Progress in Oceanography 53 (2002) 163–183
www.elsevier.com/locate/pocean
Altimeter-derived surface circulation in the large–scale NE
Pacific Gyres. Part 1. seasonal variability
P Ted Strub∗, Corinne James
College of Oceanic and Atmospheric Sciences, Oregon State University, 104 Ocean Administration Building, Corvallis, OR 973315503, USA
Abstract
The seasonal variability of sea surface height (SSH) and currents are defined by analysis of altimeter data in the NE
Pacific Ocean over the region from Central America to the Alaska Gyre. The results help to clarify questions about
the timing of seasonal maxima in the boundary currents. As explained below, the long-term temporal mean of the SSH
values must be removed at each spatial point to remove the temporally invariant (and large) signal caused by the
marine geoid. We refer to the resulting SSH values, which contain all of the temporal variations, as the ‘residual’ SSH.
Our main findings are:
1. The maximum surface velocities around the boundaries of the cyclonic Alaska Gyre (the Alaska Current and the
Alaska Stream) occur in winter, at the same time that the equatorward California Current is weakest or reversed
(forming the poleward Davidson Current); the maximum surface velocities in the California Current occur in summer.
These seasonal maxima are coincident with the large-scale atmospheric wind forcing over each region.
2. Most of the seasonal variability occurs as strong residuals in alongshore surface currents around the boundaries of
the NE Pacific basin, directly connecting the boundaries of the subpolar gyre, the subtropical gyre and the Equatorial
Current System.
3. Seasonal variability in the surface velocities of the eastward North Pacific Current (West Wind Drift) is weak in
comparison to seasonal changes in the surface currents along the boundaries.
4. There is an initial appearance next to the coast and offshore migration of seasonal highs and lows in SSH, alongshore
velocity and eddy kinetic energy (EKE) in the Alaska Gyre, similar to the previously-described seasonal offshore
migration in the California Current.
5. The seasonal development of high SSH and poleward current residuals next to the coast appear first off Central
America and mainland Mexico in May–June, prior to their appearance in the southern part of the California Current
in July–August and their eventual spread around the entire basin in November–December. Similarly, low SSH and
equatorward transport residuals appear first off Central America and Mexico in January–February before spreading
farther north in spring and summer.
6. The maximum values of EKE occur when each of the boundary currents are maximum.
 2002 Elsevier Science Ltd. All rights reserved.
∗
Corresponding author. Fax: +1-503-737-2064.
E-mail addresses: [email protected] (P.T. Strub); [email protected] (C. James).
0079-6611/02/$ - see front matter  2002 Elsevier Science Ltd. All rights reserved.
PII: S 0 0 7 9 - 6 6 1 1 ( 0 2 ) 0 0 0 2 9 - 0
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P.T. Strub, C. James / Progress in Oceanography 53 (2002) 163–183
Contents
1.
Introduction and background . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 164
2. Data and methods . . . . . . . . . .
2.1. Altimeter and tide gauge data .
2.2. Atmospheric forcing—sea level
2.3. Statistical gridding . . . . . . .
3.
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167
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Results . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170
4. Summary and discussion . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.1. Alaska Gyre . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
4.2. Connections around the boundaries of the subarctic and subtropical
4.3. Connections to the North Pacific Current . . . . . . . . . . . . . . . .
4.4. Offshore ‘propagation’ of the seasonal height and transport signals
4.5. Connections to the equatorial current systems along the boundaries
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gyres
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176
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1. Introduction and background
This is the first of a two-part analysis of temporal variability of the NE Pacific Ocean’s surface circulation, as measured by satellite altimeters. Here we examine the seasonal variability. In Part 2 (Strub &
James, 2002) we analyze the non-seasonal anomalies of the surface circulation over the 1993–1998 period,
during which the 1997–1998 El Niño creates the largest signal. Formation of the seasonal cycles discussed
here is the first step in creating the non-seasonal anomalies. The seasonal cycles themselves, however,
provide new information on the response of the NE Pacific to strong seasonal forcing, on scales not previously addressed. This analysis quantifies the degree of connection, on seasonal time scales, between the
boundary currents in the eastern subarctic and subtropical gyres, as well as the connection between the
boundaries and the interior NE Pacific. It further shows a connection to the equatorial current system.
Numerous papers describe aspects of the seasonal cycles for certain parameters in subregions of our
larger domain. Chapters in Robinson and Brink (1998) review some of the past results from the coastal
ocean in the regions between the Equator and the Alaska Gyre (Badan-Dangon, 1998; Hickey, 1998; Royer,
1998). Fig. 1 presents the climatological surface dynamic height field (relative to 500 m) in the NE Pacific,
calculated from the long-term mean climatological temperature and salinity data of Levitus and Gelfeld
(1992). The 500 m reference level is used to concentrate on the surface flow seen by altimeters. Although
this climatology is overly smooth, it shows the major currents in the area. The broad, eastward North
Pacific Current (also called the West Wind Drift) splits into the counterclockwise Alaska Gyre and the
equatorward California Current. South of 20°N in summer, the California Current turns westward and flows
into the North Equatorial Current, while in winter–spring, part of it continues along the Mexican mainland
before turning westward (Badan-Dangon, 1998; Fiedler, 1992, 2002). The long-term climatology shows
both paths. The North Equatorial Countercurrent (NECC) flows eastward between 5° –10°N to approximately 120°W, but is only weakly seen in the annual climatology from there to the cyclonic flow around
the Costa Rica Dome near 8°N, 92°W. The NECC is a shallow current (found in the upper 200 m) and
might appear more strongly if a shallower reference were used, but it is also seasonally intermittent. When
the Intertropical Convergence Zone (ITCZ) is in its northern location near 10°N (summer), surface divergences and upwelling create a zonal trough in surface height, driving the NECC along the southern side
of the trough. When the ITCZ moves south in winter, the NECC weakens or reverses.
P.T. Strub, C. James / Progress in Oceanography 53 (2002) 163–183
165
Fig. 1. Climatological dynamic height relative to 500 m in the NE Pacific. Based on the temperature and salinity climatology of
Levitus and Gelfeld (1992).
The most well-studied subregion of our larger domain is the California Current System, located between
approximately 20°–50°N within 1000 km of the coast of North America. In addition to reviews by Hickey
(1979, 1989, 1998), papers by Kelly, Caruso, and Austin (1993), Kelly et al. (1998), Brink et al. (2000)
and Strub and James (2000) describe aspects of the seasonal circulation and eddy statistics in this region,
using altimeter and surface drifter data, and provide reviews of previous work. Kelly et al. (1993), in
particular, addresses both the Alaska Gyre and the California Current and comes the closest to covering
the same range of territory and topics as in the present study.
