J Oceanogr (2012) 68:5–20
DOI 10.1007/s10872-011-0032-5

SPECIAL SECTION: REVIEWNew developments in mode-water research: Dynamic
and climatic effects

Progress of North Pacific mode water research in the past decade
Eitarou Oka • Bo Qiu

Received: 19 February 2011 / Revised: 27 April 2011 / Accepted: 28 April 2011 / Published online: 4 June 2011Ó The Oceanographic Society of Japan and Springer 2011

Abstract This article reviews the progress in research on
North Pacific mode waters of the past decade from thephysical oceanographic perspective. The accumulation ofsatellite altimeter sea surface height data, the rapid growthof the Argo profiling float array, and the advancement in
eddy-resolving ocean general circulation models havegreatly improved the traditional views on the mode watersthat were formed prior to the 1990s based on the analyses ofhistorical temperature/salinity data. Areas where significant
progress was made include: (1) descriptions of the mode
waters’ distributions and properties with fine spatial scales,
particularly in their formation regions in winter whereobservational data had been insufficient; (2) clarifications of
the mode waters’ formation and subduction processes
relating to the large-scale mean circulation, as well as to thetime-varying mesoscale eddy field; (3) impacts of the mode
waters’ circulation and dissipation processes on the climate
and biogeochemical processes; and (4) dynamic versus
thermodynamic causes underlying the mode waters’ decadal changes. In addition to the review, future directions formode water research are also presented.
Keywords Mode waters  North Pacific  Physicaloceanography  Seasonal to decadal variability  Eddy
modifications  Climate and biogeochemical impacts

E. Oka (&)
Atmosphere and Ocean Research Institute,
The University of Tokyo, Kashiwa, Chiba 277-8564, Japane-mail: eoka@aori.u-tokyo.ac.jp
B. QiuDepartment of Oceanography,
University of Hawaii at Manoa, Honolulu, HI 96822, USA

1 Introduction
In the subtropical gyres and part of the subpolar gyres of
the world oceans, there are distinct water masses characterized by a nearly vertically homogeneous layer or low
potential vorticity (PV) lying just above or within thepermanent pycnocline. These water masses, called ‘‘mode
waters,’’ are formed as a deep mixed layer on the warm
side of a strong current or front in late winter because of
convective mixing, and are left in the subsurface as apycnostad after being capped by the seasonal pycnocline inspring. Some part of the formed mode waters is entrained
into the mixed layer in the following winter, modulating
the sea surface temperature, while the remaining part enters
the permanent pycnocline (this process is called subduction) and is then advected to spread over a much wider area
than its formation regions, carrying temperature, salinity,
and PV anomalies. The formation, circulation, and dissipation of mode waters and their variability are related toocean-atmosphere interaction and various upper-ocean
dynamic and thermodynamic processes on a wide range oftime scales. Moreover, they are believed to play animportant role in biogeochemical processes such as theoceanic uptake of atmospheric CO2 (e.g., Bates et al. 2002)
and the nutrient cycling in the oligotrophic subtropicalgyres (e.g., Palter et al. 2005; Krémeur et al. 2009;
Sukigara et al. 2011).
The subtropical gyre in each ocean basin contains three
types of mode waters formed in the western, eastern, andpoleward parts of the gyre, except that the South Hemisphere gyres share the Subantarctic Mode Water formed
just north of their common southern boundary, the subantarctic front (Hanawa and Talley 2001). In the North
Pacific, a thermostad of 16–18°C lying in the upper permanent thermocline in the northwestern part of the

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subtropical gyre was identified and named the subtropicalmode water (STMW) by Masuzawa (1969) because it
produces a mode in the volume distribution on the temperature-salinity diagram for the subtropical gyre.1 Since
then STMW was little explored for nearly 2 decades, butdrew attention again in the late 1980s when the Ocean
Mixed Layer Experiment (Toba et al. 1991) was implemented as one of the Japanese activities in the World
Climate Research Program. Possible roles of STMW as aheat reservoir memorizing the wintertime ocean-atmosphere interaction have inspired a number of studies basedon shipboard observations and data analyses that helped
clarify the formation, circulation, and long-term variationof STMW.
STMW is formed as a deep winter mixed layer justsouth of the Kuroshio and the Kuroshio Extension (KE)between *132°E and near the dateline, where warm surface water supply from the Kuroshio/KE underlying thestrong East Asian winter monsoon results in large oceanic
buoyancy loss in winter (Hanawa 1987; Hanawa andHoshino 1988; Suga and Hanawa 1990; Bingham 1992).
Due to the downstream cooling of the Kuroshio/KE, colder
and denser STMW tends to be formed progressively in thefarther eastern part of the formation region (Suga and
Hanawa 1990; Bingham 1992). After spring, the main bodyof STMW, capped by the seasonal pycnocline and isolated
from atmospheric contact, is advected southwestward by
the ‘‘Kuroshio Countercurrent’’ and is subducted into thepermanent pycnocline (Hanawa 1987; Suga et al. 1989;
Suga and Hanawa 1990, 1995a; Bingham 1992).
The volume and properties of STMW formed every
winter exhibit significant year-to-year variations (Hanawa
1987; Suga et al. 1989). On interannual time scales, thicker
and colder STMW tends to be formed in years withstronger winter monsoons because of larger oceanic heatloss and increased southward Ekman transport (Suga and
Hanawa 1995b; Yasuda and Hanawa 1997, 1999; Taneda
et al. 2000; Hanawa and Kamada 2001; Hanawa andYoritaka 2001). On decadal to interdecadal time scales, on
the other hand, intensification of the westerlies leads to theformation of warmer STMW through gyre spin-up andincreased warm water advection by the Kuroshio (Hanawa
and Kamada 2001). The local existence of STMW south ofJapan is also influenced by path variations of the Kuroshio(Yoshida 1964; Kawabe 1995). The westward advection of
STMW from the region southeast of Japan is blocked when1

The term ‘‘mode water,’’ introduced by Masuzawa (1969), wassubsequently applied to any thick, broadly distributed, near-surface
layer characterized by low PV (Hanawa and Talley 2001). Therefore,
mode waters nowadays are not necessarily characterized by uniformity in terms of temperature and salinity. In other words, temperatureand salinity might be stratified in a compensating way within apycnostad of mode waters.

