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Formation and Subduction of Central Mode Water Based on Profiling
Float Data, 2003–08
EITAROU OKA,*,1 SHINYA KOUKETSU,1 KATSUYA TOYAMA,# KAZUYUKI UEHARA,@TAIYO KOBAYASHI,1 SHIGEKI HOSODA,1 AND TOSHIO SUGA1,#
1* Atmosphere and Ocean Research Institute, The University of Tokyo, Kashiwa, Japan
Research Institute for Global Change, Japan Agency for Marine-Earth Science and Technology, Yokosuka, Japan#
Department of Geophysics, Graduate School of Science, Tohoku University, Sendai, Japan@
Department of Marine Science, School of Marine Science and Technology, Tokai University, Shizuoka, Japan(Manuscript received 24 December 2009, in final form 18 August 2010)
ABSTRACT
Temperature and salinity data from Argo profiling floats in the North Pacific during 2003–08 have been
analyzed to study the structure of winter mixed layer north of the Kuroshio Extension and the subsurface
potential vorticity distribution in the subtropical gyre in relation to the formation and subduction of thecentral mode water (CMW). In late winter, two zonally elongated bands of deep mixed layer extend at 338–398
and 398–438N, from the east coast of Japan to 1608W. These correspond to the formation region of the lighter
variety of CMW (L-CMW) and that of the denser variety of CMW (D-CMW) and the recently identified
transition region mode water (TRMW), respectively. In the western part of the L-CMW and D-CMW–
TRMW formation regions west of 1708E, the winter mixed layer becomes deeper and lighter to the east (i.e.,to the downstream). As a result, the formed mode water is reentrained into the mixed layer in the farther eastin the following winter and modified to the lighter water and is thus unable to be subducted to the permanent
pycnocline. In the eastern part of the formation regions between 1708E and 1608W, on the other hand, thewinter mixed layer becomes shallower and lighter to the east. From these areas, the L-CMW with potential
density of 25.7–26.2 kg m23 and the D-CMW–TRMW (mostly the former) of 26.1–26.4 kg m23 are subducted to the permanent pycnocline, and they are then advected anticyclonically in the subtropical gyre.
These results imply that during the analysis period large-scale subduction to the permanent pycnocline occurs
in the density range up to 26.4 kg m23 in the open North Pacific, whereas the winter mixed layer density
reaches the maximum of 26.6 kg m23. This is supported by the vertical distribution of apparent oxygen utilization in a hydrographic section in the subtropical gyre.

1. Introduction
Central mode water (CMW) is characterized by a pycnostad lying in the lower ventilated pycnocline in the North
Pacific subtropical gyre. Motivated by Talley’s (1988)
potential vorticity (Q) map on the su 5 26.2 kg m23 isopycnal (su is potential density) indicating a horizontal Q
minimum in the central part of the gyre, Nakamura (1996)
and Suga et al. (1997) separately identified a thermostad of
8.58–11.58 and 108–138C using different climatological data
and termed it CMW. Recent isopycnal Q maps of Sugaet al. (2004), which are based on the isopycnally averaged

Corresponding author address: Eitarou Oka, Atmosphere andOcean Research Institute, The University of Tokyo, Kashiwa,
Chiba 277-8564, Japan.
E-mail: eoka@aori.u-tokyo.ac.jp
DOI: 10.1175/2010JPO4419.1
Ó 2011 American Meteorological Society

North Pacific HydroBase climatology (Macdonald et al.
2001), delineate that the low-Q signature of CMW extends
anticyclonically from the northern to western boundary of
the subtropical gyre on isopycnals between su 5 25.9 and26.5 kg m23.
The distribution and properties of CMW are observed
to change on decadal and interdecadal time scales (Yasuda
and Hanawa 1997; Suga et al. 2003). Therefore, the CMW
circulation in the subtropical gyre can cause the observed
and modeled propagation of temperature anomalies in thepermanent pycnocline (e.g., Deser et al. 1996; Schneider
et al. 1999; Nonaka et al. 2000), which might be importantfor decadal–interdecadal climate variability in the Pacificregion (Latif and Barnett 1994; Gu and Philander 1997).
CMW also possibly affect the long-term climate variability by transporting anthropogenic carbon both downward
and meridionally because the formation region of CMW,114