The seasonal circulation of the California Current System has been best defined off central and southern
California, using the long CalCOFI data set (Chelton, 1984; Lynn & Simpson, 1987). Some of the early
CalCOFI cruises also cover northern California and are shown by Hickey (1979), but were not repeated
often enough to statistically define the seasonal cycle there. Hickey (1979) also presents fields derived
from ship drift records. Other studies have inferred seasonal changes in the circulation over the shelf or
farther offshore from shorter records of measured currents (Strub, Allen, Huyer, Smith, & Beardsley, 1987;
Wickham, Bird, & Mooers, 1987) or from specific cruises covering one or more years (Kosro et al., 1991;
Rienecker & Mooers, 1989). From these and other works, we expect an equatorward current that extends
from Vancouver Island to the Southern California Bight in summer (Strub & James, 1995; Strub, Kosro,
Huyer, & CTZ Collaborators, 1991), partially entering the cyclonic flow in the Bight and partially continuing along Baja California. An inshore poleward countercurrent develops next to the coast in mid-summer
(Lynn & Simpson, 1987; Strub & James, 2000) and eventually extends to the Canadian border in winter,
becoming the poleward Davidson Current. In spring, upwelling and equatorward flow begin in the south,
next to the coast, and this also expands offshore and to the north to create the summertime California
Current jet, completing the cycle (Strub & James, 2000). Thus, both equatorward and poleward phases of
the seasonal currents begin next to the coast off southern California and migrate offshore and to the north,
at least as far as Vancouver Island.
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The seasonal wind forcing over the California Current has been well-defined by Bakun and Nelson
(1991) at scales of 100–200 km and more, using merchant ship data. They show the generally upwellingfavorable (equatorward) winds south of approximately 37°N, maximum in summer. North of 37°N, there
are seasonally alternating poleward winds in winter and equatorward winds in summer associated with the
expansion of the Aleutian Low (winter) and North Pacific High (summer) pressure systems. Other papers
present similar seasonal cycles derived from measured and proxy coastal winds (Halliwell & Allen, 1987;
Strub et al., 1987), model winds (Strub, James, Thomas, & Abbott, 1990), or scatterometer winds (Thomas,
Carr, & Strub, 2001).
There are far fewer historical studies of the seasonal circulation in the Alaska Gyre (Royer, 1998). Much
of the past and recent work has focused on the Alaska Coastal Current (ACC), a narrow surface current
found within 20–50 km of the coast in the north and northwest part of the Gyre. It is driven by both fresh
water inputs along the boundaries and cyclonic wind systems over the basin, which help retain the fresh
water at the coast. The maximum transport is thought to occur in autumn, at the time of maximum fresh
water input, rather than at the time of maximum wind forcing in winter (Royer, 1981a; Stabeno, Reed, &
Schumacher, 1995). This would agree with the October–November maximum in coastal tide gauge sea
levels, which was first observed by Pattullo, Munk, Revelle, and Strong (1955). Velocities are maximum
at the surface next to the coast at speeds of 0.5–1.0 m s⫺1 or even greater (Johnson, Royer, & Luick,
1988; Stabeno et al., 1995). Mean transports are of order 0.6–0.8 Sv and are correlated with the wind
(Stabeno et al., 1995).
The Alaska Current (AC) flows in the region farther offshore and is what we expect the altimeter to
see. In the western half of the basin, it is called the Alaska Stream. Royer (1981b) uses a series of hydrographic sections along the western boundary of the Gyre to calculate the total transport (relative to 1500
m) in the AC as 9.2 Sv, with a seasonal variation of ±1.2 Sv. He finds the maximum transport to occur
in late-winter or spring (March or even later), lagging the wind forcing. Bograd, Thomson, Rabinovich,
and LeBlond (1999) report a somewhat different result, based on analysis of 5 years of WOCE drifter
tracks in the Alaska Gyre. Along the shelfbreak in the northern gyre between 150°–180°W, there are twice
as many high speed observations (velocities over 0.4 m s⫺1) in winter as in summer, with most of those
coming from the very surface (drifters that had lost their drogue). They define winter as October–March,
which leaves a fair degree of uncertainty as to the time of the actual maximum. Bograd et al. (1999) also
report a maximum in mean kinetic energy in winter (40 cm2 s⫺2) in the North Pacific Current at 50°N,
150°W, compared to values less than 20 cm2 s⫺2 in summer (a decrease of 40–50%).
In the eastern Gyre, the measurements are even more sparse and seem more often to indicate a vigorous
eddy field rather than a continuous flow (Crawford, Cherniawsky, & Foreman, 2000; Crawford & Whitney,
1999; Meyers & Basu, 1999; Royer, 1998; Tabata, 1982; Thomson & Gower, 1998). Royer (1998) characterizes the flow next to the coast off Vancouver Island and farther north as entirely poleward, even in
summer, based on measurements of a bouyancy driven current in the 20 km next to the Vancouver coast
in summer (Hickey, Thomson, Yih, & LeBlond, 1991). Hickey et al. (1991), however, show equatorward
currents of up to 15 cm s⫺1 offshore of this current. Thus, the degree of connection between the eastern
Alaska Gyre and the California Current in summer is not well-defined by earlier studies.
Previous altimeter analyses of the Alaska Gyre have used Geosat data. Combining data from the Geodetic
and Exact Repeat Mission (ERM) periods to form a nearly 4-year time series, Bhaskaran, Lagerloef, Born,
Emery, and Leben (1993) identify an annual mode through the use of Empirical Orthogonal Function
(EOF) analysis. This mode consists of a cell of low SSH values north of 55°N, maximum in December–
January. Other cells are found south of 55°N. This mode is correlated with the first EOF of atmospheric
sea-level pressure (SLP), with no lag. Bhaskaran et al. (1993) suggest that this could either represent Ekman
pumping of the cyclonic Gyre or errors in the atmospheric corrections to the altimeter data, which use the
SLP. Kelly et al. (1993) use the 2.5 years of ERM data to form seasonal cycles of SSH by fitting annual
and semiannual harmonics. The patterns presented (their Fig. 9) differ from Bhaskaran et al. (1993) and
P.T. Strub, C. James / Progress in Oceanography 53 (2002) 163–183
167
indicate SSH values that decrease toward the north of the Gyre in January, suggesting a decrease in the
Gyre’s strength in winter. The SSH residuals become positive in summer and autumn, implying an increase
in the Gyre’s strength. Kelly et al. (1993) interpret this to represent an autumn maximum in the Gyre’s
strength, as opposed to Bhaskaran et al.’s (1993) winter and Royer’s (1981b) spring maximum. One goal
of this study is to clarify this timing, using the longer record of the more accurate TOPEX altimeter.