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the Kuroshio takes a large meander path (Bingham et al.
1992; Suga and Hanawa 1995b, c).
In the late 1990s, the other two mode waters in the North
Pacific were identified in succession, as expected from isopycnal PV maps presented in an earlier study (Talley1988). Nakamura (1996) and Suga et al. (1997) separately
described a thermostad of 8.5–11.5 and 10–13°C lying in
the lower permanent pycnocline in the central part of thesubtropical gyre using different climatological data andtermed it the central mode water (CMW). They inferredthat CMW is formed between KE and its northern bifurcation (Suga et al. 1997) and between the northern bifurcation and the subarctic front2 (Nakamura 1996), which
renders CMW much colder, fresher, and denser than
STMW. The eastern subtropical mode water (ESTMW)
was identified by Hautala and Roemmich (1998) as a weak
thermostad of 16–22°C lying in the upper permanent
pycnocline in the eastern part of the subtropical gyre. It is
formed as a relatively deep winter mixed layer southwestof the subarctic front, which extends in the northwestsoutheast direction in this region (Yuan and Talley 1996;
Suga et al. 2004). It is often stratified in terms of temperature and salinity in a compensating way within its pycnostad, possibly because of cross-frontal intrusions(Hautala and Roemmich 1998; Sprintall and Roemmich
1999).
The two new mode waters formed and advected anticyclonically in the central to eastern part of the gyre contrast with the classical STMW that is largely confined to
the Kuroshio recirculation region in the northwestern partof the gyre. They are interesting from two points of view
(Hautala and Roemmich 1998; Hanawa and Talley 2001).
First, their low PV signature is a good tracer of the ventilation process in the subtropical gyre. Second, if theirvolume and properties formed every winter change fromyear to year, their advection in the permanent pycnocline
might transport temperature and salinity anomalies fromthe mid-latitude ocean surface to the equatorial upwelling
region and cause interdecadal climate variability, as
hypothesized by Hanawa (1996) and Gu and Philander
2

In this review, the term ‘‘subarctic front’’ is used for a densitycompensating front between the warmer, saltier water in thesubtropics and the colder, fresher water in the subarctics, characterized by the outcrop of the 33.0–33.8 isohalines (Roden 1970, 1972;
Zhang and Hanawa 1993; Yuan and Talley 1996), as in the literatureof physical oceanography. This front often separates into two or more
fronts, particularly in the eastern North Pacific, and is also called thesubarctic frontal zone. In the literature of fisheries oceanography (e.g.,
Favorite et al. 1976; Yasuda 2003), the term ‘‘subarctic front’’ isreferred to the front at the southern boundary of subsurface
temperature inversions characterizing the subarctics, represented by
the 4°C isotherm standing almost vertically below the 100-m depth(Uda 1963; Favorite et al. 1976). This front is called the ‘‘polar front’’in this review, as in the literature of physical oceanography.

Progress of North Pacific mode water research(1997). An initial attempt using historical temperature data
(Yasuda and Hanawa 1997) demonstrated that CMW
becomes colder during the decade after the 1976/1977
regime shift (e.g., Nitta and Yamada 1989; Trenberth 1990)
than the previous decade, probably because of larger oceanic heat loss and increased southward Ekman transport
associated with the intensification of the westerlies. On theother hand, Hautala and Roemmich (1998) did not find a
substantial change in the ESTMW volume and temperature
between 1970–1979 and 1991–1997, although the lack ofdata during the intervening period might hinder thedetection of ESTMW changes associated with the
1976/1977 regime shift.
The aforementioned studies on the North Pacific mode
waters prior to around 2000 were comprehensively
reviewed by Hanawa and Talley (2001). In the past decadefollowing their review, mode water research has been
advanced greatly because of the rapid development ofresearch tools. One of the biggest factors is the accumulation of satellite-based data such as sea surface height andthe development of subsurface observing systems represented by Argo (Roemmich et al. 2001). Compared to thehistorical data until the 1990s that had insufficient coverageaway from coasts and were spatially limited to a number of
repeat sections, the unprecedented Argo array of 3,000
profiling floats provides us now with temperature andsalinity data down to 2,000-dbar depth at 10-day intervals
with a horizontal resolution of 3° in latitude and longitude(Roemmich et al. 2009; Freeland et al. 2010). This makes it
possible to take snapshots of the entire structure of mode
waters at various times of year, providing new insights on
their formation, circulation, dissipation, and temporalvariations. In addition, the long-term satellite-based data
have revealed the variability of currents/fronts and eddy
fields that are closely related to mode water variability,
particularly in the energetic Kuroshio-Oyashio Extension
region (e.g., Qiu 2002) where STMW and CMW areformed (Qu et al. 2002). Another important factor is theintroduction and progress of modeling studies of the North
Pacific mode waters. As the model resolution improved
from non-eddy-permitting to eddy-permitting and further
to eddy-resolving (e.g., Hasumi et al. 2010; Masumoto2010), the structures of the currents/fronts and the associated mode waters became reproduced more realistically,
helping to understand the underlying dynamics, particularly on mesoscales that are not fully resolved by observations. It is also worth noting that while most of thestudies until the 1990s were performed by a single researchgroup at Tohoku University, Japan, their activities havestimulated many scientists to participate in the mode water
research, generating a variety of views and approaches.
This paper reviews the progress of physical oceanographic research on the North Pacific mode waters in the7past decade following the review of Hanawa and Talley(2001). Readers may refer to the review of Joyce (2011) inthis issue, and Kelly et al. (2010), for the mode water
research progress in the North Atlantic and the comparison
with the North Pacific. Research progress on the formationand subduction of the North Pacific mode waters is
described in Sect. 2, and that on the other processes such as
their circulation, modification, and dissipation is summarized in Sect. 3. After reviewing the studies on their decadal variability in Sect. 4, we present future directions forthe mode water research in Sect. 5.

2 Formation and subduction
In the classical ventilated thermocline theory (Luyten et al.
1983), water in the subtropical gyre is pushed down (orsubducted) from the base of the Ekman layer into thepermanent pycnocline because of Ekman pumping and is
subsequently transported by the anticyclonic Sverdrup flow
along isopycnal surfaces, conserving its PV. Practically,
water is subducted not from the base of the Ekman layer
but from that of the late winter mixed layer (Stommel1979), which generally shoals equatorward within thesubtropical gyre. Consequently, lateral induction throughthe sloping mixed layer base increases the subduction rate
as compared to the estimation from Ekman pumping alone(Williams 1989, 1991; Marshall et al. 1993; Huang and Qiu
1994; Qiu and Huang 1995).
The PV (Q) of water subducted at the base of the wintermixed layer is expressed as,
Q¼

f ub  rqm
q0 wb þ ub  rh

ð1Þ

where f is the planetary vorticity, q0 the reference density,
qm the mixed layer density, h the mixed layer depth(MLD), ub and wb the horizontal and vertical velocities atthe base of the winter mixed layer, respectively, and r is
the horizontal differential operator (Williams 1989, 1991).
The distribution of -wb (called the vertical pumping term)is relatively uniform over the North Pacific subtropicalgyre (Huang and Qiu 1994), and the water acquires low
Q if -ub  rh is large or -ub  rqm is small. In otherwords, mode waters are formed in the case of large lateralinduction associated with an MLD front (Kubokawa
and Inui 1999; Kubokawa 1999; Nishikawa and Kubokawa
2007) or small density advection associated with asmall downstream gradient of the mixed layer density(Fig. 1).
In the early 2000s, the mechanisms of large-scale formation and subduction of the three mode waters in the
North Pacific were examined, based mainly on the subduction theory including Eq. 1 and by non-eddy-permitting

