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Formation and Subduction of North Pacific Tropical Water and Their
Interannual Variability
SHOTA KATSURA AND EITAROU OKA
Atmosphere and Ocean Research Institute, The University of Tokyo, Kashiwa, JapanBO QIU
Department of Oceanography, University of Hawai‘i at Manoa, Honolulu, Hawaii

NIKLAS SCHNEIDER
Department of Oceanography, and International Pacific Research Center, University of Hawai‘i at Manoa,
Honolulu, Hawaii(Manuscript received 7 February 2013, in final form 25 August 2013)
ABSTRACT
Formation and subduction of the North Pacific Tropical Water (NPTW), its interannual variability, and its
associated mechanisms were investigated by using gridded Argo-profiling float data and various surface flux
data in 2003–11. The NPTW has two formation sites in the center of the North Pacific subtropical gyre,corresponding to two regional sea surface salinity maxima. Mixed layer salinity variations in these two NPTW
formation sites were found to be significantly different. While seasonal variation was prominent in the easternformation site, interannual variation was dominant in the western site. The mixed layer salinity variation in
the eastern site was controlled mainly by evaporation, precipitation, and entrainment of fresher water
below the mixed layer and was closely related to the seasonal variation of the mixed layer depth. In thewestern site, the effect of entrainment is small due to a small vertical difference in salinity across the mixed
layer base, and excess evaporation over precipitation that tended to be balanced by eddy diffusion, whosestrength varied interannually in association with the Pacific decadal oscillation. After subduction, denser
NPTW that formed in the eastern site dissipated quickly, while the lighter one that formed in the westernsite was advected westward as far as the Philippine Sea, transmitting the interannual variation of salinity
away from its formation region.

1. Introduction
Salinity, along with temperature, is one of the mostfundamental parameters in physical oceanography. In
open oceans, its variations are mostly controlled byprocesses near the sea surface, particularly the surface
freshwater flux, and its long-term changes reflect thosein climate and the hydrological cycle. At the same time,salinity also determines water density and controls theocean circulation, particularly at high latitudes wherewater temperature is low. Knowledge of salinity changes
is important for a better understanding of climate andCorresponding author address: Shota Katsura, Atmosphereand Ocean Research Institute, The University of Tokyo, 5-1-5
Kashiwanoha, Kashiwa, Chiba 277-8564 Japan.
E-mail: katsura@aori.u-tokyo.ac.jp
DOI: 10.1175/JPO-D-13-031.1
Ó 2013 American Meteorological Society

the hydrological cycle, particularly because evaporationand precipitation, which occur mostly over the ocean(Schmitt 1995), have not been fully quantified.
Long-term salinity changes have been investigated byprevious studies despite limited observations comparedto those of temperature in the past. Wong et al. (1999)
and Curry et al. (2003) contrasted the World Ocean
Circulation Experiment Hydrographic Program (WHP)
section data from around 1990 with the historical datafrom roughly 20 yr earlier to indicate a freshening ofintermediate waters characterized by a salinity minimum and salinification of tropical waters characterized
by a salinity maximum in the Pacific, Indian, and Atlantic Oceans. Their results implied an intensification of
the global hydrological cycle. A similar intensificationof the salinity contrast between the subtropical andsubarctic World Ocean has been presented by comparing

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KATSURA ET AL.