In the southern part of our domain off Central America, Badan-Dangon (1998) summarizes the few
direct measurements and Fiedler (1992, 2002) provides useful gridded maps of physical variables between
30°N–20°S. These include seasonal maps of ship-drift surface velocities, winds and thermocline depth (the
depth of the 20° isotherm). Seasonal residuals (the deviation from the long-term mean) of ship-drift surface
velocities and thermocline depth provide the closest comparison to the altimeter residual heights. These
show a decrease and reversal of the NECC in winter-spring, after the ITCZ moves south, and an increase
in NECC strength in summer–autumn, after the ITCZ moves north, in agreement with other observations
(Badan-Dangon, 1998; Tomczak & Godfrey, 1994). Current residuals next to the coast between 10°N–
30°N are equatorward in winter and early spring (the Costa Rica Current is weak or reversed); poleward
velocities are maximum in summer and autumn.
Putting these various descriptions together, we expect the counterclockwise boundary currents in the
Alaska Gyre to be strongest some time between autumn and spring (depending on the reference). The
altimeter fields will not usually see the Alaska Coastal Current but should resolve the broader Alaska
Current in the east and the Alaska Stream in the west. Drifter observations suggest a moderate winter
maximum in the eastward North Pacific Current, part of which serves as the southern boundary of the
Alaska Gyre. In the California Current, we expect the development and offshore/northward movement of
a spring–summer equatorward jet and winter poleward flow north of the Southern California Bight. Off
Central America, we expect the eastward North Equatorial Countercurrent and the poleward Costa Rica
Current to be strongest in summer and autumn.
Thus, the goals of the examination of seasonal variability in the surface circulation in the NE Pacific
are: (1) To determine the timing and structure of the seasonal maximum in transport in the Alaska Gyre;
(2) To determine the extent and timing of direct transport between the subtropical and subarctic gyres
along the boundary currents; (3) To determine the nature of the seasonal variability in the strength and
position of the North Pacific Current and its contribution to the seasonal cycles of the boundary currents
in the subarctic and subtropical gyres; (4) To look for connections between the equatorial currents and the
boundary currents farther north.
2. Data and methods
2.1. Altimeter and tide gauge data
The majority of altimeter studies use gridded fields of SSH, inferring the geostrophic circulation as
occurring along constant heights. Here we also use geostrophic transports calculated between altimeter
crossover points. The transports use the heights as recorded on the altimeter tracks directly, avoiding the
effects of statistical gridding, although some alongtrack smoothing is applied. The height differences also
reduce the effects of most environmental errors and seasonal steric changes in height, which have scales
larger than the 100–250 km differences in locations of the grid points. Over the open ocean, the SSH
values come only from the TOPEX and POSEIDON altimeters on the TOPEX/POSEIDON (T/P) satellite.
The TOPEX altimeter provides most of the data and has the highest signal to noise ratio of all altimeters.
Next to the coast where high quality tide gauge data are available, altimeter data from crossovers 200–
300 km offshore are combined with SSH values from the coastal tide gauges (Fig. 2). This allows a
calculation of the boundary transports with no gap between the altimeter data and the coast. This is done
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Fig. 2. T/P Altimeter tracks and sea level locations (filled circles). The crossover points used in the transport calculations are
indicated by filled squares.
because the altimeter data become noisier on the parts of the altimeter tracks within 25–50 km of the coast.
Where no tide gauge data are available, the altimeter must be used for the most inshore height measurement,
leaving a gap of order 20–50 km next to the coast.
Fig. 2 shows the T/P ground tracks, along which altimeter height data are available approximately every
6 km, repeating every 10 days. This provides approximately 3 height samples per month along each track,
with 6 height samples each month at the crossover points. The tracks are separated by 316 km at the
equator, which goes to zero at the inclination latitude of 66°. The T/P data used to calculate transports are
processed at JPL and made available by V. Zlotnicki. Standard environmental corrections are applied
(inverted barometer, wet and dry troposphere, ionosphere and EM bias). Ocean tides are removed using
the UT 3.0 tide model; poletides and solid earth tides are also removed. The data are regridded to a regular
alongtrack grid with 6 km spacing. We form and remove a five-year (1993–1997) ‘long-term’ mean from
each data point to remove the marine geoid and create a residual sea surface height (SSH) data set used
in calculating transports. The crossover points used to calculate transport are indicated by filled squares
in Fig. 2.
To avoid missing features in the surface circulation due to the choice of the specific grid used here to
calculate the transports, a different altimeter data set is used to calculate gridded 2-D fields of seasonal
SSH for 2-month calendar periods. We use alongtrack SSH data from four satellite missions included in
the NASA/NOAA Pathfinder data set: the T/P mission during the period October 1992–December 1997:
the GEOSAT mission from November 1986–September 1989 (although data dropouts become numerous
after autumn 1988), the ERS-1C mission from April 1992–December 1993, and the ERS-2 mission from
April 1995–June 1997. The Geosat and ERS data provide higher spatial resolution and a longer data set
to define the annual cycle of the 2-D SSH gradient fields. Their processing is similar to the processing of
the primary T/P data set from JPL. See Strub and James (2000) for plots of the tracks for all three altimeters
in the California Current region.
P.T. Strub, C. James / Progress in Oceanography 53 (2002) 163–183
169
In order to combine the heights from different satellites, it is necessary to first remove any offsets in
mean height between the orbits of the satellites. The approach we have found most successful is to remove
the spatial mean from each complete data set for each specific period covered (here the two-month calendar
periods for all years of data) before gridding. Besides eliminating height biases between the data sets, this
procedure also removes the spatial average of the large-scale mean seasonal cycle in height, which is
dominated by hemispheric steric height changes associated with seasonal heating (Leuliette & Wahr, 1999;
Stammer, 1997). What remains are the two-month fields of height gradients, displayed as contours of height
in a field with zero spatial mean. Before removing the mean, we also eliminate outliers using a 2-iteration
scheme that examines the time series at each point and removes points that deviate by more than 4 and
then 5 standard deviations from the mean (this primarily removes noisy points next to the coasts).
Hourly tide gauge data were obtained from NOAA/NOS and the University of Hawaii. Tides are removed
using a 46-hour half power filter and the hourly data are averaged to form a daily value. An inverse
barometer correction is applied using sea level pressure data from the U.S. Navy atmospheric model
(provided by the NOAA Pacific Fisheries Environmental Group). A 20-day filter is then applied to the
corrected daily average sea level to reduce the effect of coastally trapped waves, which have a strong
signal at the coast (in the tide gauge data) and much smaller signals 25–50 km offshore in the altimeter
signals. Temporal averages from the same 5-year period as used with the altimeter data are removed from
the tide gauge time series. Finally, this daily time series is interpolated in time to the midpoint of each
T/P cycle. The tide gauge data are from five locations along the California Current between 35°–48°N and
at three locations around the Alaska Gyre (Fig. 2).