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E. Oka, B. Qiu

Fig. 1 Schematic illustrating
the two mechanisms of mode
water subduction: a large lateralinduction and b small density
advection. Solid and dashedlines denote isopycnals and thebase of the winter mixed layer,
respectively

models with a horizontal resolution of 1° or 2°. The STMW
formation region is characterized by the deep winter mixed
layer south of the Kuroshio and KE accompanied by a
sharp MLD front at its southern end (Suga and Hanawa
1990; Oka and Suga 2003), and STMW is subducted into
the permanent pycnocline by crossing this front to the south
because of large lateral induction (Kubokawa and Inui
1999; Tsujino and Yasuda 2004). Similarly, CMW enters
the permanent pycnocline by crossing an MLD front at theeastern end of the formation region, namely the deep wintermixed layer north of KE (Xie et al. 2000; Qu et al. 2002;
Tsujino and Yasuda 2004; Hosoda et al. 2004). In contrast,
the ESTMW formation region is characterized by a weak
winter MLD maximum centered at 30°N, 140°W that ismuch shallower than those in the STMW and CMW formation regions. This maximum is formed in spite of relatively uniform, weak wintertime cooling over the eastern
North Pacific because stratus clouds (Ladd and Thompson
2000, 2001) and the saline Ekman transport from the south(Toyoda et al. 2004) preclude the development of theseasonal pycnocline in summer and precondition the
deepening of winter mixed layer. The ESTMW formation
region is also characterized by a small horizontal gradientof the winter mixed layer density because the horizontal
changes of the mixed layer temperature and salinity,
between warmer, saltier water to the southwest and colder,fresher water to the northeast, compensate each other.
Consequently, low PV of ESTMW is generated by thesmall density advection mechanism (Xie et al. 2000;
Hosoda et al. 2001; Ladd and Thompson 2001). These
subduction mechanisms proposed by the modeling studies
are supported by a winter mixed layer climatology constructed by Suga et al. (2004) for STMW and ESTMW, but
not for CMW. In their climatology, the winter MLD
gradually decreases eastward in the eastern part of the
CMW formation region, which implies that the water issubducted mainly because of small density advection (Sugaet al. 2004, 2008). This discrepancy between the models
and the observations is probably because the past models
tend to produce too deep winter mixed layers in the CMW

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formation region, particularly in its eastern part (Ladd andThompson 2001; Tsujino and Yasuda 2004).
As the horizontal resolution of ocean general circulation
models improved from non-eddy-permitting to eddy-permitting, they better reproduced the frontal structures in the
Kuroshio-Oyashio Extension region that are essential forthe formation of STMW and CMW. Qu et al.’s (2002)
model with a resolution of 1/4° (longitude) 9 1/4° (latitude) reproduced the KE and subarctic fronts, and theformation of STMW and CMW south of the respective
fronts. Furthermore, Tsujino and Yasuda’s (2004) modelwith a resolution of 1/4° 9 1/6° reproduced the northernbifurcation of KE (Kuroshio bifurcation) as well and captured the formation of two varieties of CMW, the lighter
variety (L-CMW) south of the Kuroshio bifurcation front
and the denser variety (D-CMW) south of the subarctic
front. Although the Kuroshio bifurcation front in the modelseems somewhat too strong (Tsujino and Yasuda 2004), theformation of two varieties of CMW is actually observed ina synoptic hydrographic section along 179°E (Mecking andWarner 2001) and a repeat section along 165°E (Oka andSuga 2005).
In the latter 2000s, accumulation of in situ observation
data including those from Argo profiling floats enables us
to capture more detailed, unsmoothed structures of mode
waters at various times of year, particularly at higher latitudes in winter where an insufficient number of historical
data had existed because of unfavorable weather conditions
for shipboard measurements. An intense hydrographic
survey carried out northeast of Japan in July 2002 revealed
a new type of pycnostad with h = 5–7°C, S = 33.5–33.9,
and rh = 26.5–26.6 kg m-3 (h and rh are potential temperature and density; S is salinity) around 43°N, 160°E
south of the subarctic front and the polar front3 (Saito et al.
2007), as suggested by earlier studies4 (Yasuda 2003; Oka
3

An intense hydrographic survey conducted in spring 2003 indicatesthat the subarctic front and the polar front coincide with each other inthis longitude range (Eitarou Oka, personal communication 2011).
4
In Yasuda (2003), TRMW is referred to as ‘‘Dense Central Mode
Water.’’

Progress of North Pacific mode water research

9and Suga 2005). This pycnostad is colder, fresher, and alittle denser than D-CMW, and was named the transition
region mode water (TRMW) by Saito et al. (2007) because
its formation region corresponds to the transition region
described in the literature of fisheries oceanography (e.g.,
Favorite et al. 1976; Yasuda 2003). TRMW is often stratified in terms of temperature and salinity within its pycnostad as in the case of ESTMW, probably because of
isopycnal intrusion that characterizes the transition region(Saito et al. 2007).
The TRMW distribution region presented by Saito et al.
(2007) corresponds to the region where the deepest wintermixed layer in the North Pacific appears in previous climatologies (Huang and Qiu 1994; Ladd and Thompson
2000; Suga et al. 2004), although MLD there might beoverestimated because the smoothing process in the climatologies mixes different water types across the subarctic
front, causing artificial cabbeling near the sea surface(Ohno et al. 2004; Suga et al. 2004). This MLD maximum
is possibly generated for the following reasons. First, it liesin the ‘‘stability gap’’ described by Roden (1970) and Yuan
and Talley (1996), a zonal band of lateral minimum in thevertical stability located south of the subarctic front (Laddand Thompson 2000; Suga et al. 2004; Saito et al. 2007).
Second, the quasi-stationary jet associated with the polar
front (Isoguchi et al. 2006) brings warmer and saltier waterfrom the south, enhancing the oceanic heat loss and making
the surface water denser in winter (Saito et al. 2007).
Actually, recently developed high-resolution surface heat
flux data and Argo float data revealed locally enhanced