Argo-profiling float data from after 2000 and historicalmeasurements from before 2000 (Boyer et al. 2005;
Hosoda et al. 2009; Roemmich and Gilson 2009; Durack
and Wijffels 2010; Helm et al. 2010). Isopycnal freshening in the North Pacific permanent thermocline wasalso reported through comparison between the Argo and
WHP data (Ren and Riser 2010).
In addition to the long-term changes, salinity variations on shorter time scales and their mechanisms havebeen explored by using data available at limited locations. At the station A Long-Term Oligotrophic Habitat
Assessment (ALOHA) in the subtropical North Pacific,
monthly shipboard observations during 1991–97 showeda freshening of upper thermocline water in association
with a decadal increase in net freshwater flux (Lukas
2001). In the tropics, mooring observations under the
Tropical Ocean and Global Atmosphere Coupled Ocean–Atmosphere Response Experiment (TOGA-COARE)
and Prediction and Research Moored Array in the Tropical Atlantic (PIRATA) projects made it possible to investigate seasonal-to-interannual variations of salinity inthe western equatorial Pacific (Sprintall and McPhaden
1994; Feng et al. 1998; Cronin and McPhaden 1998) andthe tropical Atlantic (Segschneider et al. 2000; Foltz et al.
2004), respectively. Observations along repeat hydrographic sections and shipping tracks have also clarified
broad-scale salinity variations and their zonal differencesin each tropical ocean (Delcroix and H
enin 1991; Qiu andJoyce 1992; Dessier and Donguy 1994; Donguy and Meyers
1996; Delcroix et al. 1996; Delcroix and Picaut 1998).
In recent years, such exploration of salinity fields has
become possible for most parts of the global upper
ocean, thanks to the accumulation of Argo-profiling
float data since 2000 (e.g., Roemmich et al. 2001). Ren
and Riser (2009) used these data to estimate the mixed
layer salinity (MLS) budget in the northeastern North
Pacific where precipitation exceeds evaporation, andpointed out the importance of the entrainment of saltier
water below the mixed layer for the increase of MLS
from fall to winter. Similar analyses demonstrated that
not only entrainment but also horizontal advection andsea ice processes are important for MLS variations in theSouthern Ocean (Dong et al. 2009; Ren et al. 2011).
Motivated further by the recent launches of the Soil
Moisture Ocean Salinity (SMOS) satellite with Microwave Imaging Radiometer using Aperture Synthesis(MIRAS) by the European Space Agency (Kerr et al.
2001) and the Satelite de Aplicaciones Cientificas-D(SAC-D) satellite with the Aquarius sensor by the National Aeronautics and Space Administration (NASA;Le Vine et al. 2007), upper-ocean salinity variability has
attracted increasing attention, which led to the initiation of intensive studies such as the ongoing Salinity

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Processes in the Upper Ocean Regional Study by NASA.
Clarification of the seasonal and interannual variations ofsalinity and their mechanisms is not only important forunderstanding air–sea heat and freshwater exchanges but
also for accurate estimation of long-term salinity changes.
In this study, we focus on the North Pacific Tropical
Water (NPTW), a water mass characterized by a surface
ocean salinity maximum. NPTW is formed at the sea
surface in the subtropics where evaporation exceeds
precipitation (Cannon 1966), and is then subducted andadvected westward by the North Equatorial Current asa subsurface salinity maximum (Tsuchiya 1968; Qu et al.
1999). When NPTW reaches the Philippine coast (Li andWang 2012), it is split into two parts. One part is advected northward by the Kuroshio, while the other partflows southward in the Mindanao Current. The latterpart partly passes the Celebes Sea (Fine et al. 1994) andthen flows into the Indian Ocean (Ffield and Gordon
1992) or back to the North Pacific (Lukas et al. 1991;
Kashino et al. 1996). When NPTW reaches the equatorial region carrying salinity or spiciness anomalies fromthe formation region in the subtropics, it can influence
the sea surface temperature variability through the formation of a barrier layer (Lukas and Lindstrom 1991) andclimate in the Pacific region (Gu and Philander 1997;
Schneider et al. 1999; Sasaki et al. 2010). NPTW carriednorthward by the Kuroshio can affect the salinity/spiciness
distribution in the permanent pycnocline of the subtropical gyre through the formation of mode waters in
the Kuroshio Extension region (Oka 2009; Oka et al.
2011; Oka and Qiu 2012).
Long-term variation of NPTW was investigated by
Suga et al. (2000), who analyzed repeat hydrographic
section data along 1378E during 1967–95, which is located to the west (downstream) of the NPTW formation
region. They demonstrated salinification of NPTW in
association with the climate regime shift in the 1970s(Nitta and Yamada 1989; Trenberth 1990) that is characterized by a southward shift and intensification of the
Aleutian low in winter. Specifically, the northern part of
NPTW in the 1378E section increased both its salinity
and thickness, while the southern part increased only its
salinity. By using limited surface flux data, Suga et al.
(2000) argued that the southern salinification is possibly
due to the changes of evaporation over precipitation andEkman salt convergence in the NPTW formation region,
while the northern one may be due to the increase of the
NPTW formation rate in the formation region. Because
climatological streamfunction implies that the northern(southern) part of NPTW at 1378E is formed in thewestern (eastern) site of the formation region, the different pattern of NPTW change within the 1378E sectionin the 1970s suggests that the NPTW signals varied

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FIG. 1. Time series of the interpolation error in 10-dbar salinityof MOAA GPV, averaged in the subtropical North Pacific at 158–
308N, 1658E–1258W.

between the western and eastern sites of their formation
region.
In this study, we investigate seasonal and interannual
variations of NPTW in the two sites of its formation
region and their mechanisms using Argo data and thelatest surface flux data. Also explored is how NPTW is
subducted and advected to the downstream as the subsurface salinity maximum, transmitting salinity anomalies generated in the formation region. We believe thatthe results of this study are important for understanding
not only air–sea interaction but also sea surface salinity(SSS) variations that are now continuously monitoredby SMOS/MIRAS and SAC-D/Aquarius. The data
and methods are described in section 2. Variations of
NPTW in its formation region and after subduction as
well as its mechanisms are examined in section 3. The
NPTW formation mechanism and its zonal difference
between the western and eastern sites of the formation
region are further discussed in section 4. A summary isgiven in section 5.