Monthly geostrophic surface velocities over the large-scale NE Pacific are calculated from the height
differences between selected crossovers and tide gauge locations. The velocities are multiplied by the
distance between points to form surface transports. The reader can multiply by an assumed depth of the
current to produce volume transports. Since the distance between crossovers decreases from approximately
250 km at the lower latitudes to 100 km at the highest latitudes, the scale arrow in some figures showing
5000 m2s⫺1 represents velocities of approximately 2 cm s⫺1 at low latitudes and 5 cm s⫺1 at high latitudes.
The selection of the location of the points used here is the result of tests of different grids, with the
goal of defining the boundary currents in the two gyres and the zonal currents in the interior North Pacific
Current, in as simple a fashion as possible. The final choice of points defines the alongshore boundary
currents around the basin, using 3–4 points oriented roughly perpendicular to the coast. The onshore–
offshore currents into this band around the basin margin are also included, to look for changes in the inflow
to the boundary currents. Finally, four north–south lines are included to monitor the zonal interior flow
between 140°–150°W north of 46°N, a pair near 141°W and a pair near 146°W (Fig. 2). These lines allow
us to see any connections between the zonal interior flow and the inflow to the margins. The geometry of
adjacent pairs of north–south altimeter lines produces east–west transports that are staggered in space,
increasing meridional resolution.
2.2. Atmospheric forcing—sea level pressure
Over the large-scale NE Pacific basin, surface momentum flux is due to winds associated with the
atmospheric sea level pressure (SLP) fields. We use the pressure fields from the NCEP/NCAR Reanalysis
atmospheric models to represent the basin-scale forcing, over the five complete years of JPL altimeter data
(1993–1997). Daily fields are used to form the mean fields presented here.
2.3. Statistical gridding
Five-year means (1993–1997) of SSH are formed at each along-track grid point and subtracted from the
time series at each grid point. This removes the marine geoid (which is the largest SSH signal and must
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be removed) and also the temporal mean dynamic SSH and circulation field. The resulting ‘residual’ SSH
data sets contain all of the temporal variability, including the seasonal cycles. These are used to construct
the time series of 2-D horizontal fields of residual SSH with 10-day spacing that are used in the EOF
analysis (centered on each time step using 35-days of data, in order to include all of the ERS tracks). The
gridding method is that of successive corrections (Bratseth, 1986; Vazquez, Zlotnicki, & Fu, 1990). In this
method, heights are estimated at each point on a regular grid (0.5° is used here). A weighted mean of the
data surrounding each grid point forms a first guess of the height at the grid point, which is modified
during three iterations, reducing the spatial scale of the quadratic weighting function from 1.25° (first guess)
to 1.0°, 0.75°, and 0.5° (Vazquez et al., 1990). The resulting field is smoothed with a final Laplacian filter,
which reduces the amplitude of features with scales less than approximately 80–100 km. See the Appendix
of Strub, Chereskin, Niiler, James, and Levine (1997) for a discussion of the statistical aspects of this
gridding. The same data set is used to form the seasonal means of residual SSH, also using successive
correction, but combining all data from a given season from all altimeters.
3. Results
Fig. 3 shows the 3-month seasonal residuals of surface geostrophic transports between all chosen points,
based on five complete years, 1993–1997. Winter consists of January –March, etc. Transports of 2000 m2
s⫺1 are darkened to emphasize the dominant circulation patterns. Similar 3-month seasonal means formed
from only the first three and four years of data produce similar fields. When the 5-year mean altimeter
height is subtracted to remove the marine geoid, the mean circulation is also removed. To restore the mean
Fig. 3. Seasonal residuals of transport between selected T/P crossovers and tide gauges. Winter is JFM, spring is AMJ, summer is
JAS and autumn is OND. Based on 5 years, 1993–1997, after removing the 5-year mean from all heights. Transports greater than
2000 m2 s⫺2 are darkly shaded to highlight the main pattern of flow.
P.T. Strub, C. James / Progress in Oceanography 53 (2002) 163–183
171
circulation, one should mentally add back a more intense version of the circulation shown in Fig. 1: cyclonic
flow around the boundaries of the Alaska Gyre, equatorward flow in the California Current to around 20°N,
poleward flow in the Costa Rica Current between 5°–15°N, and eastward flow in the North Pacific Current
between 35°–50°N and the NECC between 5°–10°N.
The winter transport residuals show a strong counter-clockwise Alaska Gyre and a weak or reversed
California Current (the poleward Davidson Current) as far south as Baja California. Strengthened eastward
return flow occurs along the southern flank of the Alaska Gyre in the interior ocean at scattered locations
between 50°–57°N. Equatorward flow is found in the offshore regions of the California Current. In spring,
equatorward transport residuals appear next to the California coast, strongest south of 40°N, as the poleward
flow moves offshore. By summer, equatorward transport residuals next to the coast are strongest north of
40°N and extend around the Alaska Gyre. Since the long-term mean is removed, this represents a weakening
(or reversal) of the normal counterclockwise flow in the Alaska Gyre and a strong California Current,
especially north of 40°N. The seasonally weakened Alaska Gyre produces westward transport residuals on
its southern border in the interior at some locations between 50°–57°N. Poleward seasonal residuals begin
to appear next to the California coast south of 40°N even in summer, representing the ‘inshore countercurrent’ described by Lynn and Simpson (1987) and Strub and James (2000). By autumn, poleward residuals
begin to dominate, as the equatorward California Current moves offshore. The residual transports become
consistently poleward around the basin in winter. The magnitude of the largest transports in both summer
and winter are equivalent to velocities of order 5–15 cm s⫺1. These are the correct order for geostrophic
climatological currents in the California Current (Chelton, 1984; Lynn & Simpson, 1987).
One striking aspect of the seasonal circulation residuals in Fig. 3 is the degree to which the strongest
and most coherent transports are concentrated around the boundaries. This makes the whole system look
more like a single basin-scale gyre, rather than the two-gyre system seen in the mean flow of Fig. 1. The
primary connection to the interior zonal flow across the four north–south lines is in the Alaska Gyre
between 50°–57°N. Thus the northern portion of the Alaska Gyre appears to strengthen and weaken without
a corresponding change in the interior flow in the subtropical gyre. If the North Pacific Current simply
changed its seasonal location, moving north in winter to preferentially feed the Alaska Gyre and south in
summer to feed the California Current, the expected residual pattern in winter would consist of eastward
transport residuals north of westward residuals. The opposite pattern (westward north of eastward residuals)
would occur in summer. The interior north–south lines primarily show only the part of this pattern associated with the Alaska Gyre. Extending these north–south lines farther south does not change this result.