oceanic heat loss and deep mixed layer formation justsoutheast of the polar front in winter (Hiroyuki Tomita,
personal communication, 2011).
The entire structure of the CMW and TRMW formation
regions and the subduction mechanism of these waters were
presented by a recent study using Argo float data during
2003–2008 (Oka et al. 2011a). To the north of KE, twozonally elongated regions of deep winter mixed layer
extend along 33°–39°N and 39°–43°N from the east coast ofJapan (*142°E) to *160°W. The southern region corresponds to the formation region of L-CMW, while thenorthern region to that of D-CMW and TRMW, in thewestern (eastern) part of which TRMW (D-CMW) is predominantly formed (Fig. 2). Both regions extend beyond
the longitude range of the Kuroshio bifurcation front to thewest and east, suggesting that the two regions are not separated by the bifurcation front, but rather enhance it,
because these two adjacent low PV waters should accompany sharply inclined isopycnals in between. From theeastern part of both regions east of 170°E where the wintermixed layer becomes gradually shallower, warmer, andlighter to the east/downstream, L-CMW with h = 11–15°C,
S = 34.3–34.6, and rh = 25.7–26.2 kg m-3 and D-CMW
and TRMW (mostly the former) with h = 7–11°C,
S = 33.6–34.2, and rh = 26.1–26.4 kg m-3 are subducted
into the permanent pycnocline because of small density
advection as inferred by the previous climatological studies(Suga et al. 2004, 2008). On the other hand, D-CMW and
TRMW (mostly the latter) with rh = 26.5–26.6 kg m-3
formed only in the western part of the formation region

Fig. 2 Schematic illustrating the relationship between the mode
waters and frontal structures in the western to central North Pacific.
PF polar front, SAF subarctic front, KBF Kuroshio bifurcation front,
KEF Kuroshio Extension front. NSTF, SSTF, and ESTF are thenorthern, southern, and eastern subtropical fronts identified in

Kobashi et al. (2006), respectively. Red, green, and blue ovals denotethe formation regions of STMW, L-CMW, and TRMW/D-CMW,
respectively. Pink (light blue) shadings with arrows indicate thespreading of STMW (L-CMW and D-CMW) after subduction

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where winter MLD gradually increases eastward/downstream are entrained into the mixed layer in the following
winter and cannot be subducted into the permanent
pycnocline, which likely explains the large discrepancy inthe temperature-salinity relations between the winter mixed
layer and the permanent pycnocline in this rh range, argued
by Suga et al. (2008). The downstream thickening of CMWs
in the western part of their formation regions preconditions
their permanent subduction in the eastern part, as speculated
by earlier studies (Ladd and Thompson 2000; Mecking andWarner 2001).
In addition to the large-scale mode water formation andsubduction mentioned above, effects of mesoscale eddies(Marshall 1997; Hazeleger and Drijfhout 2000) and theanticyclonic recirculation gyre, both of which characterize
the highly variable and nonlinear KE system (e.g., Qiu
1999, 2002; Ebuchi and Hanawa 2001), on the STMW
formation and subduction have been clarified using Argo
float data combined with altimetric sea surface height data
(Uehara et al. 2003; Pan and Liu 2005; Qiu et al. 2006; Oka
2009; Oka et al. 2011b; Kouketsu et al. 2011) and state-ofthe-art, eddy-resolving models with a horizontal resolutionof 1°/10° or better (Rainville et al. 2007; Nishikawa et al.
2010). The recirculation gyre consists of several anticyclonic circulations associated with crests of the Kuroshio/
KE meander (e.g., Kawai 1972), in each of which thick
STMW with a characteristic temperature is formed in
winter, possibly because water is effectively cooled andhomogenized there (Qiu et al. 2006; Rainville et al. 2007;
Oka 2009; Nishikawa et al. 2010). The formation process isalso modulated by prevailing mesoscale eddies; thicker(thinner) STMW tends to be formed in anticyclonic(cyclonic) eddies, in which background stratification is
weaker (stronger) in association with the deeper (shallower)
permanent pycnocline and oceanic heat loss is larger(smaller) because of warmer (colder) sea surface temperature (Uehara et al. 2003; Pan and Liu 2005; Rainville et al.
2007; Nishikawa et al. 2010; Kouketsu et al. 2011).
After spring, STMW continues to show highly variable
distributions created by mesoscale activities, being thicker(thinner) in the anticyclonic (cyclonic) part of the flow
(Uehara et al. 2003; Rainville et al. 2007; Nishikawa et al.
2010; Oka et al. 2011b). While the recirculation gyre tends
to confine STMW in its formation region (Oka 2009),
anticyclonic eddies transport part of it southward throughthe southern boundary of the region across the westward
mean flow because of southward migration of eddies that
trap STMW, as observed by Takikawa et al. (2005), and to
eddy mixing (Rainville et al. 2007; Nishikawa et al. 2010).
This also causes STMW to cross the winter MLD front andto be thereby subducted into the permanent pycnocline.
Such eddy subduction contributes to approximately half ofthe total STMW subduction rate (Nishikawa et al. 2010).

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Thus, the STMW subduction revealed by Argo float dataand high-resolution models is quite different from the traditional climatology-based picture that the thick body of
STMW is advected by the southwestward mean flow as theseason progresses (Sect. 1).
Subduction across the mean flow has also been reported
for CMW. In synoptic meridional sections in the formation
region, L-CMW (D-CMW) occasionally intrudes southward into the KE (Kuroshio bifurcation) front or exists as
isolated patches south of the front (Mecking and Warner
2001; Oka and Suga 2005). As an extreme case, a highresolution hydrographic survey southeast of Japan found a
D-CMW patch trapped in a subsurface mesoscale eddy
approximately 1,000 km south of KE in fall, which wasestimated to be subducted only a half year ago in association with the instability of the KE front (Oka et al. 2009).
Such cross-frontal subduction could modulate the CMW
distribution determined by the large-scale subduction
by the mean flow, particularly in the density range of
rh = 26.5–26.6 kg m-3, which is not ventilated throughthe large-scale CMW subduction (Oka et al. 2011a).
Furthermore, the CMW formation can also be dominated by mesoscale activities, as inferred from intermittent
CMW distributions based on Argo float data from a single
year (Bingham and Suga 2006) and from eddy-resolving
numerical models (Nishikawa et al. 2010). In fact, a recent
composite analysis of Argo float data and altimetric sea
surface height data revealed that roughly half of wintermixed layers deeper than 150 dbar in the CMW formation
region are formed in anticyclonic eddies (Kouketsu et al.
2011). Past shipboard observations also demonstrated that
inside anticyclonic eddies (or warm-core rings) pinched offfrom the KE, thick STMW is modified to CMW or TRMW
because of winter cooling and interaction with the ambient
water while the eddies are migrating in the region north of
the KE for a few years (Tomosada 1986; Yasuda et al.
1992). To further clarify the role of mesoscale activities in
the CMW formation and subduction, high-resolutionhydrographic surveys north of KE in conjunction witheddy-resolving models are strongly desired.
In addition to the mesoscale variability, shorter timescale variability has also been demonstrated to affect themode water formation. The timing of the deepest mixed
layers in the North Pacific varies among regions, being
February–March (February–April) in the STMW and
L-CMW (D-CMW and TRMW) formation regions (Okaet al. 2007; Ohno et al. 2009). The mixed layer deepening
prior to these ‘‘late winter’’ months is dominated by episodic events on time scales of several days that are related
to the passages of synoptic low-pressure disturbances, as
revealed by in situ buoy observations and high-resolution
numerical model simulations (Qiu et al. 2004; Rainvilleet al. 2007; Cronin et al. 2008; Jensen et al. 2011). The

Progress of North Pacific mode water research

importance of synoptic-scale atmospheric forcing is further
supported by short-term MLD variations observed byprofiling floats (Yoshida and Hoshimoto 2006; Oka et al.
2007), with the use of a relatively small value of temperature and density as a threshold to determine MLDs (e.g.,de Boyer Montégut et al. 2004).