2. Data and methods
a. Data
We used monthly mean temperature and salinity
data based on an Argo-profiling float, moored buoy, andshipboard observations, named Grid Point Value of theMonthly Objective Analysis using the Argo data
(MOAA GPV; Hosoda et al. 2008), in 2003–11. Thisdataset was made by optimally interpolating the observed temperature and salinity anomalies from theWorld Ocean Atlas 2001 climatology (Conkright et al.
2002) onto 18 3 18 grid points at each standard depth. In
this study, the temperature and salinity data between 10
and 2000 dbar at each grid point were vertically interpolated at an interval of 1 dbar using the Akima spline(Akima 1970). The analysis period is validated by the

FIG. 2. Distributions of (a) S at 10-dbar depth and (b) E 2 P(cm month21), (c) E (cm month21), and (d) P (cm month21) at thesea surface in the North Pacific, averaged in 2003–11. Dashedcontours in (a),(b) indicate S 5 35.35 and E 2 P 5 9 cm month21,
respectively. Gray shading in (b) denotes negative values. Fourthick rectangles indicate the NW, NE, SW, and SE boxes.

interpolation error of MOAA GPV salinity in the subtropical North Pacific, which rapidly decreased after
2001 and stayed small after 2003 corresponding to theimproved Argo float coverage (Fig. 1).
We also used reanalysis evaporation data from theobjectively analyzed air–sea heat fluxes project (Yu andWeller 2007; Yu et al. 2008), precipitation from the Climate Prediction Center Merged Analysis of Precipitation(CMAP; Xie and Arkin 1996, 1997), and wind stress fromthe National Centers of Environmental Prediction(NCEP; Kistler et al. 2001), whose horizontal resolutions
are 18 3 18, 2.58 3 2.58, and 1.8758 3 ;28, respectively. Allof these surface forcing datasets were not available when
Suga et al. (2000) conducted their analysis. Sea surface
height data from the Archiving, Validation, and Interpretation of Satellite Oceanographic data (AVISO) on
a 0.258 3 0.258 grid were also used. These data for eachmonth from 2003 through 2011 were linearly interpolated
onto the same grid points as MOAA GPV, after averaging the daily data from NCEP and AVISO for eachmonth. CMAP precipitation data in December 2011

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FIG. 3. Time series of MLS in the (a) NW, (b) NE, (c) SW, and (d) SE boxes in 2003–11 (thin curve). Thick curve
indicates the 12-month moving average.

were not available at the time of this analysis, and weused the December average of 2003–10 as a substitute.
It should be mentioned that the analysis of this study
was repeated using the evaporation data from NCEP

and the precipitation data from NCEP and the Global
Precipitation Climatology Project (Huffman et al. 1997;
Adler et al. 2003). The results were not significantly
different and are not shown.

FIG. 4. Monthly average of MLS in 2003–11 in the (a) NW, (b) NE, (c) SW, and (d) SE boxes. Vertical bars indicate
the 95% confidence interval.

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FIG. 5. (a) Time series during 2003–11 of rate of change in MLS [lhs of Eq. (1); thick curve] and the sum of forcing
terms [rhs of Eq. (1); thin curve] and (b) their monthly average in the NE box. (c) Time series during 2003–11 of each
forcing term and (d) their monthly average in the NE box: E term (red curve), P term (blue curve), Ekman advection
term (green curve), geostrophic advection term (orange curve), and entrainment term (purple curve). Vertical bars in(b) indicate the statistical error estimated from the 95% confidence interval of the forcing terms.

b. Mixed layer salinity budget
To investigate the mechanism of MLS variation, weused the budget equation of Ren and Riser (2009)w DS›S (E 2 P)S
2 uE  $S 2 ug  $S 2 e ,
5
›t
hm
hm

(1)

where S is MLS, t is time, E is evaporation, P is precipitation, hm is mixed layer depth (MLD), uE is Ekman

velocity, ug is geostrophic velocity, we is entrainment
velocity through the mixed layer base, $ is the horizontal
differential operator, and DS is the salinity difference
between the mixed layer base and 20 dbar below (former
minus latter; e.g., Hayes et al. 1991). Here, MLD at eachgrid point was determined as the depth at which potential density su increases by 0.125 kg m23 from the
10-dbar depth (Levitus 1982).
Ekman velocity was estimated from wind stress dataas

FIG. 6. Monthly average of MLD in 2003–11 in the (a) NW and (b) NE boxes. Vertical bars indicate the 95%
confidence interval.

