- American Meteorological Society

FEBRUARY 2011
KASHINO ET AL.
287
Observed Ocean Variability in the Mindanao Dome Region
YUJI KASHINO, AKIO ISHIDA, AND SHIGEKI HOSODA
Research Institute for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokosuka, Japan
(Manuscript received 31 July 2009, in final form 6 October 2010)
ABSTRACT
Ocean variations at semiannual, annual, and interannual time scales in the Mindanao Dome (MD) region
of the southern Philippine Sea were examined using data derived from underwater sensors on Triangle
Trans-Ocean Buoy Network (TRITON) buoys at 88N, 1378E; 58N, 1378E; and 88N, 1308E. Annual signal
dominated above 300-m depth in the MD region. At 58N, 1378E, saline water exceeding 35 psu was observed
at 100–200-m depth from boreal winter to spring, seemingly associated with the meridional migration of the
North Equatorial Countercurrent during these seasons. Thermocline ascent, probably related to the MD, was
also observed from boreal winter to spring. An important mechanism of the annual variation of the MD at 58N
seems to be the annual variability of local wind, as mentioned in past studies. However, annual variability at
88N seems to be due to Rossby waves originating west of 1508W rather than to local wind effects. Semiannual
variation was also observed, with its amplitude reaching 40%–70% of the annual signal. With regard to
interannual variability, ocean variation on the time scale of the El Niño–Southern Oscillation (ENSO) was
seen; upper heat content (above 300-m depth) in the Mindanao Dome region decreased during the 2002–03
and 2006–07 El Niño periods and increased between those periods. Increasing upper heat content in this
region after 2005 was probably associated with large negative anomalies of Ekman pumping (downwelling)
that appeared from 2005 to 2006 east of 1508E and north of 58N.
1. Introduction
The southern Philippine Sea (08–158N, 1258–1408E) is
an interesting region not only for oceanography but also
for climate research for a number of reasons. First, this
region encompasses the western part of a warm water
pool that has large impacts on climate variability involving
the El Niño–Southern Oscillation (ENSO). Rasmusson
and Carpenter (1982) found that sea surface temperature (SST) in this region was related to the ENSO, and
Hendon (2003) showed that SST variability was closely
correlated with the Niño-3.4 index; that is, SST in this
region drops in association with the warm phase of
El Niño. Although the change in SST is not larger than
that in the central and eastern equatorial Pacific, it is
large enough to generate wind variability in the western equatorial Pacific, a region that influences ENSO
variability (Wang et al. 1999). Furthermore, because
SST changes in the Indo-Pacific region alter regional
Corresponding author address: Dr. Yuji Kashino, Research Institute for Global Change, Japan Agency for Marine-Earth Science
and Technology, 2-15 Natsushima, Yokosuka 237-0061, Japan.
E-mail: kashinoy@jamstec.go.jp
DOI: 10.1175/2010JPO4329.1
Ó 2011 American Meteorological Society
atmospheric convection, rainfall over the Maritime Continent is also affected (Hendon 2003).
Next, this region is not only a ‘‘water mass crossroads’’
(Fine et al. 1994) but also a crossroads of ocean currents,
which flow through it in complicated current patterns
(Fig. 1). Low-latitude western boundary currents advect
several water masses that flow into this region from both
hemispheres. Some of these waters flow into the Indonesian seas as the Indonesian Throughflow, whereas others
flow eastward as the North Equatorial Countercurrent
(NECC). These currents probably play important roles
in the heat budget of the warm water pool and in global
ocean circulation (Lukas et al. 1996).
In the Northern Hemisphere (NH), the westwardflowing North Equatorial Current (NEC) bifurcates into
two western boundary currents, the Kuroshio and the
Mindanao Current (MC), near the Philippines. The bifurcation latitude of the NEC is thought to be related to
ENSO phenomena because of the high correlation between the bifurcation latitude and the Southern Oscillation index (Qiu and Lukas 1996; Kim et al. 2004). The MC,
one of the branches, flows southward along the Philippine
coast and supplies a large part of the source waters of
the Indonesian Throughflow (e.g., Gordon and Fine 1996).
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FIG. 1. Schematic of the major ocean surface currents and
structures in the western equatorial Pacific imposed on a map of
TRITON buoys denoted by solid and open stars. The Kuroshio,
NEC, NECC, MC, NGCC, South Equatorial Current (SEC),
Halmahera eddy (HE), and ME are involved in this map. The MD
is denoted by a dashed ellipse. We mainly analyzed data from
TRITON buoys T10 (88N, 1378E), T11 (58N, 1378E), and T14 (88N,
1308E), denoted by solid stars.
Because the NEC and MC transport waters that originated in the Northern Hemisphere to the equatorial
Pacific, they are potentially important factors in interdecadal climate variability (Gu and Philander 1997).
In the Southern Hemisphere (SH), the New Guinea
Coastal Current (NGCC) and New Guinea Coastal Undercurrent (NGCUC) flow along the New Guinea coast
across the equator from the South Pacific, although the
NGCC at the surface reverses direction during winter
(Kuroda 2000). Since these currents advect waters that
originated in the South Pacific to the southernmost
Philippine Sea, the South Pacific waters reach the east
coast of Halmahera Island and enter the Indonesian
seas as part of the Indonesian Throughflow (Ffield and
Gordon 1992; Kashino et al. 1996). Some of the South
Pacific waters flow to the east as the dominant source
water of the Equatorial Undercurrent (Tsuchiya et al.
1989). The South Pacific tropical water (SPTW) around
the subsurface layer (100–300 m) is particularly notable;
the potential vorticity of this water is lower than that of
North Pacific waters around the same depth. Therefore,
these waters create a potential vorticity front at the axis of
the NECC, where they meet (Gouriou and Toole 1993).
The Mindanao Dome (MD) simulated by Masumoto
and Yamagata (1991), Tozuka et al. (2002), and Suzuki
et al. (2005) is another interesting oceanic phenomenon
east of the Philippines. The MD is sometimes confused
with the Mindanao Eddy (ME) (Tozuka et al. 2002).
Their numerical studies showed that the MD develops
in boreal (boreal is omitted hereafter) winter due to
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positive wind stress curl in the Philippine Sea and decays
due to warm-anomaly propagation from the east. The
MD phenomenon is thought to be important for heat
content variability in the western tropical Pacific (Tozuka
et al. 2002).
Thus, there are many interesting oceanographic research themes to address in the southern Philippine Sea,
including regional/global climate variability from intraseasonal to decadal time scales involving ENSO, ocean
structures ranging from ocean eddies to global ocean circulation, and ocean dynamics. However, research progress
has been hindered by limited ocean observation data
compared with other equatorial regions. Time series data
are particularly limited. Therefore, even though numerical and theoretical studies have derived some interesting
results, most findings have not been fully confirmed; for
example, no observational study of the MD has been
published, with the exception of the recent study by
Kashino et al. (2009).