3 Circulation and dissipation
The development of observations and models in the past
decade enabled us to investigate the circulation, modification, and dissipation of the mode waters that had been
inadequately studied until the 1990s. By using data from 20
profiling floats that were concentratively deployed in the
anticyclonic recirculation gyre south of the first quasi-stationary meander of the KE under the Kuroshio Extension
System Study project (KESS; http://www.uskess.org), Qiu
et al. (2006) explored the seasonal evolution of the verticalstructure of STMW in its formation region in 2004. As theseasonal pycnocline develops and gradually thickens after
April, the underlying STMW is eroded from its top,
becoming thinner, whereas its bottom remains largely at aconstant depth of *500 dbar. From a PV budget analysis
for this STMW under the assumptions that horizontal
advection is negligible (because the KE is in a stable statein 2004, as explained in Sect. 4) and that the vertical eddy
diffusivity (Kv) at the lower boundary of STMW is
0.1 9 10-4 m2 s-1 as previously observed in the permanent pycnocline (e.g., Gregg and Sanford 1980; Ledwell
et al. 1993), Kv at the upper boundary of STMW is estimated to be 2–5 9 10-4 m2 s-1, which is an order ofmagnitude larger than that in the permanent pycnocline.
A recent dissolved oxygen budget analysis using data froma profiling float equipped with a fluorometer and an oxygen
sensor estimated a similar Kv value of 1.7 9 10-4 m2 s-1
near the upper boundary of STMW (Sukigara et al. 2011).
Such large vertical diffusivity possibly stems from the
sharp drop in stratification from the seasonal pycnocline to
the STMW pycnostad, which acts as a barrier for the
downward transmission of internal gravity waves generated
in the surface mixed layer and contributes to the enhanced
eddy mixing as these waves are reflected at the upperboundary of STMW (Qiu et al. 2006). The large vertical
diffusivity has important implications not only for themodulation of sea surface temperature in the STMW formation region (e.g., Hanawa and Sugimoto 2004), but also
for the nutrient supply from the subsurface to the oligotrophic surface layer in the subtropical gyre (e.g., Moriet al. 2008; Sukigara et al. 2011).
Meanwhile, shipboard observations using a microstructure profiler indicated that Kv is relatively uniform between
the top and bottom of STMW, and is much smaller than the

11

above estimates, being 10-6 to 10-5 m2 s-1 in summer
2006 and 10-7 to 10-5 m2 s-1 in winter 2007 (Mori et al.
2008). Why is there such a large discrepancy between thebudget-based and observed values of Kv? The most plausible explanation is that like the deepening of the mixed
layer in the fall and early winter seasons (Sect. 2), erosionof STMW is dominated by episodic atmospheric disturbances on time scales of several days, during which Kv
becomes a few orders of magnitude larger but is missed by
shipboard microstructure profiler measurements due torough sea conditions. In fact, an analysis of data from an
acoustic Doppler current profiler moored in the STMW
formation region under the KESS project revealed that
following each wind storm event during July–December
2004, the upper ocean inertial velocity shear tends toincrease in amplitude and to propagate downward to the
top layer of STMW (Luc Rainville, personal communication, 2010). Another possible cause for the discrepancy isinterannual variation of Kv in STMW resulting from thechanging density structure of STMW. Since thinner
STMW is formed and more strongly dissipated in 2006
than in 2004 (Sect. 4), STMW is less vertically homogeneous during 2006, which is likely to hinder the intensive
breaking of downward-propagating inertial waves at its
upper boundary. The third possible cause is the importanceof horizontal processes that were assumed to be negligiblein Qiu et al. (2006). After spring, STMW patches originating from different regions and having different densities
are stirred by mesoscale eddies and circulations andinterleave with each other, and this helps destroy the vertically uniform structure of STMW, particularly when the
KE is in an unstable state (Oka et al. 2011b; Sect. 4).
As noted in Sect. 2, part of STMW is transported
southward from its formation region, being subducted into
the permanent pycnocline. Subsequently, it is advectedsouthwestward, carrying temperature and salinity anomalies, and reaches the western boundary one to several years
later (Oka 2009). A portion of this ‘‘old’’ STMW is thentransported back to the formation region by the Kuroshio/
KE and is obducted onto the winter mixed layer, modulating the sea surface temperature there (Qiu and Huang
1995; Qiu 2002; Endoh et al. 2006; Liu and Hu 2007).
Besides this anticyclonic circulation of STMW, eastward
advection of STMW by the KE from the region southeastof Japan to the east of the dateline over a period of 1 year
was suggested from a lag correlation analysis of winter sea
surface temperature anomalies (Sugimoto and Hanawa
2005a).
The Kuroshio and KE influence not only STMW, butalso CMW and TRMW to the north. Saito et al. (2011)
analyzed quasi-Lagrangian observations by an isopycnal
Argo float to indicate that TRMW becomes warmer andsaltier after its formation in late winter, and is modified to