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FIG. 7. As in Fig. 5, but for the NW box.

uE 5

1
(t y , 2t x ) ,r0 fhm(2)

where t x and t y are zonal and meridional wind stresses(positive eastward and northward), f is the Coriolis parameter, and r0 is the reference density of seawater,taken to be 1025 kg m23. Geostrophic velocity was calculated from sea surface height data assuming geostrophy.
Entrainment velocity was estimated as

 

›hm
›hm
(3)
1 $  hm u H
1 $  hm u ,
we 5
›t
›t
where u is the sum of geostrophic and Ekman velocity,
H is the Heaviside step function, and
1, x $ 0
H(x) 5.
(4)0, x , 0Here, we only consider the net positive entrainment
velocity because detrainment does not affect the property of mixed layer.

3. Resulta. Mixed layer salinity variation
The mean SSS distribution in the North Pacific during
2003–11 shows a maximum exceeding 35.0 in thesubtropics at 208–308N, 1558E–1358W (Fig. 2a). NPTW isformed in this region as a winter surface mixed layer
with high salinity (Cannon 1966; Suga et al. 2000; O’Connor
et al. 2002). A further inspection of Fig. 2a reveals two
separate SSS maxima exceeding 35.35 centered at 258N,
1808 and 278N, 1558W, respectively. The double SSS
maxima are also seen in the surface salinity maps ofprevious works (e.g., Suga et al. 2000, 2004; O’Connor
et al. 2002; Delcroix et al. 2005). The double maxima
roughly correspond to those of evaporation minus precipitation (E 2 P) located to their south/southeast(Fig. 2b). These double E 2 P maxima have different
evaporation and precipitation contributions; the western maximum is due to high evaporation, while theeastern one is due to low precipitation (Figs. 2c and 2d).
This difference suggests that the salinity variations andtheir mechanisms may be distinct between the westernand eastern sites of the NPTW formation region.
To investigate the regional difference in the NPTW
variation around the formation region, four boxes wereselected in association with the double SSS and E 2 P
maxima (Fig. 2): northwest (NW; 218–288N, 1658E–1728W),
northeast (NE; 248–308N, 1658–1408W), southwest (SW;
158–208N, 1658E–1708W), and southeast (SE; 168–248N,
1508–1258W). For each box, MLS was calculated foreach month by averaging salinities at all 18 3 18 grid
points within the box from the sea surface to the MLD.

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FIG. 8. As in Fig. 5, but for the SE box.

The obtained MLS variations were significantly different between the two SSS maximum boxes (Figs. 3aand 3b). Interannual variation was dominant in the NW
box, while seasonal change was prevalent in the NE box.
The MLS in the NE box was lowest in late winter around
March and highest in late summer around September–
October almost every year (Figs. 3b and 4b). The MLS
in the NW box was not significantly different among
months, although semiannual variation is discernible(Figs. 3a and 4a). Its interannual variation was characterized by a rapid increase of more than 0.2 in 2008–10.
MLS in the SW and SE boxes (Figs. 3c and 3d) showedsimilar variations to the NW and NE boxes, respectively.
MLS in the SE box changed seasonally, being lowest
around March and highest in November (Fig. 4d). In the
SW box (Fig. 3c), interannual variation was dominant as in
the NW box, and MLS increased rapidly in 2007–09, preceding that in the NW box by about one year. MLS in thisbox also changed seasonally, being highest in December
and lowest in May on the seasonal average (Fig. 4c).

b. Mixed layer salinity budget
To investigate the mechanism of MLS variation in eachbox, variations of the terms in Eq. (1) were estimated. In
the NE box, the rate of change in MLS on the left-handside of Eq. (1) and the sum of the forcing terms on theright-hand side of Eq. (1) corresponded well to eachother in terms of amplitude and phase, showing a highcorrelation both in the full time series (coefficient R 5
0.68; Fig. 5a) and in the seasonal average (R 5 0.85; Fig.
5b). The effect of Ekman and geostrophic advection was
small due to the small horizontal salinity gradient in
the SSS maximum area, and the variation of MLS was
mainly forced by vertical processes, that is, evaporation,
precipitation, and entrainment (Figs. 5c and 5d). In thisbox, evaporation almost always exceeded precipitation.
The excess evaporation increased MLS during thewarming period from March through September, whenthe mixed layer shoaled and then stayed shallow (Fig. 6b).
On the other hand, during the cooling period from
October through February when the mixed layer deepened, the effect of entrainment of underlying fresher
water into the mixed layer exceeded that of the excess
evaporation, decreasing MLS. Thus, MLS fluctuated inassociation with MLD, being highest (lowest) in
September–October (March) when the mixed layer was
shallow (deep), as demonstrated in Fig. 4b. Such an MLS
variation and its budget contrast with those in the subarctic North Pacific (Ren and Riser 2009), where salinity
increases with depth and, resultantly, MLS is highest(lowest) in March–April (September–October).
In the NW box, the rate of change in MLS showedsemiannual variation, while the sum of forcing terms