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described in detail below, is an intense sink for atmospheric CO2 (e.g., Takahashi et al. 2009).
CMW is formed in the deep winter mixed layer1 northof the Kuroshio Extension (Ladd and Thompson 2000;
Suga et al. 2004; Oka et al. 2007; Ohno et al. 2009). An
analysis of data from a repeat hydrographic section along
1658E during 1991–2003 (Oka and Suga 2005) demonstrated that a lighter (denser) variety of CMW with su 5
25.8–26.2 (26.3–26.4) kg m23 is formed between the Kuroshio Extension and its northern bifurcation (the northernbifurcation and the subarctic front2). The differential formation of CMW was also observed in a hydrographic
section along 1798E (Mecking and Warner 2001) and was
reproduced in a high-resolution ocean general circulation
model (Tsujino and Yasuda 2004).
The CMW formation region was initially inferred toextend zonally from 1758E to 1608W based on Suga et al.’s(1997) temperature climatology, but more recent data
have demonstrated that the formation region extends
much farther westward. The ever presence of CMW in the
1658E repeat section and the surrounding flow field (generally eastward) imply that the formation region extends atleast as far west as 1558E (Oka and Suga 2005). Furthermore, upgraded winter mixed layer climatology of Sugaet al. (2004) demonstrated that the probable formation
region of CMW, which is defined as the winter mixed layer
having the same potential temperature and salinity as the
CMW pycnostads in the permanent pycnocline, extends as
far west as 1438E, almost reaching the east coast of Japan.
How is CMW subducted from this zonally elongatedformation region into the permanent pycnocline? The
mechanism is important in determining the CMW properties and causing their variability but has not been wellclarified. Numerical models demonstrated that CMW is
subducted east of the date line by crossing a sharp front of
the mixed layer depth (MLD) at the eastern end of theformation region (e.g., Kubokawa and Inui 1999; Xie et al.
2000; Hosoda et al. 2004; Tsujino and Yasuda 2004). On
1
In this study, the term mode water indicates a pycnostad lying
in the subsurface during the warm season as well as that outcropping to the sea surface in late winter (i.e., deep winter mixed
layer). The formation of such deep winter mixed layer is referred toas the formation of a mode water. The subduction of a mode watermeans a situation in which a mode water or part of it is advected
into the permanent pycnocline during the warm season and is notreentrained into the mixed layer in the subsequent winters. The
subduction occurs in the whole or part of the formation region of
the mode water.
2
In this study, the term subarctic front is used for a densitycompensating front between the warmer, saltier water in the subtropical gyre and the colder, fresher water in the subarctic gyre
(Roden 1970, 1972; Zhang and Hanawa 1993; Yuan and Talley
1996). Some studies refer to this front as the subarctic boundary(e.g., Favorite et al. 1976; Yasuda 2003).