A number of studies, including the United States–
People’s Republic of China Cooperative Studies of Air–
Sea Interaction in the Tropical Western Pacific (US-PRC)
project (e.g., Toole et al. 1990), the Western Equatorial
Pacific Ocean Climate Study (WEPOCS) (e.g., Lukas
et al. 1991), and the World Ocean Circulation Experiment (WOCE) (e.g., Kashino et al. 1996) have examined
ocean currents and water masses in the southern Philippine Sea using hydrographic observations. In addition,
analyses of historical data have helped clarify the mean
states of currents and water masses in this region (e.g.,
Qu et al. 1998, 1999). Because of these works, a zeroorder description of the ocean in this region seems to
have been completed up to the end of the twentieth
century (Lukas et al. 1996). However, a full description
of ocean variability has not been achieved because of
the lack of time series data, as mentioned above. Ocean
variability in this region has been mainly studied using
numerical methods, with the exception of some studies
using sea level data and other types of observations (e.g.,
Lukas 1988; Mitchum and Lukas 1990; Kessler 1990).
The Japan Agency for Marine-Earth Science and
Technology (JAMSTEC, formerly the Japan Marine
Science and Technology Center) has been conducting
observations using the R/Vs Kaiyo, Mirai, and Yokosuka
in the western equatorial Pacific and eastern Indian
Ocean under the Tropical Ocean Climate Study (TOCS)
project since 1993. To address the lack of time series
data, several moorings with current meters and acoustic
Doppler current profilers (ADCPs) have also been deployed at 4819N, 1278319E; 38119N, 1288279E (Kashino
et al. 1999); 68509N, 1268439E (Kashino et al. 2005);
28N, 1388E; 08, 1388E (Kashino et al. 2007); 08, 1428E;
28309S, 1428E (Kuroda 2000; Ueki et al. 2003); 08, 1478E
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KASHINO ET AL.
(Matsuura 2002); and 08, 1568E. These current observations have revealed some interesting phenomena, including intraseasonal variability (50 days) at the Pacific
entrance of the Indonesian Throughflow (Kashino et al.
1999), current reversal (southeastward flow) of the NGCC
during winter (Kuroda 2000), change of the NGCC/
NGCUC during the 1997–98 El Niño (Ueki et al. 2003),
and current speed increase of the MC during the onset of
the 2002–03 El Niño (Kashino et al. 2005).
JAMSTEC has also been deploying Triangle TransOcean Buoy Network (TRITON) surface buoys (Kuroda
and Amitani 2001) in the western equatorial Pacific and
eastern Indian Ocean to measure meteorological and
oceanographic parameters (Fig. 1). The underwater conductivity and temperature sensors, which provide data
not only on temperature variations but also on salinity
variations in the ocean, are installed in TRITON buoys.
For the southern Philippine Sea, TRITON buoy data
have been used to examine ocean variability north of
New Guinea (Kashino et al. 2007). However, oceanic variability has not been previously described for the MD region. TRITON buoys have been deployed in the MD
region (88N, 1308E; 88N, 1378E; and 58N, 1378E) since
2001, and the collected data can be used to describe ocean
variability in this region.
Furthermore, in 2000, the international Argo Project
was launched with the goal of deploying more than 3000
profiling floats to observe temperature and salinity profiles from the sea surface to 2000 dbar throughout the
World Ocean (Argo Science Team 1998, 2001). The number of deployed floats has been increasing each year,
making it possible for researchers to investigate basinscale ocean variability. Although the temporal resolution
of Argo data is not better than that of data from moorings
such as the TRITON buoys, data from Argo floats in the
western equatorial Pacific are useful for interpreting the
variability seen in the TRITON buoy data.
This paper uses TRITON buoy data to describe ocean
variability having semiannual and longer signals, with a
focus on the MD region of the southern Philippine Sea.
The following section presents the data used in this study.
Section 3 describes oceanic variability identified from
TRITON buoy data, while section 4 presents the variability found using TRITON buoy data combined with
Argo float data and surface wind data. A summary is
provided in section 5.
2. Data sources
As described in section 1, we used data collected by
TRITON buoys in the western tropical Pacific. TRITON
buoys (Fig. 2) have been deployed in this region since
1998. The collected data are used for monitoring and
289
FIG. 2. Mooring configuration of a TRITON buoy.
understanding atmospheric and oceanic conditions of
the equatorial Pacific and Indian Oceans, in conjunction with data obtained by Tropical Atmosphere Ocean
(TAO) buoys of the National Oceanic and Atmospheric
Administration (Kuroda and Amitani 2001). For these
purposes, TRITON buoys are surface buoys equipped
with both meteorological and underwater sensors. The
TRITON buoys have also been designed considering the
serious damage sometimes inflicted on monitoring equipment by vandalism activities in the equatorial region.
Additionally, both temperature and conductivity sensors
are installed on the buoys to take salinity measurements.
Therefore, TRITON data can be used to examine the
variability of the ocean above the intermediate layer
(.750 m).
In the western equatorial Pacific, 16 TRITON buoys
have been deployed west of 1568E (Fig. 1), although observations by buoy T15 (58N, 1308E) ended in November
2002. Of these buoys, we mainly used data from three
buoys, located at 88N, 1378E (T10); 58N, 1378E (T11); and
88N, 1308E (T14) in the MD region, with supplementary
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data from buoys at 88N, 1568E (T01); 58N, 1568E (T02);
and 58N, 1478E (T07). These real-time data are available
from the TRITON Web site (available online at http://
www.jamstec.go.jp/jamstec/TRITON/real_time/top.html).
Note that data from T10 and T11 are also available at
the TAO Web site (available online at http://www.pmel.
noaa.gov/tao/) because these two buoys are part of the
TAO/TRITON buoy array. In this study, we analyzed
data from 28 September 2001 to 12 January 2008 (T10),
from 29 September 2001 to 11 January 2008 (T11), and
from 13 August 2002 to 18 January 2008 (T14). These
periods included the 2002–03 El Niño, 2006–07 El Niño,
and 2007–08 La Niña events.
We analyzed data from conductivity–temperature–
depth (CTD) sensors at 300- and 750-m depths and from
conductivity–temperature (CT) sensors at depth 1.5, 25,
50, 75, 100, 125, 150, 200, 250, and 500 m along the
mooring lines of the TRITON buoys for this study. Data
from one extra CT sensor on buoy T11 at 175-m depth
were also used. Original temperature, conductivity, and
pressure values were stored in the sensors every 10 min.