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FIG. 9. As in Fig. 5, but for the SW box.

showed annual variation (Figs. 7a and 7b). As a result,
the rate of change in MLS and the sum of the forcing
terms were poorly correlated both on the full time series(R 5 20.02; Fig. 7a) and on the seasonal average (R 5
20.03; Fig. 7b). In this box, the effect of entrainment wasmuch smaller than the NE box (Figs. 5c, 5d, 7c, and 7d).
The MLD and its temporal change in this box were smaller
than the NE box (Figs. 6a and 6b), as previously demonstrated (Delcroix et al. 2005), which makes both we and hm
in the entrainment term [Eq. (1)] smaller. The small entrainment was primarily due to a smaller DS in this box,
which will be discussed in section 4. Thus, the forcing
terms in this box were dominated by evaporation andprecipitation (Figs. 7c and 7d). Because evaporation wasalways higher than precipitation, the sum of forcing terms
was always positive and the MLS budget did not close.
In the SE box, the sum of the forcing terms was almost
always larger than the rate of change in MLS and had alarge value in December (Figs. 8a and 8b), likely because of the underestimation of MLD in the southernpart of the box using MOAA GPV data. Without
the December values, the correlation between the twoterms was moderately positive both on the full time series (R 5 0.40; Fig. 8a) and on the seasonal average (R 5
0.69; Fig. 8b). The forcing terms are notably different
from the NE box in two aspects. First, the Ekman term
had a large contribution to decrease MLS (Figs. 8c and8d), unlike the negligible contribution in the NE box(Figs. 5c and 5d). The large Ekman term is caused bya large northwestward gradient of MLS (Fig. 2a) andthe strong northeasterly trade winds in this box; ittransports salinity northward to the SSS maximum area,that is, the NPTW formation region. Second, the entrainment term worked to increase MLS, which was
opposite to the NE box. This indicates the existence ofsaltier water below the mixed layer, which is NPTW
subducted from the formation region. In the SW box,
the rate of change in MLS and the sum of the forcing

FIG. 10. Distribution of 10-dbar salinity (solid contours), averaged in January–March 2003–11. Dashed contours denote potential density of 23.0, 23.5, 24.0, 24.5, and 25.0 kg m23 at 10-dbar
depth from the south. Two thick rectangles indicate the NW andNE boxes.

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FIG. 11. Meridional section of salinity with respect to pressure, averaged in
February 2003–11 along (a) 1508W, (c) 1758E, (e) 1558E, and (g) 1358E (black solidline). Dashed lines indicate isopycnals of su 5 23.5, 24.0, and 24.5 kg m23. Graysolid line denotes the mixed layer base. Meridional section of salinity with respect
to su , averaged in August 2003–11 along (b) 1508W, (d) 1758E, (f) 1558E, and(h) 1358E (black line). Gray line indicates the outcrop surface in February.

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FIG. 12. Distribution of salinity (solid contours) and acceleration
potential (m2 s22; dashed contours) on the su 5 24.0 kg m23 isopycnal, averaged in March 2003–11. Gray shading indicates salinity
higher than 35. Two thick rectangles indicate the NW and NE
boxes.

terms did not correspond with each other (Figs. 9aand 9b), with a low correlation both on the full timeseries (R 5 0.07; Fig. 9a) and on the seasonal average(R 5 20.06; Fig. 9b), as in the NW box. The Ekman
term was important for MLS variation (Figs. 9c and9d) and the entrainment term was positive, as in the
SE box.