123

12

D-CMW within 1 year. They attributed this change partly
to double-diffusive salt-finger convection between TRMW
and the warmer, saltier, and equally dense surface water ofthe Kuroshio origin that is advected from the south. This
scenario is consistent with the observed facts that the upper
portion of TRMW is characterized by a high Turner angle
implying active salt-finger convection and by inhomogeneity in terms of temperature and salinity relative to thelower portion (Toyama and Suga 2010, 2011) and that the
proportion of TRMW (D-CMW) decreases (increases)
eastward/downstream in their formation region (Oka et al.
2011a; Saito et al. 2011). Thus, TRMW is likely a transientfeature in the TRMW/D-CMW formation region and canbe regarded as the most upstream portion of D-CMW.
After L-CMW and D-CMW are subducted from theeastern part of the formation regions (Oka et al. 2011a),they are advected anticyclonically in the subtropical gyre
toward the western boundary (Suga et al. 2004; Tsujino and
Yasuda 2004; Oka et al. 2011a). As they circulate southwestward in the southern part of the gyre, D-CMW takes amore zonal path than L-CMW because of the beta spiraleffect (Stommel and Schott 1977; Kubokawa 1999), whichmakes the two waters overlap each other (Fig. 2; Oka andSuga 2005; Kobashi et al. 2006; Xie et al. 2011). This stackof CMWs in the western to central part of the gyre, as well
as STMW in the western part, has been shown to generate
the Subtropical Countercurrent (Uda and Hasunuma 1969)
and the associated subtropical fronts (Kubokawa 1997,
1999; Kubokawa and Inui 1999; Aoki et al. 2002; Kobashi
et al. 2006; Yamanaka et al. 2008; Xie et al. 2011), whose
thermal effects on the atmosphere locally enhance precipitation and generate positive wind stress curl anomalies in
winter to spring (Kobashi et al. 2008; Xie et al. 2011). Thisis a new dynamic role of the mode waters in climate, which
contrasts with the traditional role as a heat reservoir
memorizing the wintertime ocean-atmosphere interaction(Xie et al. 2011; see also the review of Kobashi andKubokawa 2011 in this issue for details). Another interesting feature is that as L-CMW and D-CMW circulate
anticyclonically, their pathways shift to the inner side of
the subtropical gyre represented by geostrophic streamlines(Oka and Suga 2005). This is consistent with the modelingstudy of Nonaka and Xie (2000) that demonstrated thattemperature anomalies subducted from the northern part of
the gyre in association with the subduction of low PV water
circulate more zonally in the southern part of the gyre ashigh baroclinic mode Rossby waves than a passive tracer
subducted from the same location, reaching the westernboundary at a latitude higher than that of the North
Equatorial Current bifurcation (e.g., Qiu and Chen 2010a),
and are therefore unable to flow into the equatorial region.
ESTMW is also advected southwestward from the formation region along the outer path in the subtropical gyre

123

E. Oka, B. Qiu(Hautala and Roemmich 1998), but its low PV signal is
rapidly dissipated to the downstream (Ladd and Thompson
2001; Hosoda et al. 2001; Suga et al. 2004; Sugimoto andHanawa 2005b). This is because (1) ESTMW is muchthinner than the other mode waters in the North Pacific andis characterized by a weak PV minimum (Suga et al. 2004;
Sugimoto and Hanawa 2007); (2) owing to the subsurface
temperature/salinity stratification in the eastern part of thesubtropical gyre, salt-finger convection occurs vigorously
between ESTMW and the underlying permanent pycnocline water, which erodes ESTMW from its bottom after
spring (Sugimoto and Hanawa 2005b, 2007; Shimada et al.
2007; Toyama and Suga 2010, 2011); and (3) owing tostratus clouds over the formation region, ESTMW is not
capped by the strong seasonal pycnocline after spring,which allows atmospheric disturbances to continually
modify ESTMW from its top (Sugimoto and Hanawa
2007). Nevertheless, subducted temperature/salinity
anomalies in association with ESTMW might be transported to the downstream even after the low PV signal of
ESTMW is lost. A recent analysis of Argo float data
showed that isopycnal h–S (or spiciness) anomalies propagate in the range of rh = 25.0–25.5 kg m-3 from theeastern subtropical Pacific to the western tropical Pacific,
along a path and with a speed corresponding to the mean
geostrophic flows (Sasaki et al. 2010). This propagation
might be associated with the subduction and circulation of
ESTMW, and its influence on the equatorial region, which
has been hypothesized to be important for interdecadal
climate variability of the North Pacific (Hanawa 1996; Gu
and Philander 1997), should be examined quantitatively infuture studies.

4 Decadal variability
The ocean-atmosphere system in the mid-latitude North
Pacific exhibits significant decadal to interdecadal variability (e.g., Nitta and Yamada 1989; Tanimoto et al.
1993; Graham 1994; Trenberth and Hurrell 1994). Its stateis often described using the Pacific decadal oscillation(PDO) index, which is defined as the leading principal
component of North Pacific sea surface temperature variability poleward of 20°N (Hare 1996; Zhang 1996). A
positive (negative) PDO index indicates a stronger(weaker) Aleutian low, stronger (weaker) westerlies, and alower (higher) sea surface temperature except near the
North American coast. An abrupt change of sign in the
PDO index and the subsequent persistence for a few
decades, called a regime shift, occurred in 1925, 1947, and1977, exerting widespread impacts on the climate andmarine ecosystems in the North Pacific (Mantua et al.
1997; Mantua and Hare 2002).

Progress of North Pacific mode water researchThis observed long-term ocean-atmosphere variability is
widely believed to be controlled by the ocean, which has
much larger heat content and longer time scales than theatmosphere. As an important step for clarifying its mechanism, the response of the ocean interior to the decadal/
interdecadal surface forcing has been explored. Pioneering
works of Deser et al. (1996) and Schneider et al. (1999)
analyzed historical temperature data to demonstrate that
following the 1976/1977 regime shift when the PDO index
turned positive, cold anomalies were subducted from thesea surface in the central North Pacific and then propagateddownward in the permanent thermocline. Since the subduction area corresponded roughly to the central locationof decadal sea surface temperature variability (Tanimoto
et al. 1993) and also to the CMW subduction region, CMW
was expected to play a key role in transmitting the decadalatmospheric changes to the upper ocean in the mid-latitude
North Pacific (Suga et al. 1997). Following the data analysis of Yasuda and Hanawa (1997), numerical simulations
using various surface forcing were performed in the early
2000s to investigate decadal variability of CMW, withparticular attention to the 1976/1977 regime shift (Inuiet al. 1999; Xie et al. 2000; Kubokawa and Xie 2002; Ladd
and Thompson 2002; Hosoda et al. 2004). Their results,
with some differences among models, showed that changes
in the thermodynamic forcing (increases of upward surface
heat flux and southward Ekman transport) and the dynamic
forcing (increase of Ekman pumping) associated with theregime shift resulted in the formation of thicker and colder
CMW and the eastward shift of its circulation path in thesubtropical gyre, respectively, and their combined effects
altered the subsurface thermal structure of the gyre. An
analysis of the 180° repeat hydrographic section across the
CMW formation region consistently indicated that the
CMW temperature abruptly increased by 1°C from 1988 to
1989 when the PDO index transiently turned negative(Suga et al. 2003). Ladd and Thompson (2002) further
generalized that when the PDO index is positive, thicker
and colder CMW tends to be formed and the gyre circulation transporting it tends to be strengthened, and viceversa. Such a relation between the PDO index and themode waters with no time lag was also demonstrated byrecent modeling studies for CMW (Qu and Chen 2009) and
STMW (Davis et al. 2011).
On the other hand, recent observational studies proposeda new mechanism of decadal STMW variability, which
involves delayed oceanic response to the PDO-related
surface wind forcing. Specifically, Qiu and Chen (2005)
identified a decadal oscillation of the KE current system
between two dynamic states based on the satellite altimetersea surface height data during 1993–2004. This oscillation
originates from the large-scale wind stress curl forcing in
the central North Pacific around 160°W where the forcing