The original data were corrected based on the results of
sensor calibrations conducted by the Mutsu Institute of
Oceanography of JAMSTEC (Ando et al. 2005) after
recovery of the moorings. After the sensor calibrations
and data correction shown by Ando et al. (2005), the
mean and standard deviation of the difference between
onboard CTD salinity and TRITON CT-sensor salinity
measurements were 20.001 and 0.033 psu, respectively.
We then compiled daily datasets from the corrected
10-min-interval data.
The buoys have been maintained using the R/Vs Kaiyo,
Mirai, and Yokosuka under the TOCS and TRITON
projects of JAMSTEC. During these cruises, CTD observations were conducted near the buoy locations to
collect data for checking sensor performance.
We used Argo data to discuss the time series appearing in the TRITON buoy data. Because Argo floats drift
freely and do not remain in the same location, we used
monthly optimal analysis data combined with data from
the TRITON buoys and some onboard CTD observations.
For this study, data from January 2002 to December 2007
were examined. These optimal analysis data had a horizontal resolution of 18 3 18 and 25 levels in the vertical
from 10 to 2000 dbar. The optimal analysis method has
been described by Hosoda et al. (2006, 2008), and the data
are available at the JAMSTEC Argo Web site (available
online at http://www.jamstec.go.jp/ARGO/argo_web/
argo/index_e.html).
We also used surface wind data provided by the European Centre for Medium-Range Weather Forecasts
for discussion of the ocean observations. These wind data
are from one of the sets of operational surface analysis
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data archived by the ECMWF. To calculate wind stress,
we used a drag coefficient of 1.5 3 1023 and an air density
of 1.18 kg m23.
3. Results
a. Mean state
Before describing the observed time series, we describe the mean states at the sites of TRITON buoys
T10 (88N, 1378E), T11 (58N, 1378E), and T14 (88N, 1308E)
to allow for better understanding of the time series results (Fig. 3). Note that the vertical profiles are plotted
using averaged values at each sensor depth, and the potential temperature–salinity relationships are plotted using isopycnal-coordinate data averaged in 0.1 su bins.
Thermocline depths at T10 and T14 (around 140–150 m)
were shallower than at T11 (around 200 m) because the
latitude of T10 and T14 (88N) is close to the boundary
between the NEC and NECC and the thermocline ridge
is located around this latitude (Kessler 1990; Qiu and
Joyce 1992). The shallowest thermocline depth in this
region occurred at the westernmost site, T14 (Tozuka
et al. 2002). However, the SST and mixed layer depth
(around 30 m) did not largely differ at these stations.
Some salinity maxima and minima were found in salinity
profiles and potential temperature–salinity relationships.
For example, shallow salinity maxima were seen between
24.0 and 24.5 su (100 and 150 m). The salinity maximum
at T11 occurred in a deeper layer than at T10 and T14.
Salinity minima were observed in the intermediate layer
between 26.3 and 26.5 su, and the density of the salinity
minimum at T14 was larger than the densities of the
salinity minima at T10 and T11.
Water masses in this region have been discussed by
Fine et al. (1994), Bingham and Lukas (1994, 1995),
Kashino et al. (1996), Kaneko et al. (1998, 2001), Qu et al.
(1999), and others. These studies reported two salinity
maxima, North Pacific Tropical Water (NPTW) at 24.0
su and SPTW at 24.8 su, and two salinity minima, North
Pacific Intermediate Water (NPIW) at 26.5 su and Antarctic Intermediate Water (AAIW) at 27.2 su. On the
basis of those findings, the shallow salinity maxima
observed at the TRITON sites seem to represent the
NPTW, SPTW, or their mixture. The location and density indicate that the salinity maximum observed at T10
and T14 was the NPTW, which originates from the
subtropical Pacific near Hawaii and is advected by
the westward-flowing NEC to the Philippine Sea
(Tsuchiya 1968). Maximum salinity of the NPTW at
T10 was higher than that at T14, probably owing to
diffusion during advection of the salinity maximum
from T14 to T10.
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KASHINO ET AL.
291
FIG. 3. Mean vertical profiles of (a) temperature and (b) salinity and (c) potential temperature–salinity relationships at T10 (thin line), T11 (dashed line), and T14 (thick line). Arrows on the right axes of (a),(b) indicate the sensor
depths.
At T11, the salinity maximum was higher and had
greater density than the maxima at T10 and T14. The
NECC flows near this site (Kashino et al. 2007). Because
the northern boundary of the SPTW corresponds with
the axis of the NECC (Gouriou and Toole 1993), the
high salinity observed at T11 can be attributed to the
appearance of the SPTW, which has higher salinity than
the NPTW.
The salinity minimum at 26.5 su at T14 represents the
NPIW, as evidenced by the good agreement of the curve
of the potential temperature–salinity relationship around
that density with that described by Bingham and Lukas
(1994), who examined water masses near the Mindanao
coast (T14 is the closest TRITON buoy to the Mindanao
coast). However, the intermediate salinity minimum at
26.3 su at T10 and T11 seems to differ in density from
the NPIW. As described by Bingham and Lukas (1995),
there is another intermediate water in this region: the
Northern Pacific Tropical Intermediate Water (NPTIW),
with a salinity maximum and an oxygen minimum around
26.8 su. Therefore, the salinity minimum at T10 and
T11 might have been the NPIW that had been eroded
by an intrusion of NPTIW, as in the case of the tropical
salinity minimum (TSM), which seems to be created by
an intrusion of high-salinity water into the mixed waters
of the NPIW and shallow salinity minimum (Yuan and
Talley 1992). Kaneko et al. (2001) also called this salinity
minimum the TSM.
b. Variability
Depth–time plots of 120-day low-pass filtered time series of temperature, salinity, and potential density (Fig. 4)
show that the seasonal signal dominated at all sites. At
T10 and T14, the thermocline was deep from summer to
autumn and shallow from winter to spring. Harmonic
analysis results for the 208C isotherm depth indicate that
the thermocline at T10 (T14) was shallowest in February
(March). The thermocline ascended from winter to spring,
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FIG. 4. Depth–time plots of 120-day low-pass filtered time series
of (a) temperature, (b) salinity, and (c) potential density at T14;
(d) temperature, (e) salinity, and (f) potential density at T10; and
(g) temperature, (h) salinity, and (i) potential density at T11. Contour intervals are 18C, 0.1 psu, and 0.5 su for temperature, salinity,
and potential density; thick lines are plotted every 58C, 1.0 psu, and
1.0 su.
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KASHINO ET AL.
likely in relation to the MD generation. The depth of the
salinity maximum layer around 100 m at these sites
(representing the NPTW) varied seasonally with thermocline depth.
