the eddy transport (v'T'). The interannual
variability in eddy heat transport, from Roemmich and Gilson (1999), is
shown in Fig 8b. Its range is about 0.1 pW, a substantial fraction of the
total interannual heat transport variability. Moreover, the upward trend
in eddy heat transport tends to balance a decrease due to diminishing Ekman
transport (Fig 6) so that the total (Fig 8a) does not show a trend.
What is the significance of interannual anomalies in ocean heat transport
of order 0.3 pW? This can be addressed in a couple of different ways. First,
it amounts to a substantial modification of the meridional overturning
depicted in Fig 4. The amplitude of this circulation changes by more than
30% on interannual periods. Second, because the area of the North Pacific
to the north of the XBT/XCTD transect is about 3.6 x 1013 m2, the anomalous
heat transport amounts to nearly 10 w/m2 on average over the whole of this
area. This heating is undoubtedly distributed unevenly over the basin, so
it can be quite significant regionally - to cause anomalous warming or cooling
and possibly feedback to the atmosphere. Unlike the annual cycle variability,
which we argued is not highly significant because of the short displacement
of water parcels over a few months, the interannual anomalies penetrate
thousands of kilometers even at interior ocean velocities.
Figure 8. a. Time-series of interannual variability in northward heat
The three lines are for the barotropic (red line) and baroclinic (solid
line) mass closure schemes described in the text. The blue line is northward
heat transport as inferred from ECMWF heat fluxes integrated over the ocean
surface north of the XBT section. b. Interannual variability in eddy heat
transport (from Roemmich and Gilson, 1999), smoothed by 4-cruise running
Figure 9. Eddy locations in T/P and XBT data as a function of longitude
and time. Warm (cold) core eddies with sea surface height maxima (minima)
are shown as red (green) symbols for XBT data and gray (black) shading for
8. Discussion and conclusions
In the present work, the value of regularly repeating measurements of the
upper-ocean temperature and geostrophic shear fields has been demonstrated.
The mean of geostrophic transport in the upper 800 m, 17.5 ± 0.8 Sv
southward during 1993-1998, is well determined by the time-series of 25
cruises. This approximately balances the Ekman transport, 15.9 ± 0.8
Sv northward for the same 6-year period. The heat engine of the North Pacific
consists of northward Ekman transport of warm surface water plus northward
geostrophic transport in the warm Kuroshio. The balancing southward geostrophic
flows in the ocean interior occur at a broad range of thermocline temperatures
centered on 17-18°C Subtropical Mode Waters (Fig 4). This is the shallow
meridional overturning circulation of the subtropical North Pacific. Moreover,
the 6-year time-series demonstrates that geostrophic and Ekman transports
remain in approximate balance on interannual periods, while the amplitude
of the overturning circulation varies by over 30% (Fig 6).
The mean northward heat transport is 0.77 ± 0.12 pW across the XBT/XCTD
transect, which crosses the North Pacific at an average latitude of 22°N
(Fig 1). The error bounds on this estimate result from uncertainty in balancing
the 1.6 Sv difference between northward Ekman and southward geostrophic
flows and from random sampling errors in the 25-cruise, 6-year time series.
Systematic errors in the wind field have been addressed (Fig 3) but may
still be significant, and are not included in the error bounds.
It is interesting to note that the error bounds on heat transport assigned
to the one-time hydrographic surveys at 24°N and 10°N (Fig 2)
span a wide range of values including both the climatological estimate and
the operational model estimates for 1993-1998. The present study finds good
agreement with the net air-sea heat flux integral from ECMWF (Fig 2) but
the lower error bounds marginally exclude the estimates from NCEP and the
COADS climatology. This result is for a single region, so it should not
be taken as a general confirmation or correction of the flux products. However,
it illustrates the potentially powerful use of the oceanic heat transport
estimates for constraining or testing models (e.g. Wilkin et al,
1995) and operational analyses. Moreover, it clearly emphasizes the high
premium attached to accurate determination of the wind field. Errors in
the wind field affect both air-sea flux estimates (via latent heat exchange)
and the oceanic heat transport (via the Ekman layer).
Interannual variability in ocean heat transport was found to be of order
0.3 pW (Fig 8a) - amounting to nearly 10 w/m2 of anomalous heating/cooling
over all of the North Pacific to the north of this transect. This finding
highlights the need to close the time-varying heat budget including air-sea
fluxes and storage, with errors less than 10 w/m2 on interannual time-scales.
This will reveal the patterns of coupled variability in the air-sea system
and bring into focus the role of ocean circulation as a potential feedback
mechanism. For now, the similarity of geostrophic and Ekman transport
on interannual time-scales demonstrates that this climatically significant
variability in ocean circulation can be measured and tracked.
Collection of the XBT/XCTD data was supported by the National Science
through Grants OCE90-04230 and OCE96-32983 as part of the World Ocean
Experiment. TOPEX/Poseidon data were kindly provided by the Jet Propulsion
Laboratory, and the assistance of L.-L. Fu and A. Hiyashi is gratefully
acknowledged. Analysis was supported by the NASA JASON-1 project through
JPL Contract 961424. We are grateful for the dedicated efforts of many ship
riders in collecting the XBT/XCTD data, under the management of G. Pezzoli.