c. Subduction of NPTW
Given the significantly different variations of MLS
between the western and eastern sites of the NPTW
formation region (Figs. 3a and 3b), it is natural to ask
how NPTW formed in these sites is subducted and thentransported westward by the North Equatorial Current
to impact the downstream regions such as the one observed by Suga et al. (2000) at 1378E. A winter sea surface density map shows that isopycnals of 24.0 and24.5 kg m23 traversed the center of the SSS maximum inthe western and eastern sites of the NPTW formation
region, respectively (Fig. 10). Consistently, at 1508W in
the middle of the NE box, NPTW with su ;24.5 kg m23
was subducted (Figs. 11a and 11b). Where 1758E crosses
the center of the NW box, NPTW with su 5 ;24.0 kg m23
was subducted; however, NPTW of 24.5 kg m23, which isthought to be subducted from the NE box, no longer
appeared as the salinity maximum at this longitude(Figs. 11c and 11d). NPTW subducted from east of Hawaii
might be dissipated rapidly, as suggested in a numericalsimulation (R. Furue 2012, personal communication).
Farther west at 1558 and 1358E, NPTW did not outcrop to
the surface and became gradually lighter to the downstream centered at su 5 23.5–24.0 kg m23 (Figs. 11e–h). In
addition, the isopycnal salinity map on su 5 24.0 kg m23
indicates that the high-salinity tongue departed from the
NW box and extended southwestward along the streamlines (Fig. 12). Thus, NPTW observed as the subsurface

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FIG. 13. Time series during 2003–11 of the 12-month moving
average of the max salinity value at each longitude lying north of
78N and below the mixed layer base.

salinity maximum to the west (downstream) of the formation region originates mainly in the western site of theformation region.
This is confirmed by comparing interannual variations of maximum salinity of NPTW at each longitude(Fig. 13). The interannual variations at 1508W and 1758E
correspond to those of MLS in the NE and NW boxes(Figs. 3a and 3b), respectively. Farther west, the interannual variation in the NW box propagated westward,
decreasing salinity values as a result of dissipation. For
example, a salinity minimum observed around the endof 2007 at 1758E was seen in early 2008 at 1658E, in late2008 at 1558 and 1458E, and in early 2009 at 1358E;
a salinity maximum observed mid-2006 at 1758E wasseen in early 2007 at 1658E, in late 2007 at 1558E, in early2008 at 1458E, and in mid-2008 at 1358E. This westward
propagation was also seen in a salinity anomaly distribution on the su 5 24.0 kg m23 isopycnal (Fig. 14). A
positive anomaly was generated around 1608E in 2003
and advected westward, reaching 1358E in mid-2005. A
negative anomaly occurred around 1658E in early 2008
and arrived at 1358E in mid-2009. These propagations
from the NW box to 1358E took about two years, being
consistent with Fig. 13. This advection time is much
shorter than that (;5 years) estimated from the World
Ocean Atlas climatology (Bingham et al. 2002) and alsothan that (;5 years) observed on denser isopycnals ofsu 5 ;25.5 kg m23 (Sasaki et al. 2010; Kolodziejczyk
and Gaillard 2012). The westward advection speed estimated from the propagations in Figs. 13 and 14 is;10 cm s21, which is comparable to the westward current speed of the North Equatorial Current on the isopycnal surface of su 5 24.0 kg m23 estimated from MOAAGPV data, for example, one derived from the acceleration potential distribution in Fig. 12.

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FIG. 14. Longitude–time plot of 12-month moving average ofsalinity anomaly on the su 5 24.0 kg m23 isopycnal. The anomaly isrelative to the 2003–11 average at each longitude and is averaged
between 158 and 208N. Gray shading denotes negative values.

4. Discussion
MLS showed a clear seasonal cycle in the NE box(Figs. 3b and 4b) and not in the NW box (Figs. 3a and4a). In order for the seasonal cycle to exist, there should
be both periods with positive and negative forcing during the course of a year. In the NE box, the effect ofentrainment became large because of the increase of
MLD and made the sum of the forcing terms negativefrom fall to winter (Fig. 5). In the NW box, on the other
hand, the entrainment term did not become large evenin fall–winter, and the sum of the forcing terms was always positive (Fig. 7). To clarify the reason for a small
entrainment term in the NW box, we compared thevertical profile of salinity in these boxes (Fig. 15a). As
these boxes are located in the NPTW formation region,
salinity was at a maximum in the mixed layer and