The 258C isotherm depth at T11 was generally shallow
during winter and in relation to the MD generation at
T10 and T14. However, the pattern of variation at T11
below the thermocline differed from those at T10 and
T14; most notably, high-salinity water exceeding 35 psu
frequently appeared in the 100–200-m layer. A salinity
maximum always occurred around this depth, but this
more saline water was observed from winter to spring.
In association with the appearance of this high-salinity
water, the intervals of isothermal and isopycnal lines increased. This result indicates that the saline water exceeding 35 psu from winter to spring has lower potential
vorticity than water with a salinity maximum during other
seasons (NPTW). As described above, the saline water
with low potential vorticity is the SPTW that originated
in the South Pacific. Because the NECC core, which coincides with a property front (e.g., salinity and potential
vorticity) between the NPTW and SPTW (Gouriou and
Toole 1993), flows near T11 at this longitude (Kashino
et al. 2007), both tropical waters were observed at this
site.
This result indicates that ocean variability at this site is
not only associated with an ascending/descending thermocline but also with water mass change. Furthermore,
this result also indicates a seasonal horizontal shift of the
NECC axis; the NECC in the thermocline layer probably migrates north from winter to spring. Ocean surface
current analysis–real time (OSCAR) data also show the
northward migration of the NECC at the surface during
this period (not shown).
These features can be clearly seen in the annual cycle
of salinity variability (Fig. 5). Salinity variations at T10
and T14 were small compared with the salinity variation
at T11 for all layers and had amplitudes on the order
of the salinity measurement accuracy of the TRITON
buoys. This result implies that ocean variability below
the upper ocean (0–50 m) at T10 and T14 can be mostly
attributed to the vertical motion of the thermocline rather
than to variations of waters. On the other hand, the amplitude of seasonal salinity variability at the subsurface
(100–300 m) at T11 was 0.15 psu. This amplitude largely
exceeded the observed seasonal salinity variations of the
cores of the NPTW and SPTW, which were 0.03 (Fig. 5a)
and 0.06 psu (derived from TRITON at 08, 1388E; not
shown). Therefore, salinity variability at the subsurface
at T11 is not solely due to the salinity variability of the
NPTW and SPTW but likely due to the water mass
change associated with the seasonal meridional migration of the salinity front at the NECC axis. Notably, large
293
FIG. 5. Seasonal cycle of salinity variability on potential density
surfaces at 24.0 (thick line), 25.0 (dashed line), and 26.5 su (thin
line) at (a) T10, (b) T11, and (c) T14. Monthly averaged values of
anomaly from the annual mean are plotted.
annual and semiannual signals are also seen in the intermediate layer (26.5 su) of T11, compared with those
at T10 and T14.
Qu et al. (2008) recently discussed semiannual variation in the western tropical Pacific. Using altimeter data
and numerical simulation results, they found that the
semiannual signal was comparable to the annual signal
in that region. A semiannual signal was also seen in time
series data from the TRITON buoys (Fig. 4), although
it was not larger than the annual signal (Fig. 6). The
semiannual signal was particularly large at T10 and T11,
where thermocline descent occurs around spring. At
T14, no clear semiannual signal was detected around the
thermocline (around 100-m depth), although this signal
was found below 200 m.
To evaluate the annual and semiannual signals, the
harmonic analysis results for these two signals in dynamic
height relative to 500 dbar were used (Fig. 6). We selected
1 July 2003 to 30 June 2006 as the period of this analysis,
during which normal-type annual cycles were observed;
according to Mitchum and Lukas (1990), the annual signal is modulated by ENSO phenomena in the western
equatorial Pacific. This figure includes results from
TRITON sites farther to the east: T01 (88N, 1568E), T02
(58N, 1568E), and T07 (58N, 1478E). As shown in Fig. 6,
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FIG. 6. (top) Amplitude and (bottom) phase of (left) annual and (right) semiannual harmonics for time series of
dynamic height relative to 500 dbar derived from TRITON buoy data along 58 (thin lines) and 88N (thick lines). The
analysis period was from 1 Jul 2003 to 30 Jun 2006.
the amplitude of the semiannual signal reaches 40%–70%
of the annual signal at sites west of 1408E. This result is
consistent with Fig. 1 of Qu et al. (2008). To check the
results presented in Fig. 6, harmonic analysis results using
Archiving, Validation, and Interpretation of Satellite
Oceanographic (AVISO) sea surface anomaly data (available online at http://www.aviso.oceanobs.com/) for the
same period as in the figure are plotted in Fig. 7. Both
results almost agree, although some minor peaks shown
in Fig. 7 were not resolved in Fig. 6.
Interestingly, the amplitude of annual and semiannual
signals at 58N increased toward the west, but not at 88N.
The westward propagation speeds of the annual signal
estimated from the phase difference are 0.27 m s21 at
88N and 0.75 m s21 at 58N. Propagation speeds agree
well with the theoretical speeds of baroclinic Rossby
waves at each latitude, Cr 5 bc2/f 2 (Kessler 1990), where
b is the meridional derivative of the Coriolis parameter
f and c is the baroclinic gravity wave speed. This result
suggests that the variation in dynamic height was generated
FIG. 7. As in Fig. 6 but of the sea surface height anomaly derived from AVISO along 4.98 (thin lines) and 7.98N
(thick lines).
FEBRUARY 2011
KASHINO ET AL.
not only by local forcing but also by the propagation of
Rossby waves from the east. Propagation speed at 58N
was also similar to that of the gravest meridional mode
(n 5 1) of equatorial-trapped baroclinic Rossby waves,
Cr 5 c/(2n 1 1) (Philander 1990). Ekman pumping due
to wind stress is a primary forcing for the westward radiation of Rossby waves in the off-equatorial and midlatitude regions, but equatorial-trapped Rossby waves
may also play a role in the variability at 58N. The relationship between wind forcing and variation in heat
content is discussed in the next section.
Semiannual variability in temperature and annual variability in salinity were large near the surface at all sites
(Fig. 4). Near-surface temperature always fell during
winter (January–March) at these sites, probably in association with the MD, which occurred due to positive wind
stress curl in this season. This result agrees with that of
Qu (2003), who showed that incoming surface heat flux
is balanced mainly by vertical entrainment in the MD
region. When near-surface salinity was low, near-surface
temperature tended to be high. This result was often
observed in the region where the barrier layer was developed (e.g., Ando and McPhaden 1997).
Next, we focus on observed interannual variability.
Note that there were three warm episodes (2002–03,
2004–05, and 2006–07) and one cold episode (2007–08)
during the observation period (see online at http://www.
cpc.ncep.noaa.gov/products/analysis_monitoring/ensostuff/
ensoyears.shtml). Among the three warm events, that in
2004–05 was relatively weak.