Five-year (1993–1997) seasonal means of surface forcing, as represented by the SLP fields, are shown
in Fig. 4. Note that the long-term mean is not removed from these fields. The forcing is consistent with
the circulation being primarily wind-driven. The Aleutian Low is strong in winter, representing the net
result of cyclonic storms that drive the counterclockwise Alaska Gyre and oppose the equatorward flow
in the northern half of the California Current System. In spring, the strengthened North Pacific High creates
equatorward wind stress in the California Current. By summer the pressure gradient causes eastward wind
stress in the northern Coastal Gulf of Alaska and equatorward wind stress along the entire eastern margin.
This reverses in the autumn north of central California, beginning the return to counterclockwise seasonal
wind forcing around the basin in winter. Note that the summer and winter patterns are more basinwide,
while the spring and autumn patterns are more like the mean double-gyre pattern expected in the ocean.
The region where the seasonal circulation residuals (Fig. 3) flow counter to the winds most consistently
is in the southern half of the California Current System, where poleward transport residuals are found next
to the coast in summer and autumn, despite strong equatorward winds. The tendency for poleward flow
next to the coast in summer off California is documented and referred to as the ‘Inshore Countercurrent’
(IC) by Lynn and Simpson (1987) and is also seen in higher resolution seasonal height fields from the
combined Geosat, ERS and TOPEX altimeters, presented by Strub and James (2000). The in situ data of
Lynn and Simpson (1987) indicate that it may be concentrated closer to the coast than depicted by Fig.
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Fig. 4.
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Seasonal SLP fields, based on 1993–1997 NCEP fields. Seasons are as in Fig. 3 and the 5-year mean is not removed.
3. A tendency for poleward flow in summer off central and southern California has also been noted in
previous studies using in situ data, as discussed further below.
Fig. 5 shows the first two Empirical Orthogonal Functions (EOFs) of the time series of transports over
the entire basin. These represent a total of 33% of the total variance and more compactly summarize the
principal features of the seasonal surface circulation seen in Fig. 3, with higher (monthly) temporal resolution. They also give some indication of the year-to-year variability in the seasonal cycles. The spatial
pattern of the first EOF (representing 23.5% of the variance) consists of poleward/counterclockwise boundary currents at the 1–2 locations next to the coast. There are no regions of strong transports in the interior
or into the margins. The time series for the first EOF forms a clear seasonal cycle, peaking in December–
February in each year. This represents the strongest period of the circulation in the Alaska Gyre and the
weak or reversed California Current. The negative summer minimum of the time series occurs between
June–August and corresponds to the strongest period of equatorward flow in the California Current and
the weakest period of poleward flow in the Alaska Gyre. Looking forward to the analysis of interannual
variability in Part 2 (Strub & James, 2002), we note that the summer minimum is weakest in mid-1997
and that the winter maximum is strong and longer in duration during the 1997–1998 El Niño.
In the California Current, the spatial pattern for the second EOF (representing 10% of the variance) is
equatorward for the transports next to the coast south of 37°N, poleward farther offshore between 30°–
40°N. In the Gulf of Alaska, the pattern strengthens the counterclockwise rotation in the northwest part
of the Alaska Gyre next to the coast and weakens it slightly farther offshore. The time series is usually
positive in spring and negative during late-summer and autumn. Thus the second EOF is uncorrelated with
the first, by virtue of a three month phase shift in the time series (in approximate quadrature). The second
EOF corresponds to the earlier (spring) beginning of equatorward flow next to the coast in the southern
half of the California Current System, as the winter poleward flow moves offshore. In late summer and
autumn, the pattern represents the early poleward flow next to the coast and the offshore migration of the
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Fig. 5. EOFs of transport residuals: (a) first EOF (23.5% of the variance); (b) second EOF (9.7% of the variance). In the spatial
pattern, transports with EOF weightings greater than 1 m2 s⫺2 are darkly shaded to highlight the main pattern of flow.
equatorward California Current in the same region. In the Gulf of Alaska, the second EOF continues the
westward flow next to the coast in the northwest during spring, prior to the reversal of the seasonal residuals
in summer.
To check whether the fairly specific grid of transports used above misses aspects of the seasonal flow,
more complete spatial maps derived from all available altimeter surface heights are used. Two-month
seasonal height residuals in Fig. 6, based on over seven years of altimeter data, cover the very large-scale
NE Pacific, from the equator to 61°N and from 80°–160°W. As stated above, the spatial mean is removed
from each two-month field, eliminating the expected large-scale steric rise and fall of hemispheric sea level
with seasonal heating and cooling. The remaining seasonal gradients in height are represented by the twocentimeter contour spacing. The magnitude of the transports are inversely proportional to the distance
between height contours at a given latitude, while arrows on the height contours indicate the direction in
which geostrophic currents flow.
The height residuals confirm that most of the strong, coherent variability north of 20°N is confined to
the boundaries. There is only the weakest indication that the strength of the zonal North Pacific Current
varies systematically on seasonal time scales. Strong variations in its strength or position would appear as
zonally oriented changes in these seasonal residuals. This type of variation is apparent south of 20°N, due
to seasonal changes in the Equatorial Current System, especially in the zonal NECC near 5°–10°N (strongest
in September–December and absent or reversed in March–June). No strong zonal patterns such as these
appear between 30°–55°N. However, if one averages the fields from November–April (autumn–winter),
the resulting field is smoother but similar to the November–December residual in Fig. 6f, with a slight
tendency for an eastward flow from 150°W to 130°W between 40°–50°N, due to a north–south gradient
in SSH caused by generally lower values to the north. Similarly, the average of the May–October (spring–
summer) fields produces slightly higher SSH values in the north, creating a weak tendency for a westward
spring–summer transport residual in this region (somewhat similar to the May–June residual in Fig. 6c,
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Fig. 6. Two-month seasonal SSH residuals based on Geosat, T/P, ERS-1 and ERS-2. Arrows indicate the direction of geostrophic
flow and geostrophic surface transports are inversely proportional to the contour spacing. The contour interval is 2 cm.
but smoother). Thus, there is a slight increase in surface transport in the North Pacific Current in winter,
as found in the drifter velocities by Bograd et al. (1999), but the seasonal changes are much weaker than
changes in the boundary currents.
The strongest pattern that relates the boundaries to the interior in California Current and Alaska Gyre
is the general offshore migration of the alongshore bands of height (and the associated bands of alongshore
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transports along height contours). Thus, in winter (January–February) there are high values of SSH next
to the coast, associated with poleward seasonal transport residuals extending from southern California to
the Alaska Gyre. At the same time, there is an offshore band of low SSH, associated with equatorward
transport on its western side. This band of low SSH is the expression of the band of low SSH found next
to the coast in July–August, which has moved offshore. The band of high SSH next to the coast in January–
February, in turn, moves offshore by July–August. At times there is a weak third band, farther offshore,
of the same sign as found next to the coast, representing the previous year’s coastal signal.