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decreased with depth in the underlying layer in both
boxes, with a larger vertical gradient in the NE box thanin the NW box. The salinity difference between themixed layer base and 20 dbar below [DS; Eq. (1)] in the
NE box was about twice larger than in the NW box(Fig. 16). This suggests that the subduction of fresher
water from the north into the layer below the mixed
layer is strong (weak) in the NE (NW) box. Such subduction of fresher water must be important for the seasonal variation of NPTW in the eastern part of theformation region.
A similar zonal tendency in SSS variations has beenobtained for the equatorial Pacific (Delcroix and Henin1991). That is, interannual variation was dominant in
the western equatorial Pacific owing to the El Ni~no–Southern Oscillation (ENSO)-related precipitation,
while seasonal variation was prominent in the easternequatorial Pacific. The seasonal cycle of SSS in theeastern equatorial region might be related to the entrainment of the South Pacific Tropical Water (Johnson
and McPhaden 1999; Kessler 1999) and/or South PacificEastern Subtropical Mode Water (Wong and Johnson
2003; Sato and Suga 2009), which are subducted from theeastern South Pacific subtropical gyre and then advectedto the equatorial region, in a similar manner that the
subduction and entrainment of freshwater are important
for the seasonal variation of the NPTW in the eastern partof its formation region.
In the SE box, the entrainment term increased the
MLS (Fig. 8). This is because there is a subsurface salinity maximum layer, that is, NPTW subducted fromthe NE box (Fig. 15b). The mixed layer in the SE box
takes in part of this NPTW, and is then advected back to
the NE box by the Ekman flow, under the positive E 2 P
flux that continuously increases the MLS (Fig. 2b). Because the effect of northward Ekman advection is large
in the SE box and small in the NE box (Figs. 5 and 8),
salinity convergence occurs between two boxes, leading
to the formation of the salinity maximum (i.e., NPTW)
to the north of the E 2 P maximum (Figs. 2a and 2b).
Such a meridional overturning circulation of NPTW inthe eastern regions is schematically illustrated in Fig. 17.
In this circulation, the positive E 2 P flux that increases
the NPTW salinity in the mixed layer is balanced bydissipation and, more importantly, the freshwater flux
from the lower, fresher layer in the formation region.
In the NW box, the salinity budget was not closed(section 3b). This means that the dissipation term, which
is neglected in the MLS budget equation [Eq. (1)], has aconsiderable effect in this box. Actually, past observations based on the dispersion of surface drifters demonstrated that lateral diffusivity is about twice larger in
the NW box than in the NE box (Zhurbas and Oh 2004).

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FIG. 15. Vertical profile of salinity averaged in the (a) NW and NE and (b) SW and SE boxes for January–March
2003–11.

If we assume that the lateral diffusivity is 1 3 104 m2 s21
based on the map from Zhurbas and Oh and use theaverage =2S of 5 3 10212 m22 in the NW box (Fig. 2a),the decrease of MLS by diffusion is estimated to be;0.1 month21. This matches in magnitude with thediscrepancy between the rate of change in MLS and thesum of the forcing terms (Figs. 7a and 7b). Furthermore, Qiu and Chen (2013) analyzed satellite altimeterdata during 1992–2012 and detected that eddy kinetic
energy (EKE) in the subtropical North Pacific exhibitsa larger variation in the NW box than in the NE box andwas stronger in 2004 and 2007, when MLS in the NW
box was relative low (Fig. 3a). Hence, we estimated theresidual A of Eq. (1) asw DS›S (E 2 P)S5
2 uE  $S 2 ug  $S 2 e 2 A, (5)
hm
›t
hm

and EKE, following Qiu and Chen (2013), as

EKE 5

1
2


 
 
g ›h0 2g ›h0 2
1
,
f ›x
f ›y

(6)

where x and y are the zonal and meridional coordinates,
h0 is sea surface height anomaly, and g is the gravity
constant. The obtained EKE and A showed similar interannual variation and are highly correlated with eachother (R 5 0.71; Fig. 18a). Thus, the interannual variation of MLS in the NW box was related to that of EKE.
Moreover, EKE in this box also showed annual variation and was highest around May–June (Fig. 18b), suggesting that semiannual variation of MLS in the NW box(Fig. 4a) is due to the decrease of MLS around June byenhanced eddy diffusion.

FIG. 16. Monthly average of DS in 2003–11 in the (a) NW and (b) NE boxes. Vertical bars indicate the 95%
confidence interval.

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VOLUME 43

FIG. 17. Schematic diagram of the meridional overturning circulation of NPTW in the NE andSE boxes.