13

has its largest amplitude (Qiu 2003). When the PDO index
is positive (negative), negative (positive) sea surface heightand permanent thermocline depth anomalies are generated
there and propagate westward to the southeast of Japanover a period of *3 years at the speed of first-mode
baroclinic Rossby waves. As a result, the upstream KE jet
becomes unstable and weak (stable and strong), and isaccompanied by a weak (strong) southern recirculation
gyre and high (low) regional eddy activity, which is against
our expectation that a stronger mean flow favors baroclinic
instability and leads to higher eddy activity (Qiu and Chen
2005, 2010b, 2011; Kelly et al. 2007; Taguchi et al. 2007).
In the past 18 years, the KE was in the stable state in
1993–1994, 2002–2005, and 2010 and in the unstable statein 1995–2001 and 2006–2009.
The decadal KE variability is closely related to the
STMW formation in the recirculation gyre. An analysis oftemperature profiles in the KE region during 1993–2004
revealed that the STMW thickness also exhibited significant decadal variability and was large (small) when the KE
was in the stable (unstable) state (Qiu and Chen 2006),
which has been confirmed by several studies based on
observational data (Pan and Liu 2005; Qiu et al. 2007;
Sugimoto and Hanawa 2010; Oka et al. 2011b) andreanalysis product (Miyazawa et al. 2009). This occurs
because during an unstable KE period, active southwardtransport of high PV water from the Mixed Water region
across the KE due to the stronger eddy activity (Qiu andChen 2006; Qiu et al. 2007), as well as the higher background stratification associated with the shallower permanent pycnocline in the weaker recirculation gyre (Sugimotoand Hanawa 2010), is unfavorable for the development of
deep winter mixed layer, and the reverse is true during thestable KE period. For the period of 1993–2004, Qiu andChen (2006) also demonstrated that the STMW thickness
was highly correlated with the pre-existing upper-ocean
stratification and had little correlation with the wintertime
surface heat flux forcing. In other words, the thickness isdetermined by the dynamic forcing rather than the thermodynamic forcing, which contrasts with our traditionalview that the latter forcing is dominant. A subsequent
analysis of long-term temperature data during 1971–2007
showed that the contributions from the pre-existing stratification and the wintertime surface heat flux are comparable, with the former (latter) being more important after(before) around 1990 (Iwamaru et al. 2010). It has also
been pointed out that the pre-existing stratification isdetermined not only by the dynamic state of the KE current
system, but also by the summer surface heating, whose
intensity depends on the activity of tropical cyclones (Kakoand Kubota 2007; Tomita et al. 2010).
The decadal KE variability is important not only for theformation of STMW, but also for its subsequent evolution.

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14

Recent analyses of Argo float and high-resolution shipboard observation data (Oka 2009; Oka et al. 2011b)
showed that the STMW circulation in 2008 was much moreturbulent than that in 2006. In 2006, the STMW formed inseveral anticyclonic circulations in the recirculation gyretended to be continually trapped in its respective circulations after spring and to remain in the formation region
until late fall (Oka 2009). In 2008, in contrast, it tended togradually migrate southward. Simultaneously, STMWs
with different temperatures formed in different longitude
ranges were gradually stirred, then interleaved with eachother, and were finally vertically mixed to form a STMW
with an intermediate temperature (Oka et al. 2011b). The
difference in the STMW circulation between the 2 years is
attributable to the dynamic state of the KE current system.
The KE was in the unstable state in both 2006 and 2008,
but its path was more variable in 2008, which was the peakyear of the unstable period of 2006–2009. The associatedstronger eddy activity in the STMW formation region in2008 was hypothesized to enhance the eddy transportof STMW in both the meridional and zonal directions(Oka et al. 2011b).
Although the seasonal evolution of the STMW distributions in the stable KE period has not been investigated indetail (because of the lack of Argo floats in 2002–2005), it
is possible to put forth the following hypothesis about the
decadal variation in the STMW subduction and dissipation.
In the unstable KE period, the volume of STMW formed in
winter is smaller than in the stable KE period because theformation region is partially masked by cyclonic eddies(Rainville et al. 2007) and the STMW thickness is smaller(as mentioned above), but the volume of STMW subducted
into the south of the formation region is larger because of
the enhanced eddy transport. Also, STMW in the formation
region is dissipated after spring, mainly by vertical processes in the stable KE period and by horizontal processes
in the unstable period. This is because the thicker and more
uniform STMW in the stable period is more susceptible tothe intensive breaking of downward-propagating inertial
waves at its upper boundary (Sect. 3), while the stronger
eddy activity in the unstable period helps destroy the vertically uniform structure of STMW (Oka et al. 2011b).
Such hypothesized STMW variability would possibly
affect the climate and primary production through theimpact on the Subtropical Countercurrent (Yamanaka et al.
2008; Xie et al. 2011) and the nutrient supply from thesubsurface to the surface layer (Sukigara et al. 2011),
respectively.
The decadal variability of STMW temperature has also
been explored. Modeling studies have demonstrated that on
decadal time scales, stronger westerlies in the central North
Pacific lead to the formation of warmer STMW several
years later through an increase of the Kuroshio heat123

E. Oka, B. Qiu

transport (Ladd and Thompson 2002; Yasuda and Kitamura
2003; Lee 2009), as previously shown by observational
studies (Yasuda and Hanawa 1997; Hanawa and Kamada
2001). In addition, the Argo float data analysis (Oka 2009)
suggested that the zonal temperature change in the STMW
formed east of the Izu-Ogasawara Ridge (*140°E) issmall in the stable KE period and large in the unstable KEperiod. This might be explained as follows. In the stable
period, the KE takes a relatively straight path and isaccompanied by a single, contiguous southern recirculation
gyre, within which STMW having a zonally uniform
temperature is formed. In the unstable period, on the other
hand, the southern recirculation gyre is broken into several
anticyclonic circulations, within each of which STMW
with a different temperature is formed. In this latter case,the longer KE path results in a greater oceanic heat loss
along the path, leading to a gradual decrease of the STMW
temperature to the downstream. The float data analysis also
revealed a significant interannual change in the STMW
salinity (Oka 2009), which may be related to the variabilityof the North Pacific Tropical Water (Cannon 1966; Sugaet al. 2000). With the accumulation of additional Argo float
data in the coming years, we can expect the decadal variability in the STMW properties and its mechanisms to beclarified further.
The decadal variability of ESTMW has been least
explored among the three mode waters. A hindcast run for
1965–1993 (Ladd and Thompson 2002) showed that the
ESTMW density fluctuated on decadal time scales due
primarily to the buoyancy flux change, while a recent data
assimilation experiment for the 1990s (Toyoda et al. 2011)
demonstrated that the ESTMW formation rate variedinterannually owing to changes not only in the surface
cooling in winter and the pre-existing stratification controlled by the low-level cloud coverage in summer, but also
in the Ekman convergence of salt. These studies, however,pertain to the pre-Argo period during which we had inadequate salinity data. Given the important roles played bysalinity in the ESTMW formation and dissipation (Sects. 2
and 3), future analyses of Argo float data and the reanalysis
product assimilating them are expected to reveal new features and mechanisms of the ESTMW decadal variability.