Previous studies of the California Current using altimeter data and surface drifters have also documented
the offshore seasonal movement of the eddy kinetic energy (EKE) generated in the vicinity of the seasonal
jet (Kelly et al., 1998). Strub and James (2000) use ‘intraseasonal’ variances of crosstrack geostrophic
velocity, calculated from the alongtrack gradients of altimeter SSH, to demonstrate this offshore movement
of EKE in the California Current. In Fig. 7, we present similar maps of ‘intraseasonal’ variance of crosstrack
geostrophic velocity in the Alaska Gyre during winter and summer. Since these variances only include one
component of velocity, they are only equivalent to the complete EKE field if the velocities are isotropic.
Still, they show the general pattern of the offshore movement. To create these maps, the mean crosstrack
velocity at each point along the T/P tracks is removed for each season. This leaves the variability caused by
changes within each season and by changes between years (representing both intraseasonal and interannual
Fig. 7. ‘Intraseasonal’ variances of geostrophic cross-track velocities, based on 5 years of T/P data. Only the winter (January–March)
and summer (July–September) periods are shown. Variances are formed by subtracting the 5-year seasonal mean from each observation. Thus, these variances include both intraseasonal and interannual variability.
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variability, although we call these ‘intraseasonal’ variances for simplicity). In the California Current, this
proxy for EKE is greatest in summer, surrounding the equatorward jet (Strub & James, 2000). In the Alaska
Gyre, Fig. 7 shows that the proxy EKE is greatest in winter in the 100–200 km next to the coast, where
the Alaska Current also has a winter maximum. In spring and summer, the variances move farther offshore
and diminish in magnitude, similar to the fate of the EKE in the California Current in autumn and winter.
This is clearest in the north and northeast regions of the Alaska Gyre, less apparent in the Alaska Stream
(where the data is often flagged as suspect, due to the wide shelf). The net result is that EKE is greatest near
the strongest flow in both the Alaska Gyre and the California Current, suggesting dynamical instabilities as
a possible source of the EKE. In both cases the EKE diminishes and moves offshore, similar to the movement of the SSH signals.
Besides the offshore migration of SSH, alongshore currents and EKE in both the Alaska Gyre and
California Current, the larger picture provided by Fig. 6 also suggests a connection between height and
transport residuals in the southern part of the California Current and those in the eastern tropical Pacific
farther south. The seasonal height residuals in January–February (Fig. 6a) indicate low sea level and weaker
poleward flow (or actual equatorward flow) along the Mexican mainland and Central America. These extend
as far north as southern Baja California in January–February. The equatorward residuals further strengthen
off Central America and Mexico in March–April. This winter and early spring period is the time when
the normal eastward flow in the NECC and the northward flow along Central America (the Costa Rica
Current) weaken or reverse in ship-drift currents shown by Fiedler (1992, 2002). By March–April, the
residual flow is strongly equatorward from Central America to Baja California and weakly equatorward
as far north as northern California. Equatorward residuals off Central America next to the coast reverse
in May–June and by July–August strong poleward residuals occur as far north as the tip of Baja California,
with weaker poleward residuals as far north as central California. The strongest poleward residuals move
to Baja California in September–October, then extend around the Gulf of Alaska in November–December,
while they weaken off Mexico. Thus, the observed earlier poleward (and equatorward) flow in the southern
part of the California Current appears to be connected to similar, stronger flow off Mexico and Central
America.
The poleward progression of the seasonal signals is most clearly defined by the timing of the annual
maximum SSH at locations 50–100 km from the coast, as a function of latitude (reconstructed from annual
and semiannual harmonics). The combination of these two harmonics identify the steady progression of
peak SSH (not the first arrival of the SSH signal, but the peak) from Central America in June to mainland
Mexico (July–August), Baja California (September–November), southern California (December) and then
around the Alaska Gyre in January–March. A similar analysis of the month of minimum SSH also starts
off Central America (in February–March) and progresses to Oregon–British Columbia by July–September.
This pattern of northward seasonal progression of coastal SSH from Central America to the Alaska Gyre
also appears in the seasonal cycles of coastal tide gauge records presented by Enfield and Allen (1980).
4. Summary and discussion
With respect to the four objectives stated at the end of the Introduction, we summarize the most certain
of our results, then discuss them further.
1. The maximum surface currents in the Alaska Current and Alaska Stream occur in winter (December–
February), as demonstrated in Figs. 3, 5 and 6. This is in phase (approximately) with large-scale surface
forcing (Fig. 4).
2. The primary seasonal variability in surface circulation and transport in the NE Pacific north of 20°N is
in the boundary currents stretching from mainland Mexico around the Alaska Gyre (Figs. 3 and 6).
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Seasonal residuals of surface transport are equatorward in summer around the Alaska Gyre and in the
California Current, except for an inshore poleward countercurrent south of approximately 40°N after
July. In winter, the seasonal residuals flow poleward from Baja California around the Alaska Gyre.
3. Seasonal residuals in transport between the interior of the NE Pacific (the North Pacific Current) and
the boundaries of the subarctic and subtropical gyres are very weak compared to the residuals along
the boundaries (Fig. 6). The most consistent pattern around the basin is the seasonal offshore movement
of bands of high and low SSH, with corresponding cells of alongshore geostrophic current and EKE.
The offshore scale of these cells is approximately 200–500 km.
4. Seasonal transport residuals along Central America and the Mexican mainland lead those in the California Current off central California by approximately six months (Fig. 6). This creates an out-of-phase
relationship between SSH in the low latitudes and SSH in the mid- and high-latitudes. Transport residuals
are equatorward off Central America in January–April, while residuals are mostly poleward in the California Current and Alaska Gyre. Residuals are poleward off Central America in May–October, while
residuals in the California Current are mostly equatorward, except in the inshore countercurrent.
4.1. Alaska Gyre
Our finding of a winter seasonal maximum in surface transport in the Alaska Gyre confirms the tentative
conclusion of Bhaskaran et al. (1993) and is consistent with the general finding of an October–March
maximum in the Alaska Stream from surface drifters (Bograd et al., 1999). It is somewhat at odds with
the March or later date of Royer (1981b) and the autumn date of Kelly et al. (1993). The transport fields
in Figs. 3 and 5 are based on five years of consistently processed data from the high quality TOPEX
altimeter. The height fields in Fig. 6 are based on over seven years of data from four altimeters of varying
quality and differing orbits, but still with more regular sampling than the hydrography available to Royer
(1981b) and a much longer record than available to Bhaskaran et al. (1993) and Kelly et al. (1993).