The most interesting result found in the depth–time
figures (Fig. 4) was a gradual deepening of the thermocline depth from 2003 to 2006 at T14; the depth of 258C in
autumn occurred at approximately 60 m in 2003, 70 m
in 2004, 80 m in 2005, and 100 m in 2006. Similar deepening was also found at T10 but was not clearly identified
at T11.
To clarify variability, we plotted 120-day low-pass
filtered time series of temperature at 100 m and the heat
content anomaly above the 300-m depth in Fig. 8. In this
study, heat content anomaly is defined as the vertically
averaged temperature from the sea surface to 300-m
depth minus its time average for the observation period.
The heat content anomaly at all sites increased annually
from 2003 to 2006 and was particularly notable after
2005. The temperature at 100-m depth at T10 and T14
also rose in the same manner as the heat content anomaly. However, the interannual variability of temperature
at 100 m was different from the heat content at T11;
temperature at 100 m did not rise from 2003 to 2005, but
the heat content anomaly increased during this period.
(The annual signal was large in the heat content anomaly
but not large in temperature at 100 m.) This difference
295
FIG. 8. The 120-day low-pass filtered time series of (a) temperature at 100-m depth and (b) heat content anomaly above 300-m at
T10 (thin line), T11 (dashed line), and T14 (thick line).
can be explained by the change in observed waters due
to the meridional migration of the NECC, as described
previously.
During the 2002–03 El Niño, the heat content anomaly decreased largely at all sites. This result suggests
the discharge of heat that had built up in this region until
the 2002–03 El Niño. Heat storage seems to have occurred until the 2006–07 El Niño. Because the warm
event of 2004–05 was not large, this event does not seem
to have affected the heat content of these sites. During
the 2006–07 El Niño, heat content again largely decreased.
Temperature at 100 m and the heat content anomaly did
not fall in the winter of the 2007–08 La Niña; that is, heat
content in the western equatorial Pacific seems to have
remained high. This result is consistent with that of
Kashino et al. (2009), who identified a rise of dynamic
height from the 2006–07 El Niño to the 2007–08 La Niña
using data from two ship-based observations. It is interesting that the decrease in heat content during the
2002–03 El Niño started at T11 and moved to T10 and
finally T14. Heat content increases after this El Niño
also appeared to start from eastern sites.
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JOURNAL OF PHYSICAL OCEANOGRAPHY
VOLUME 41
FIG. 9. As in Fig. 6 but of Ekman pumping velocity derived from surface wind data provided by the ECMWF along
5.18 (thin lines) and 8.48N (thick lines). To calculate wind stress, we used values of 1.5 3 1023 for the drag coefficient
and 1.18 kg m23 for air density.
The annual amplitudes of temperature variability at
100 m and of heat contents during the 2002–03 El Niño
and 2006–07 El Niño were larger than those in other
years, consistent with the findings of Mitchum and Lukas
(1990).
4. Discussion
Here, we present time series of underwater observations of the MD taken by three surface moorings, namely
the TRITON buoys. Although some time series observation results have been published for areas in and
around the Philippine Sea (e.g., Kessler 1990; Mitchum
and Lukas 1990; White and Tai 1992), these findings
were based on data from expendable bathythermograph
(XBT) observations by volunteer ships, satellite observations, and tide-gauge monitoring. Details of variability,
particularly the variability in temperature and salinity
time series, inside the ocean have not been clarified until
now. Our observations are the first long-term time series
observed below the sea surface in the Philippine Sea and
have revealed some interesting results.
One interesting finding is from the salinity observation
at T11. Because the current axis of the NECC corresponds
with the northern boundary of the SPTW (Gouriou and
Toole 1993), we can speculate on the variability of the
NECC around T11. If we assume that the NECC boundary is at the 35-psu contour [which we consider to be a
good assumption based on Fig. 2 of Kashino et al. (2007)],
then the NECC would seasonally migrate northward
from winter to spring. Interestingly, no water of salinity
greater than 35 psu was observed during these seasons
in 2003, the year of the strongest ENSO from 2002 to
2008. This result agrees with that of Qiu and Joyce (1992),
who showed that the NECC shifted southward during
ENSO years. Underwater salinity measurements are
important not only for observation of waters but also
for precise estimation of the dynamic height in the warm
water pool region (Ueki et al. 2002).
Next, we discuss observed annual and semiannual variations from TRITON buoy data. For this discussion, the
amplitudes and phases of annual and semiannual harmonics of Ekman pumping velocity, based on ECMWF
surface wind data, are plotted in Fig. 9. As in Fig. 6, the
annual signal amplitude at 5.18N increased toward
the west from around 1478N, but that at 8.48N did not. The
phase of the annual signal at 5.18N shows that Ekman
pumping was maximum in a 30–60-day phase, with no
remarkable longitudinal difference west of 1478E. This
finding is consistent with that of Qu et al. (2008), who
reported that the propagating signal in local Ekman
pumping disappears west of the date line. The results
of harmonic analysis of the buoy and wind data seem
to suggest that the local response to local wind forcing
may play a role in the annual signal of the dynamic
height.
These results can also be checked using optimal
analysis data derived from Argo floats (Fig. 10). Clear
seasonal variability appeared east of 1608W at 5.58N
(Fig. 10a), with a warm anomaly in autumn and a cold
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KASHINO ET AL.
FIG. 10. Longitude–time plots of the monthly (top) mean and (bottom) anomaly of temperature at 100-m depth
along (left) 5.58N and (right) 8.58N. For the top panels the time mean during the period from 2002 to 2007 was
subtracted from the monthly mean. The plotted values were derived from Argo float monthly optimal analysis data
combined with data from TRITON buoys and some onboard CTD observations. The contour interval is 18C.
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JOURNAL OF PHYSICAL OCEANOGRAPHY
anomaly in spring. The seasonal variation seems to have
been forced by the annual variation in Ekman pumping velocity, which was particularly significant east of
1608W (Fig. 11a). These anomalies were probably influenced by the ‘‘annual ENSO’’ (Tozuka and Yamagata
2003) signals created by westerly/easterly wind anomalies near the equator east of 1508W. As shown in Fig. 10a,
these anomalies propagated as cold and warm Rossby
waves to the west and gradually decayed. The signals
then became strong again west of 1408E. The amplification was probably a local response to the large annual
signal of wind stress curl related to the Asian monsoon
in this region. The MD at this latitude was thus generated as described by Masumoto and Yamagata (1991)
and Tozuka et al. (2002).