The development and production of VOS IMET modules has been carried out
by D. Hosom and is supported by NOAA Grant NA47GP0188 (JIMO Consortium).
The views expressed herein are the authors' and do not necessarily reflect
those of NOAA or its sub-agencies. We thank the officers and crew of SS
Sea-Land Enterprise for their continued assistance to this project. Chief
Mate M. Smith provided invaluable programming support in the installation
of the anemometers on the Enterprise. The National Center for Atmospheric
Research provided access to ECMWF analysis products. Graphics were produced
using FERRET, developed by NOAA/PMEL.
Bennett, A. and W. White, 1986: Eddy heat flux in the subtropical North
Pacific. J. Phys. Oceanogr., 16, 728-740.
Bryden, H., D. Roemmich and J. Church, Ocean heat transport across 24oN
in the Pacific. Deep-Sea Res., 38. 297-324, 1991.
Chelton, D. and M. Schlax, Global observations of oceanic Rossby waves.
Science. 272, 234-238, 1996.
DaSilva, A., C. Young and S. Levitus, Atlas of Surface Marine Data 1994,
vol. 1, Algorithms and Procedures. NOAA Atlas NESDIS 6, 83 pp., Natl. Oceanogr.
Data Cent, Silver Spring, Md., 1995.
European Centre for Medium-Range Weather Forecasts (ECMWF), The description
of the ECMWF/WCRP III - A global atmospheric data archive. Tech. Attachment
ECMWF 1-18, Reading, England, 1993.
Freilich, M. and R. Dunbar, The accuracy of the NSCAT 1 vector winds:
with National Data Buoy Center buoys. J. Geophys. Res., in press,
Gilson, J., D. Roemmich, B. Cornuelle and L.-L. Fu, Relationship of
altimetric height to steric height and circulation in the North Pacific.
J. Geophys. Res., 103, 27947-27965, 1998.
Hautala, S. and D. Roemmich, Subtropical mode water in the Northeast Pacific
Basin. J. Geophys. Res., 103, 13055-13066, 1998.
Hellerman, S. and M. Rosenstein, Normal monthly wind stress over the World
Ocean with error estimates. J. Phys. Oceanogr., 13, 1093-1104,
Hsu, S., E. Meindl and D. Gilhousen, Determining the power-law wind-profile
exponent under near-neutral stability conditions at sea. J. Applied
33, 757-765, 1994.
Kalnay, E., M. Kanamitsu, R. Kistler, W. Collins and 18 others, The NCEP/NCAR
40-year reanalysis project. Bulletin of the American Meteorological
77, 437-471, 1996.
Large, W. and S. Pond, Open ocean flux measurements in moderate to strong
winds. J. Phys. Oceanogr., 11, 324-336, 1981.
Macdonald, A. and C. Wunsch, An estimate of global ocean circulation and
heat fluxes. Nature, 382, 436-439, 1996.
Masuzawa, J., Subtropical mode water. Deep-Sea Res., 16, 463-472,
Roemmich, D. and J. Gilson, Eddy transport of heat and thermocline waters
in the North Pacific: A key to interannual/decadal variability. Submitted
to J. Phys. Oceanogr., 1999.
Roemmich, D. and T. McCallister, Large scale circulation of the North Pacific
Ocean. Prog. Oceanogr., 22, 171-204, 1989.
Stephens, G., G. Campbell, and T. Vonder Haar, Earth radiation budgets.
J. Geophys. Res., 86, 9739-9760, 1981.
Suga, T., Y. Takei and K. Hanawa, Thermostad distribution in the North Pacific
subtropical gyre: the central mode water and the subtropical mode water.
J. Phys. Oceanogr., 27, 140-152, 1997.
Trenberth, K. and A. Solomon, The global heat balance - heat transports
in the atmosphere and ocean. Climate Dynamics, 10, 107-134,
Weller, R. A. and S. P. Anderson, Surface meteorology and air-sea fluxes
in the western equatorial Pacific warm pool during the TOGA Coupled
Response Experiment, J. Climate, 9(8), 1959-1990, 1996.
Weller, R. A., M. F. Baumgartner, S. A. Josey, A.S. Fischer, and J. C. Kindle,
Atmospheric forcing in the Arabian Sea during 1994-1995: observations and
comparisons with climatology and models, Deep-Sea Research,
Wijffels, S., E. Firing and H. Bryden, Direct observations of the Ekman
balance at 10-degrees-N in the Pacific. J. Phys. Oceanogr., 24,
Wijffels, S., J. Toole, H. Bryden, R. Fine, W. Jenkins and J. Bullister,
The water masses and circulation at 10oN in the Pacific. Deep-Sea Res.,
43, 501-544, 1996.
Wilkin, J., J. Mansbridge and J.S. Godfrey, Pacific Ocean heat transport
at 24oN in a high-resolution global model. J. Phys. Oceanogr.,
2204 - 2214, 1995.
Yelland, M., B. Moat, P. Taylor, R. Pascal, J. Hutchings and V. Cornell,
Wind stress measurements from the open ocean corrected for airflow distortion
by the ship. J. Phys. Oceanogr., 28, 1511-1526, 1998.