Bhaskaran et al. (1993) based their analysis on an EOF decomposition of nearly four years of Geosat
data, which found an annual cycle in one mode. This corresponded to a low SSH cell north of 55°N,
lowest in winter, similar to the cell we find in two-month means from seven years of data in Fig. 6. Kelly
et al. (1993) based their seasonal cycle on the fit of annual and semiannual harmonics to 2.5 years of
Geosat data. Their SSH field for January consists primarily of low residuals, sloping down to the north,
which may partly result from seasonal steric cooling. Examination of their fields (their Fig. 9), however,
also shows that they were forced to eliminate data within approximately 2°–3° of the coast, precisely where
the combined altimeter data find the high SSH residuals associated with the strengthened gyre in January–
February. Likewise their fields of high SSH in July and October miss the low SSH band next to the coast
to the north. It seems likely that the analysis of Kelly et al. (1993) found only the southern half of the
winter low and summer high cells, offshore of the strong boundary signals in the NE Alaska Gyre in
Fig. 6. Thus, the results of both analyses of Geosat data are consistent with our results, with somewhat
different interpretations.
With respect to the nature of the Alaska Gyre seasonal residuals, in Fig. 6 the offshore migration of the
summer coastal low and winter coastal high produces an elongated NW–SE gyre, alternating in transport
direction between winter and summer. This is in contrast to the location of the center of the climatological
gyre (Fig. 1), which is at 54°N, 150°W. Thus, the seasonal variability of the Alaska Gyre is not in the
overall gyre, but in the structure along the boundary and in the interior in the northeastern portion of the
basin. To place these seasonal residuals into perspective, it is useful to combine them with the mean
dynamic height field shown in Fig. 1. In Fig. 8 we do this for the January–February and July–August
periods from Fig. 6. This makes it clear that the effect of the seasonal residuals in the NE corner is to
expand the central low of the Gyre farther into the NE corner in winter and strengthen the gradient next
to the coast. In summer, the central low contracts to its climatological mean position and gradients next
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Fig. 8. Two-month seasonal height residuals (from Fig. 6) added to dynamic height climatology (from Fig. 1). Only winter and
summer are shown. The contour interval is 2 cm.
to the coast become weaker than in the climatology. The location and strength of the center of the gyre,
however, is not greatly affected (Fig. 6), although it may be difficult to determine that by looking at the
combined fields of residual plus long-term mean (Fig. 8). We have chosen to present the seasonal residual
fields alone in previous figures, since this isolates the temporally varying signal, which is all that the
altimeter can resolve.
4.2. Connections around the boundaries of the subarctic and subtropical gyres
In the western and northern Alaska Gyre, it is doubtful that the currents actually reverse in summer
along the boundaries, despite the appearance of eastward coastal flow in Fig. 8. The climatology is too
smooth to portray the actual strength of the mean boundary currents in those regions and so is overpowered
by the altimeter seasonal residuals, which are not as smooth. Royer (1998) cites measurements in the
Alaska Coastal Current at the apex of the Gyre, showing that velocities drop from 40–50 cm s⫺1 (or more)
in September–March to 10–20 cm s⫺1 in April–June but do not reverse (Johnson et al., 1988). More recent
measurements of the ACC farther west in April–September 1991 also show no seasonal reversal (Stabeno
et al., 1995).
In the eastern Alaska Gyre, there is a greater likelihood of seasonal reversals, offshore of a narrow,
poleward, buoyancy-driven jet. This poleward jet has been measured at the northern end of Vancouver
Island (Hickey et al., 1991) but its behavior north of Vancouver Island is not well known. Hickey et al.
(1991) also show a broader equatorward flow of over 15 cm s⫺1 offshore of the poleward jet. The altimeter
transport values and height differences correspond to velocities of 5–10 cm s⫺1 in the summer north of
Vancouver Island, similar to those measured by Hickey et al. (1991). In addition, Thomson and Gower
(1998) and Thomson, Hickey, and LeBlond (1989) depict summer currents over the slope as equatorward
north of and along Vancouver Island. Thus, we postulate that the surface velocities are equatorward over
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the slope (perhaps offshore of a narrow poleward jet) in summer south of 53°–55°N, similar to the field
obtained by combining the summer altimeter residuals with the climatological dynamic heights (Fig. 8).
In winter, poleward velocities have been found from 35°N to the British Columbia coast (Chelton,
Bratkovich, Bernstein, & Kosro, 1988; Hickey et al., 1991; Largier, Magnell, & Winant, 1993; Royer,
1998; Strub et al., 1987; Wickham et al., 1987). Because of these measurements, we believe that the
combined height fields in Fig. 8a are qualitatively correct in winter in the Alaska Gyre and the northern
part of the California Current. Thus, the boundaries of the eastern Alaska Gyre and northern California
Current are directly connected in both summer and winter.
4.3. Connections to the North Pacific Current
Here we consider the question: what is the likely origin of seasonal changes in the water properties
within the core of the California Current off central and southern California, as measured by CalCOFI
surveys? The analysis shows only slight increases and decreases in strength of the North Pacific Current
in winter and summer, making it clear that the seasonal changes in the boundary currents are not driven
by seasonal changes in the interior ocean. Still, even an absolutely constant eastward flow in the North
Pacific Current will continue to bring water from the interior to the boundaries. How will that interior
water be advected by the seasonal currents?
Figs. 6a and 8a suggest that water transported into the northern end of the California Current in winter
will turn south, offshore of the poleward Davidson Current or will join the poleward flow into the Alaska
Gyre. The water that turns south stays far offshore until it approaches the coast in the Southern California
Bight and along Baja California. In spring (Fig. 6b), equatorward flow begins next to the coast, still separated from the North Pacific Current by the poleward remnants of the Davidson Current north of the Bight.
By summer (Figs. 6c, 6d and 8b), equatorward flow off central California is more directly connected along
the boundary to the Pacific Northwest and the British Columbia Coast than to the interior NE Pacific. This
helps explain why drifters placed in the North Pacific Current turn south far offshore and rarely join the
core of the California Current (McNally, 1981), while drifters released in the coastal current off Oregon
in spring flow directly through the core of the California Current off central California in summer (Barth &
Smith, 1998; Barth, Pierce, & Smith, 2000). In autumn, water from the North Pacific Current that has
traveled south offshore of the seasonal jet may flow eastward into the northward coastal current found off
Baja California and the Southern California Bight.