The longitude–time plot of the Argo float temperature at 8.58N (Fig. 10b) exhibits different features from
that at 5.58N (Fig. 10a). There is a discontinuity in propagation from ;1508 to 1408W. The annual signal near the
eastern boundary is not clear compared to that at 5.58N.
The annual signal in the central Pacific between approximately 1608E and 1508W seems to exhibit a maximum
in May–June and minimum in November–December.
The annual signal of Ekman pumping shows a minimum (downwelling) in February–April and maximum
(upwelling) in July–September in the same longitudinal
band (Fig. 11b). This variation in annual wind should be
related to the variation in annual temperature. West of
1608E, the annual low-temperature signals propagated
westward without amplification and neared the western boundary in February–March. Thus, the MD at 88N
during the generation period seems to have been influenced by a cold anomaly propagating from the east rather
than by the local wind effect in the western boundary
region.
With regard to semiannual signals, we found that the
phase lag of the variation between 1308 and 1378E corresponded to the propagation speed at 58N, which agreed
with the theoretical value for the baroclinic Rossby waves
(Fig. 6d). This suggests that the semiannual signal at 58N
in the western region was caused by remotely generated
Rossby waves but was probably also modified by the local
wind forcing during the westward propagation, as described by Qu et al. (2008). As for the semiannual signal
at 88N, we cannot explain its mechanism because the
westward propagation speed largely exceeded the theoretical value for baroclinic Rossby waves at this latitude.
Finally, we discuss interannual variability in this region after 2002. As described in section 3, the heat
content at all the sites varied on the time scale of ENSO.
During the 2002–03 El Niño event, the heat content
greatly decreased (Fig. 8b). Thereafter, it then gradually increased and peaked in 2006. During the 2006–07
VOLUME 41
El Niño period, heat content decreased again. This interannual variability was also captured by the Argo float
observations (Figs. 10c,d). It is interesting that these
variations differed slightly at the various locations; for
example, the variation at T14 appeared to be delayed
a little compared to T10 and T11. These variations appeared to express the recharge and discharge of heat
in the western equatorial Pacific. Note that similar interannual variations were not observed at the other
TRITON buoys located south of 28N or at the TAO
buoys at 58 and 88N between 1658E and 1558W (not
shown). This result suggests that the recharge and discharge of heat associated with ENSO-scale variability
after 2002 occurred in the northern off-equatorial region (north of 58N) of the western Pacific. On the basis
of a numerical experiment, Ishida et al. (2008) reported
that the meridional warm water transport in the western
boundary region compensates for the interior transport
in the SH, whereas such compensation does not hold in
the NH. Therefore, warm water volume variation and/or
recharge–discharge may be remarkable in the NH but
not in the SH in the western boundary region. The observed results are consistent with their numerical study.
The longitude–time plots of anomaly from monthly
mean Ekman pumping velocity along 5.48 and 8.48N
(Figs. 11c,d) indicate interannual variability related to
El Niño and La Niña events. The plot along 5.48N shows
three remarkable positive (upwelling) anomalies propagating eastward during summer 2002–winter 2002/03,
winter 2003/04–spring 2005, and autumn 2006–winter
2006/07 from 1508E to 1208W. A large negative (downwelling) anomaly is shown during spring–winter 2007 for
all the longitudes. The positive anomalies in 2002/03
and 2006/07 correspond to the 2002–03 El Niño and
2006–07 El Niño, respectively. A weak warm event also
occurred in 2004–05. The negative anomaly corresponds
to the 2007–08 La Niña. Although the relationship between the Ekman pumping anomalies (Fig. 11c) and
temperature anomalies (Fig. 10c) is not straightforward,
the forcing and westward propagation of the forced signals are suggestive of some relations between them. The
three positive Ekman pumping anomalies in the central
Pacific should force the upwelling Rossby waves, which
extend to near the western boundary in about 4 months.
The low-temperature anomaly in winter 2002/03–spring
2003 west of 1608E (Fig. 10c) seems to have been generated by the Rossby waves forced in the central Pacific
during summer 2002–winter 2002/03. Although the lowtemperature anomalies during summer 2005 and spring–
summer 2007 may have been related to the positive
Ekman anomalies of 2004/05 and 2006/07, their anomalies were not so significant compared to that during
2002/03. The upwelling Rossby waves may have been
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KASHINO ET AL.
FIG. 11. As in Fig. 10 but of Ekman pumping velocity (31026 m s21) derived from the ECMWF
surface wind along (left) 5.48N and (right) 8.48N. Positive values indicate upward velocity.
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dampened by the negative Ekman pumping (downwelling)
anomalies during autumn 2005–spring 2006 and spring
2007–early 2008. The high-temperature anomalies during
spring–summer 2006 and autumn–winter 2006/07 west
of 1608E seem to be interpretable by the Rossby waves
forced by the negative (downwelling) Ekman pumping
during autumn 2005–spring 2006 and spring 2007–early
winter 2007. Similar high-temperature anomalies are
seen at 8.58N (Fig. 10d).
Kashino et al. (2009) also observed dynamic height
change from the 2006-07 El Niño to the 2007-08 La Niña.
They proposed that this change was not due to local
wind change but to a remote effect from the east. The
above result (heat charging east of 1508E after 2007) is
consistent with their explanation.
5. Summary
The southern Philippine Sea is an interesting and important region not only for oceanographic research but
also for climate research. However, ocean variability in
this region has not been fully clarified. While some researchers have examined ocean variability in the western
equatorial Pacific based on data derived from satellites,
tide gauges, and XBT observations, few have reported on
time series observations in the ocean (e.g., Kashino et al.
2005, 2007) until now. Therefore, although many numerical studies have focused on this region, including the
MD, their results had not been confirmed by observed
data.
TRITON buoys have been deployed in the Pacific and
Indian Oceans since 1998, monitoring the tropical ocean
and atmosphere. We analyzed time series data derived
from underwater sensors of the TRITON buoys at 88N,
1378E (T10); 58N, 1378E (T11); and 88N, 1308E (T14) to
describe ocean variability above the thermocline with
a focus on the MD region. Our results are summarized
as follows.
1) Among semiannual, annual, and interannual signals
in temperature and salinity time series above 300 m
in the MD region, annual signal dominated although
other signals were not negligible. In particular, there
was notable seasonal salinity variability in the subsurface layer (100–300 m) in T11; high-salinity water
(SPTW) exceeding 35 psu appeared at the subsurface from winter to spring. This variability is consistent with the northward migration of the NECC
during these seasons. Temperature and salinity variations at T10 and T14 appeared to be mainly associated with the vertical motion of the thermocline.