5 Future directions
We have reviewed the progress in the observational andmodeling studies of the North Pacific mode waters from thepast decade. It is instructive to end this review by offeringsome directions for the mode water studies of the next
decade. The first target is the role of mesoscale eddies,
which prevail in the regions surrounding the KE jet(Mizuno and White 1983; Itoh and Yasuda 2010a, b; Itoh

Progress of North Pacific mode water research

et al. 2011), in the formation and subduction of L-CMW,
D-CMW, and TRMW. ‘‘Eddy-resolving’’ synoptic shipboard observations, such as the one conducted by Oka et al.
(2011b), in conjunction with eddy-resolving simulations
will contribute greatly to the refined depiction of the largescale CMW formation and subduction revealed by Argo
float data (Oka et al. 2011a).
Nearly 5 years have passed since the Argo float array in
the North Pacific was completed, and the collected data
have improved the climatology-based, classical views on
STMW and CMW. However, the formation, subduction,and circulation processes of ESTMW have not been fully
investigated using the float data. Future studies areencouraged to verify the various processes of ESTMW put
forth by the modeling studies in the early 2000s.
During the past decade, accumulation of satellite
altimeter sea surface height data has significantly enhanced
our understanding of the decadal variability of the KE
current system (Qiu and Chen 2005, 2010b, 2011). In asimilar manner, accumulation of Argo float data willundoubtedly unravel the decadal variability of the mode
waters in the coming decade. In Sect. 4, we have presented
our hypothesis on the decadal variability of STMW associated with that of the KE current system. CMW is also
expected to exhibit decadal variability owing to both thethermodynamic (surface heat flux and meridional Ekman
transport) and dynamic forcings. We anticipate two typesof dynamic forcing on CMW that originate from the PDOrelated surface wind forcing in the central North Pacific.
One is the decadal variability of eddy activity north of theKE, which would control the number and distribution of
anticyclonic eddies that are the probable formation sites ofthick CMW (Kouketsu et al. 2011). A recent study by Qiu
and Chen (2011) found that the decadal mesoscale eddy
variability had an opposite phase between the upstream(140°–152°E) and downstream (152°–165°E) KE regions.
These downstream decadally varying eddy signals canpotentially induce the long-term CMW variability. The
other is the observed (Qiu and Chen 2005) and simulated(Nonaka et al. 2006; Taguchi et al. 2007) decadal shifts ofthe thermohaline fronts bounding the CMW formation
regions. If these shifts are accompanied by changes in thewinter mixed layer properties in the interfrontal regions,the CMW density can fluctuate without a significant
change in the thermodynamic forcing (Suga et al. 2011). It
is of interest to examine the degree to which the large-scale
CMW subduction scheme for 2003–2008 (Oka et al.
2011a) applies to other periods. The maximum rh of CMW
that is subducted into the permanent pycnocline(26.4 kg m-3 for 2003–2008) can fluctuate on decadal timescales, and this may cause the decadal apparent oxygen
utilization changes observed in the subsurface centered at
rh = 26.6 kg m-3 in the subtropical and subarctic North

15

Pacific (e.g., Andreev and Kusakabe 2001; Ono et al. 2001;
Watanabe et al. 2001; Emerson et al. 2004; Mecking et al.
2008). Investigation into the decadal mode water variability and its mechanisms will also be facilitated by further development of numerical models in the coming
decade. Models that are able to reproduce the recently
identified structures and variability of currents, eddies, andmode waters are called for.
The decadal variability of STMW and CMW can
influence the sea surface temperature in their formation
regions through re-emergence (Hanawa and Sugimoto
2004; Sugimoto and Hanawa 2005a) or obduction after
circulating in the subtropical gyre (Qiu and Huang 1995;
Qiu 2002; Endoh et al. 2006). It can also potentially affectthe sea surface temperature in the southern region throughits impact on the Subtropical Countercurrent (Yamanaka
et al. 2008; Xie et al. 2011; Kobashi and Xie 2011; Nonaka
et al. 2011) and the Hawaiian Lee Countercurrent (Sasaki
et al. 2011). On the other hand, the decadal variability of
ESTMW can possibly modulate the sea surface temperature in the equatorial upwelling region. Quantification ofthese influences and clarification of the underlying mechanisms will lead to a better understanding of the ocean’srole in the decadal climate variability.
Multidecadal variability or longer-term changes of themode waters have been examined only recently. The
STMW temperature observed in the winter formation region
at 141°–150°E and in summer at 137°E exhibited anincreasing trend of 3–5 9 10-2°C/year-1 for the period of
1971–2007 (Iwamaru et al. 2010). Coupled climate model
simulations predicted that in response to global warming,the densities of the mode waters tended to become lower(Lee 2009; Luo et al. 2009), although their subduction rates
increased in Lee’s model and decreased in Luo et al.’smodel. It has also been predicted that banded structures ofthe sea surface temperature warming that were slanted in thenortheast-southwest direction appear in response to global
warming because the shallower winter mixed layer in the
Kuroshio-Oyashio Extension region results in weaker mode
water subduction, weaker Subtropical Countercurrent, andsmaller thermal advection from the southwest (Xie et al.
2010, 2011; Xu et al. 2011). Such long-term mode waterchanges and their consequences are expected to become
increasingly important not only from the physical, but also
the biogeochemical, perspectives in the coming decades.
Acknowledgments The authors thank Atsushi Kubokawa for theopportunity to write this review and two anonymous reviewers forhelpful comments on the manuscript. This review is based on discussions by the authors in summer 2010 when EO visited the University of Hawaii at Manoa for 1 month under the Overseas Internship
Program for Outstanding Young Earth and Planetary Researchers
provided by the Department of Earth and Planetary Science, theUniversity of Tokyo. EO is also supported by the Japan Society for

123

16
Promotion of Science [KAKENHI, Grant-in-Aid for Scientific
Research (B), no. 21340133] and the Ministry of Education, Culture,
Sports, Science and Technology, Japan (MEXT; Grant-in-Aid for
Scientific Research on Innovative Areas under grant no. 22106007).
BQ is supported by NSF through grant OCE-0926594 and NASAthrough contract 1207881 as part of NASA’s Ocean Surface Topography Mission.

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