According to Qiu and Chen (2013), there is a positive
correlation between EKE in the western subtropicalNorth Pacific and the Pacific decadal oscillation(PDO; Mantua et al. 1997) index at least after 1992,
and the latter led the former by one year. This suggests that the interannual variation of MLS in the NW
box was also related to the PDO. To examine therelation for a sufficiently long period, we compared
the variation of the salinity maximum value along the1378E repeat hydrographic section of the Japan Meteorological Agency (Masuzawa 1967; Qiu and Joyce
1992; Suga et al. 2000) and that of the PDO index(available from http://www.data.kishou.go.jp/shindan/b_1/pdo/winpdo.txt). The salinity maximum value at1378E dropped several years after a period of positive
PDO index, showing the strongest negative correlation with a time lag of 3 years (R 5 20.43; Fig. 19).
This time lag agrees with the sum of a 1-yr lag between the PDO index and the EKE (Qiu and Chen
2013) and a 2-yr period needed for the MLS variation in the NW box to propagate to 1378E (Fig. 13).
Thus, the interannual variation of MLS in the NW
box is likely to be controlled by the PDO-related
eddy diffusion.

surface flux data in 2003–11. The MLS variations weresignificantly different between the western and easternsites of the NPTW formation region. Interannual variation was dominant in the western site, while seasonal
one was dominant in the eastern site.
The MLS variation in the eastern site was controlled
mainly by evaporation, precipitation, and entrainment

5. Summary
Formation and subduction of NPTW and their interannual variation have been investigated by using
the MOAA GPV gridded Argo data and the various

FIG. 18. Time series of (a) 12-month moving average and (b)monthly average in 2003–11 of A (thick curve) and EKE (thincurve) in the NW box.

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FIG. 19. (a) Time series of the winter max salinity value north of
78N and below the mixed layer base in the 1378E repeat hydrographic section (thick curve) and the PDO index (thin curve). (b)Lag correlation between the salinity max value at 1378E and the
PDO index. A positive lag indicates that the PDO index leads.

and was closely related to the mixed layer depth variation.
Specifically, MLS increased from March to September–
October when the mixed layer shoaled and stayed shallow
owing to the excess evaporation, while it decreased in
the rest of year when the mixed layer deepened owing
to the entrainment of fresher water. The existence offresher water below the mixed layer, which is subductedfrom the north of the formation region, is important forthe clear seasonal variation in the eastern site.
In the western site of the formation region where theeffect of entrainment was small due to the small vertical difference in salinity across the mixed layer base,the forcing terms were governed by evaporation minus
precipitation that was always positive throughout theyear, and the salinity budget equation [Eq. (1)] is not
closed. This imbalance is because we neglected theeffect of dissipation in the budget equation. The residual
term of the equation, defined as the difference between
the rate of change in MLS and the sum of the forcing
terms [Eq. (5)], was highly correlated with EKE in thewestern North Pacific which, in turn, was correlated with
the PDO index. Thus, the MLS variation in the westernsite is likely to be controlled by evaporation, precipitation, and PDO-related eddy diffusion, at least after
1992.
From the western and eastern sites of the formation
region, NPTW with su 5 24.0 and 24.5 kg m23 was subducted, respectively. Subsequently the NPTW formed at
the eastern site was dissipated quickly, while that formedat the western site was advected westward, transmitting
PDO-related salinity anomalies from the formation region to the Philippine Sea. It is an important future task to
examine whether theses anomalies survive the intensive
mixing in the Philippine Sea and how they are further
advected and affect upper-ocean processes in the tropical Pacific and Kuroshio Extension regions.
This study has revealed the significant zonal differences in the MLS variation and the associated mechanism within the NPTW formation region. Similar zonal
difference in the SSS variation was reported in the equatorial Pacific (Delcroix and H
enin 1991), although the interannual variation in the western subtropical North
Pacific revealed here is related to the PDO, while that in
the western equatorial Pacific is associated with ENSO
(Delcroix and H
enin 1991). Such zonal and meridional
differences suggest a necessity of further regional studies on the MLS variation. Specifically, we need to clarify
the contribution of horizontal and vertical processes
quantitatively by using Argo data and various surface
flux data in each part of the global ocean. Such studies
would be useful for understanding not only climate andhydrological cycle but also SSS variations monitored by
SMOS/MIRAS and SAC-D/Aquarius.
Acknowledgments. The authors thank Ali Belmadani,
Ryo Furue, Takuya Hasegawa, Shigeki Hosoda, Roger
Lukas, Toshio Suga, and two anonymous reviewers forvaluable comments. Part of this study was conducted during SK’s one-month visit to the University of Hawaii at
Manoa in October 2012 under the Overseas Internship
Program for Outstanding Young Earth and Planetary
Researchers provided by the Department of Earth andPlanetary Science, The University of Tokyo. He appreciates Prof. T. Hibiya, who operates this program. SK is also
supported by the Sasakawa Scientific Research Grant fromthe Japan Science Society. EO is supported by the Ministryof Education, Culture, Sports, Science and Technology,
Japan (MEXT; Grant-in-Aid for Scientific Research on
Innovative Areas under Grants 22106007 and 24121001).
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