2) Thermocline ascent, probably related to the MD, was
observed during winter and spring in all years. The
VOLUME 41
harmonic analysis results for the 208C isotherm depth
indicate that the thermocline at T10 (T14) was shallowest in February (March). Argo optimal analysis
data suggested that this variability at 58N originated
east of 1508W and propagated westward as annual
Rossby waves, with gradual decay. West of 1408E, it
intensified due to annual local wind curl variability.
Thus, an important mechanism of the annual variation of the MD at 58N is the annual variability of local
wind, as also mentioned by Masumoto and Yamagata
(1991) and Tozuka et al. (2002). However, annual
Ekman pumping velocity variability at 88N in the
western Pacific was not large compared to that at
58N (Fig. 9) The longitude–time plot of temperature
from Argo floats at 88N indicates that the annual lowtemperature signals propagated westward without
amplification and approached the western boundary.
Thus, the MD at 88N seems to be influenced by
westward propagating signals from the east.
3) The observed semiannual signal in the MD region
was consistent with the results of Qu et al. (2008).
Its amplitude in dynamic height variation was 40%–
70% that of the annual signal.
4) Interannual variability was observed on the ENSO
time scale. During the 2002–03 El Niño, the heat content in the MD region greatly decreased. Afterward,
heat content gradually increased, peaked in 2006, and
then decreased again during the 2006–07 El Niño.
Large negative anomalies of Ekman pumping velocity (downwelling) that appeared from 2005 to 2006
east of 1508E and north of 58N probably contributed
to the increasing heat content after 2005 in the western
equatorial Pacific.
Warm water volume exchange in the Northern Hemisphere has more impact than that in the Southern Hemisphere on the recharge–discharge of the equatorial warm
water volume (Ishida et al. 2008). For a clear understanding of this recharge–discharge mechanism, further
studies focused on the MD region using the TRITON buoy
data along with altimeter, Argo, numerical simulation, and
other data are needed.
Finally, intraseasonal variability was also seen in
TRITON buoy data time series (not shown). We plan to
describe and discuss the observed intraseasonal variability in a future report.
Acknowledgments. This study was conducted under
the TOCS and TRITON projects of JAMSTEC. The
TRITON buoys were maintained during cruises of the
R/Vs Kaiyo, Mirai, and Yokosuka. We thank the crews
of the R/Vs Kaiyo, Mirai, and Yokosuka as well as
technicians from Marine Works Japan Co., Ltd.; Global
FEBRUARY 2011
KASHINO ET AL.
Ocean Development Inc.; and Nippon Marine Enterprises, Ltd. and all persons related to these JAMSTEC
projects. We also extend our thanks to technicians of
Marine Works Japan Co., Ltd. at JAMSTEC’s Mutsu
Institute of Oceanography for their work with quality control for TRITON buoy data. Yuji Kashino, Akio Ishida,
and Shigeki Hosoda are also grateful for support from the
Japan Society for Promotion of Science through a grantin-aid for scientific research. We also thank three anonymous reviewers for specific and thorough remarks, which
were helpful in the revision of the manuscript.
REFERENCES
Ando, K., and M. J. McPhaden, 1997: Variability of surface layer
hydrography in the tropical Pacific Ocean. J. Geophys. Res.,
102, 23 063–23 078.
——, T. Matsumoto, T. Nagahama, I. Ueki, Y. Takatsuki, and
Y. Kuroda, 2005: Drift characteristics of a moored conductivity–
temperature–depth sensor and correction of salinity data. J. Atmos. Oceanic Technol., 22, 282–291.
Argo Science Team, 1998: On the design and implementation of
Argo: An initial plan for a global array of profiling floats.
GODAE Rep. 5, 32 pp.
——, 2001: Report of the Argo Science Team Second Meeting.
Proc. Argo Science Team Second Meeting, Southampton, United
Kingdom, Southampton Oceanography Centre, 35 pp.
Bingham, F., and R. Lukas, 1994: The southward intrusion of North
Pacific Intermediate Water along the Mindanao coast. J. Phys.
Oceanogr., 24, 141–154.
——, and ——, 1995: The distribution of intermediate water in
the western equatorial Pacific during January–February, 1986.
Deep-Sea Res. II, 42, 1545–1573.
Ffield, A., and A. L. Gordon, 1992: Vertical mixing in the Indonesian thermocline. J. Phys. Oceanogr., 22, 184–195.
Fine, R. A., R. Lukas, F. Bingham, M. J. Warner, and R. H. Gammon,
1994: The western equatorial Pacific: A water mass crossroads.
J. Geophys. Res., 99, 25 063–25 080.
Gordon, A. L., and R. A. Fine, 1996: Pathways of the water between the Pacific and Indian Oceans in the Indonesian seas.
Nature, 379, 146–149.
Gouriou, Y., and J. Toole, 1993: Mean circulation of the upper
layers of the western equatorial Pacific Ocean. J. Geophys.
Res., 98, 22 495–22 520.
Gu, D., and S. G. H. Philander, 1997: Interdecadal climate fluctuations that depend on exchanges between the tropics and extratropics. Science, 275, 805–807.
Hendon, H. H., 2003: Indonesian rainfall variability: Impacts
of ENSO and local air–sea interaction. J. Climate, 16, 1775–
1790.
Hosoda, S., S. Minato, and N. Shikama, 2006: Seasonal temperature variation below the thermocline detected by Argo floats.
Geophys. Res. Lett., 33, L13604, doi:10.1029/2006GL026070.
——, T. Ohira, and T. Nakamura, 2008: A monthly mean dataset of
global oceanic temperature and salinity derived from Argo
float observations. JAMSTEC Rep. Res. Dev., 8, 47–59.
Ishida, A., Y. Kashino, S. Hosoda, and K. Ando, 2008: North-south
asymmetry of warm water volume transport related with
El Niño variability. Geophys. Res. Lett., 35, L18612, doi:10.1029/
2008GL034858.
301
Kaneko, I., Y. Takatsuki, H. Kamiya, and S. Kawae, 1998: Water
property and current distributions along the WHP-P9 section
(1378–1428E) in the western North Pacific. J. Geophys. Res.,
103, 12 959–12 984.
——, ——, and ——, 2001: Circulation of intermediate and deep
waters in the Philippine Sea. J. Oceanogr., 57, 397–420.
Kashino, Y., M. Aoyama, T. Kawano, N. Hendiarti, Syaefudin,
Y. Anantasena, K. Muneyama, and H. Watanabe, 1996: The
water masses between Mindanao and New Guinea. J. Geophys. Res., 101, 12 391–12 400.
——, H. Watanabe, B. Herunadi, M. Aoyama, and D. Hartoyo,
1999: Current variability at the Pacific entrance of the Indonesian Throughflow. J. Geophys. Res., 104, 11 021–11 035.