From these relationships, we hypothesize that seasonal changes in water mass characteristics in the upper
core of the California Current derive primarily from advection along the boundaries and that the cool, fresh
water found in the seasonal jet in summer off central California, comes from the Pacific Northwest and
the British Columbia coast, rather than from the interior North Pacific. In the Southern California Bight
and farther south, water from the interior North Pacific flows onshore to join the mix, creating the complex
interleaving of water masses noted by numerous studies (Lynn & Simpson, 1987). The CalCOFI climatology shows that the fresh water core of the California Current becomes a subsurface feature south of the
Bight and that this subsurface salinity minimum is fresher in July–October than in January and April (Lynn,
Bliss, & Eber, 1982). We interpret the change from surface to subsurface minimum to reflect the onshore
movement of interior water in the Bight, which overlays the fresher, cooler water from the north. The
freshening of the subsurface minimum in July–October reflects the net result of increased flow from the
north in the seasonal jet in spring and summer.
4.4. Offshore ‘propagation’ of the seasonal height and transport signals
The offshore migration or propagation of SSH, isotherm depths, velocity, and eddy kinetic energy in
the California Current has been documented in a number of studies, most recently by Kelly et al. (1998)
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and Strub and James (2000). Two primary mechanisms have been hypothesized to be responsible for this
westward movement of dynamic features: propagation as first mode, annual baroclinic Rossby waves
(White, Tai, & DiMento, 1990); or a response to direct surface forcing by the large-scale curl of the wind
stress, which also moves westward on a seasonal time scale (Kelly et al., 1993).
Although White et al. (1990) found the westward propagation of SSH in one year of Geosat altimeter
data to match the dispersion relationship expected for Rossby waves, sampling aliases of tidal errors in
the Geosat data, along with the short data record, make that conclusion tentative. Likewise, although Kelly
et al. (1993) found that westward movement of ECMWF wind stress curl fields were able to explain the
westward movement of features in two years of Geosat data (using a simple dynamic model), Kelly et al.
(1998) could not find a similar relation using T/P data and winds from the ERS scatterometer or ECMWF
model. However, since Kelly et al. (1998) only had one year of overlapping wind and altimeter data, their
failure to create westward propagation with the wind and model does not eliminate direct wind forcing as
a possible mechanism. The offshore migration of low and high sea level patterns and EKE in the Alaska
Gyre adds a new wrinkle to the problem, since it occurs north of the latitude where we expect annual
Rossby waves to propagate in the North Pacific. Thus, a definitive explanation of this phenomenon remains
to be found and it seems likely that a mix of processes is responsible for the observed behavior.
4.5. Connections to the equatorial current systems along the boundaries
The poleward flow in summer off central California is directly counter to the wind and has been an
enigma for some time. The summer poleward flow within the Southern California Bight is well known,
forming one branch of the Southern California Eddy (Tsuchiya, 1980). It is sometimes called the Southern
California Countercurrent (Hickey, 1998). This poleward coastal current is less of an enigma, since it is
consistent with forcing by the positive wind stress curl in the sheltered, semi-enclosed Bight.
Poleward flow over the shelf and slope north of the Southern California Bight in summer, directly
opposing the wind, has been observed by several studies (Chelton et al., 1988; Strub et al., 1987; Wickham
et al., 1987). Most attempts to explain this flow have involved local forcing by the wind and wind stress
curl (Oey, 1999; Wang, 1997), pressure gradients originating in the Santa Barbara Channel (Harms &
Winant, 1998), rectification of alternating currents caused by fluctuating winds (Samelson & Allen, 1987),
or eddy-topography interactions (Holloway, 1987). The results presented here suggests that there may also
be a more remote component to the observed high sea levels and poleward flow off southern and central
California in summer. The altimeter fields show that higher sea levels and poleward residuals appear off
Central America in May–June, then strengthen and spread north in July–August. This follows the northward
movement of the ITCZ, which has a region of wind-driven divergence on its northern side that re-establishes
(or strengthens) a trough between the NECC and NEC east of 120°W. At this time, the NECC connects
to the cyclonic flow around the southern side of the Costa Rica Dome and feeds into the poleward Costa
Rica Current. Suggestions for the dynamics governing flow around the Costa Rica Dome include: geostrophic adjustment to the cyclonic flow imposed when the NECC reaches the coast and is forced to the
north (Wyrtki, 1964); or a response to the local wind stress curl (Hofmann, Busalacchi, & O’Brien, 1981).
Irrespective of the dynamics, the question here is whether it is possible for these annual signals to propagate
to the north and oppose local forcing?
The dynamics responsible for slow poleward propagation of low-frequency height residuals along the
eastern Pacific margin have been investigated by Grimshaw and Allen (1988) and Clarke and Van Gorder
(1994). Grimshaw and Allen (their Fig. 7) present evidence that annual signals traveling slower than 100–
150 km day⫺1 could propagate from central Mexico (17°N) to central California in a coastal trapped mode,
rather than radiating offshore as Rossby waves. Although their model is highly idealized, it uses realistic
stratification and coastal morphology, without bottom topography. Estimates of the seasonal lag in high
sea levels between central Mexico and central California from our data (Fig. 6) and from tide gauge data
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(Enfield & Allen, 1980, their Fig. 3) yield approximately 2–4 months, resulting in phase propagation speeds
below 50 km day⫺1. At these slow speeds, the model of Grimshaw and Allen suggests that poleward
propagation should be considered as a possible mechanism to explain the observations.
This question of whether the movement of the signal in the poleward direction is propagation of a
distantly forced feature, or a purely local response to winds that also develop from south to north, has
consequences for regional modelers. If the poleward currents opposing the winds in the southern region
of the California Current are generated locally, then isolated regional models, such as by Haidvogel,
Beckmann, and Hedstrom (1991), Batteen and Vance (1998) and others, should be able to reproduce those
currents. To date, this has not been the case. Batteen and Vance (1998) note that their model, which has
a southern border at 37°N, does not produce a poleward current in the southern part of the domain. They
suggest that the Davidson Current starts off Point Conception (35°N) and so can not be reproduced within
the closed boundaries of the model. We suggest that the tendency for poleward currents begins much
farther south, earlier in the year, requiring much larger domains and/or nested approaches to capture the
true large-scale variability revealed by the altimeter in this analysis.
Acknowledgements
Support for this work was provided by JPL grants 958128 (TOPEX) and 1206714 (Jason-1), NAG56604 (SeaWiFS), and NSF/NASA grants 9711344-OCE and 0000900-OCE (US GLOBEC–NE Pacific
Program). Altimeter data were provided by JPL and GSFC. Tide gauge data were provided by NOAA/NOS
and the University of Hawaii. Sea level pressure fields are from the NCEP/NCAR Reanalysis project. SLP
times series used to adjust the tide gauge data are from NCEP and the NOAA/PFEL laboratory, courtesy
of Frank Schwing and Jerrold Norton. The quality of this manuscript benefited from comments by anonymous reviewers and revisions suggested by the editor, Dr M. Angel. This is contribution number 173 of
the U.S. GLOBEC program, which is jointly funded by the National Science Foundation and the National
Oceanic and Atmospheric Administration.
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