——, A. Ishida, and Y. Kuroda, 2005: Variability of the Mindanao
Current: Mooring observation results. Geophys. Res. Lett., 32,
L18611, doi:10.1029/2005GL023880.
——, I. Ueki, Y. Kuroda, and A. Purwandani, 2007: Ocean variability north of New Guinea derived from TRITON buoy
data. J. Oceanogr., 63, 545–559.
——, N. España, F. Syamsudin, K. J. Richards, T. Jensen, P. Dutrieux,
and A. Ishida, 2009: Observations of the North Equatorial
Current, Mindanao Current, and Kuroshio Current system
during the 2006/07 El Niño and 2007/08 La Niña. J. Oceanogr.,
65, 325–333.
Kessler, W. S., 1990: Observations of long Rossby waves in the
northern tropical Pacific. J. Geophys. Res., 95, 5183–5217.
Kim, Y. Y., T. Qu, T. Jensen, T. Miyama, H. Mitsudera, H.-W. Kang,
and A. Ishida, 2004: Seasonal and interannual variations of
the North Equatorial Current bifurcation in a high-resolution
OGCM. J. Geophys. Res., 109, C03040, doi:10.1029/2003JC002013.
Kuroda, Y., 2000: Variability of the currents off the northern coast
of New Guinea. J. Oceanogr., 56, 103–116.
——, and Y. Amitani, 2001: TRITON: New ocean and atmosphere
observing buoy network for monitoring ENSO (in Japanese).
Umi no Kenkyu, 10, 157–172.
Lukas, R., 1988: Interannual fluctuations of the Mindanao Current
inferred from sea level. J. Geophys. Res., 93, 6744–6748.
——, E. Firing, P. Hacker, P. L. Richardson, C. A. Collins, R. A. Fine,
and R. Gammon, 1991: Observations of the Mindanao Current
during the Western Equatorial Pacific Ocean Circulation
Study. J. Geophys. Res., 96, 7089–7104.
——, T. Yamagata, and J. P. McCreary, 1996: Pacific low-latitude
western boundary currents and the Indonesian Throughflow.
J. Geophys. Res., 101, 12 209–12 216.
Masumoto, Y., and T. Yamagata, 1991: Response of the western
tropical Pacific to the Asian winter monsoon: The generation
of the Mindanao Dome. J. Phys. Oceanogr., 21, 1386–1398.
Matsuura, H., 2002: Observed variations of upper ocean zonal
currents in the western equatorial Pacific and their relation to
the local wind. J. Geophys. Res., 107, 3210, doi:10.1029/
2000JC000757.
Mitchum, G. T., and R. Lukas, 1990: Westward propagation of
annual sea level and wind signals in the western Pacific Ocean.
J. Climate, 3, 1102–1110.
Philander, S. G., 1990: El Niño, La Niña, and the Southern Oscillation. Academic Press, 293 pp.
Qiu, B., and T. Joyce, 1992: Interannual variability in the mid- and
low-latitude western North Pacific. J. Phys. Oceanogr., 22,
1062–1079.
——, and R. Lukas, 1996: Seasonal and interannual variability of
the North Equatorial Current, the Mindanao Current, and
the Kuroshio along the Pacific western boundary. J. Geophys.
Res., 101, 12 315–12 330.
302
JOURNAL OF PHYSICAL OCEANOGRAPHY
Qu, T., 2003: Mixed layer heat balance in the western North Pacific.
J. Geophys. Res., 108, 3242, doi:10.1029/2002JC001536.
——, H. Mitsudera, and T. Yamagata, 1998: On the western
boundary currents in the Philippine Sea. J. Geophys. Res., 103,
7537–7548.
——, ——, and ——, 1999: A climatology of the circulation and
water mass distribution near the Philippine coast. J. Phys.
Oceanogr., 29, 1488–1505.
——, J. Gan, A. Ishida, Y. Kashino, and T. Tozuka, 2008: Semiannual variation in the western tropical Pacific Ocean. Geophys. Res. Lett., 35, L16602, doi:10.1029/2008GL035058.
Rasmusson, E. M., and T. H. Carpenter, 1982: Variations in tropical sea surface temperature and surface wind fields associated
with the Southern Oscillation/El Niño. Mon. Wea. Rev., 110,
354–384.
Suzuki, T., T. Sakamoto, T. Nishimura, N. Okada, S. Emori,
A. Oka, and H. Hasumi, 2005: Seasonal cycle of the Mindanao Dome in the CCSR/NIES/FRCGC atmosphere-ocean
coupled model. Geophys. Res. Lett., 32, L17604, doi:10.1029/
2005GL023666.
Toole, J. M., R. C. Millard, Z. Wang, and S. Pu, 1990: Observations
of the Pacific North Equatorial Current bifurcation at the
Philippine coast. J. Phys. Oceanogr., 20, 307–318.
Tozuka, T., and T. Yamagata, 2003: Annual ENSO. J. Phys. Oceanogr., 33, 1564–1578.
VOLUME 41
——, T. Kagimoto, Y. Masumoto, and T. Yamagata, 2002: Simulated multiscale variations in the western tropical Pacific: The
Mindanao Dome revisited. J. Phys. Oceanogr., 32, 1338–1359.
Tsuchiya, M., 1968: Upper Waters of the Intertropical Pacific Ocean.
Johns Hopkins Oceanographic Studies, Vol. 4, The Johns
Hopkins Press, 50 pp.
——, R. Lukas, R. A. Fine, E. Firing, and E. Lindstrom, 1989:
Source waters of the Pacific Equatorial Undercurrent. Prog.
Oceanogr., 23, 101–147.
Ueki, I., K. Ando, Y. Kuroda, and K. Kutsuwada, 2002: Salinity
variation and its effect on dynamic height along the 1568E in the
Pacific warm pool. Geophys. Res. Lett., 29, 1689, doi:10.1029/
2001GL013993.
——, Y. Kashino, and Y. Kuroda, 2003: Observation of current
variations off the New Guinea coast including the 1997–1998
El Niño period and their relationship with Sverdrup transport.
J. Geophys. Res., 108, 3243, doi:10.1029/2002JC001611.
Wang, B., R. Wu, and R. Lukas, 1999: Role of the western North
Pacific wind variation in thermocline adjustment and ENSO
phase transition. J. Meteor. Soc. Japan, 77, 1–16.
White, W. B., and C. Tai, 1992: Reflection of interannual Rossby
waves at the maritime western boundary of the tropical Pacific. J. Geophys. Res., 97, 14 305–14 322.
Yuan, X., and L. D. Talley, 1992: Shallow salinity minima in the
North Pacific. J. Phys. Oceanogr., 22, 1302–1316.