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the other hand, climatological data indicated that the water is mainly subducted from a wide area in the formation
region where the winter MLD gradually decreases eastward while the mixed layer density changes little (Sugaet al. 2004, 2008). Ladd and Thompson (2000) and Mecking
and Warner (2001) speculated that the CMW formed in thewestern part of the formation region is not subducted into
the permanent pycnocline but reentrained into the mixed
layer in the following winters, thereby preconditioning the
subduction in the eastern part of the formation region. Thatis, the eastward subsurface advection of low Q from thewestern part and the resultant weak density stratification
below the seasonal pycnocline in the eastern part enable
the winter mixed layer there to deepen to a considerable
depth in spite of relatively weak surface cooling.
The CMW subduction mechanism so far has been investigated using climatological data by most of the previous
studies, but these analyses had intrinsic difficulties (Sugaet al. 2004). If we suppose a synoptic Q map on an isopycnal
in the CMW density range, the lowest Q of CMW immediately after subduction should be located just south or eastof the winter outcrop of the isopycnal, neighboring high Q
in the seasonal pycnocline on the other side of the outcrop.
The smoothing process of climatology inevitably averages
these low and high Q across the winter outcrop, thus considerably weakening the low-Q signature of CMW. In fact,
Suga et al. (2004) demonstrated that the most upstream
portion of CMW on each isopycnal is not directly connected to the winter outcrop, even though they employed as
small smoothing scale as possible for their climatology. The
low Q of the CMW in the formation region should also bevulnerable to the smoothing process because of its meridional extent of only a few degrees in latitude (Oka and Suga
2005). In addition, the winter MLD in the formation region
tends to be overestimated because the smoothing process
mixes different water types across the subarctic front,
causing artificial cabbeling in the surface layer (Ohno et al.
2004; Suga et al. 2004). It is thus suggested that we need toreexamine the CMW subduction mechanism using in situ
observation data unaffected by the smoothing process.
Since 2000, profiling floats have been deployed in
the global ocean under the international Argo program(Roemmich et al. 2001). The unprecedented array of 3000
floats, which has a horizontal resolution of 38 in latitude
and longitude, was completed in 2007 (Roemmich et al.
2009) and is providing temperature and salinity data down
to 2000-dbar depth in near–real time with relatively small
regional biases and almost no seasonal bias. In this study,
we analyze the float data in the North Pacific during 2003–
08 to investigate the detailed structure of winter mixed
layer north of the Kuroshio Extension and the subsurface
Q distribution in the subtropical gyre in relation to the
CMW formation and subduction. By using the unsmoothed

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FIG. 1. Distribution of MLD in February–April of 2003–08.

float data, we will demonstrate a direct connection between the deep winter mixed layer and the subsurface low
Q, clarifying the location and mechanism of the CMW
subduction. Note that in this study we will not consider
interannual variation of the CMW formation–subduction
during the analysis period. It would be ideal to do thepresent analysis using float data from a single year, but
their spatial density is not enough to resolve the detailed
structure of winter mixed layer in the CMW formation
region (Bingham and Suga 2006).
The reminder of this paper is organized as follows: Thedata and their processing are explained in section 2. In
section 3, the winter mixed layer structure around the
CMW formation region is investigated using late winterfloat data, from which the CMW subduction mechanism isinferred. The inferred mechanism is verified in section 4,
by examining the CMW circulation in the subtropical gyreusing float data during the warm season. Summary anddiscussion are given in section 5.

2. Data and processing
We used temperature (T ) and salinity (S) data obtainedby Argo profiling floats in the North Pacific in 2003–08.
These floats drift freely at a predetermined parking pressure (typically 1000 dbar), and conduct T–S measurements
between near the sea surface (4 dbar) and an intermediate
depth (2000 dbar) at a predetermined interval (10 days).
The collected data at about 70–120 sampling pressures,with a typical interval of 5 dbar at depths shallower than
200 dbar, 10–25 dbar at 200–1000 dbar, and 50–100 dbar atdeeper than 1000 dbar, are transmitted from the surfaced
floats to satellites and are made freely available within
24 h, after passing through the Argo real-time quality
control (Argo Data Management Team 2009).
We downloaded the real-time quality controlled float
data from the ftp site of Argo Global Data AssemblyCenter (available online at ftp://usgodae1.fnmoc.navy.
mil/pub/outgoing/argo; ftp://ftp.ifremer.fr/ifremer/argo)and eliminated defective T–S profiles, such as those withmeasurements flagged as bad and those lacking intermediate layers for certain depths, following the procedures of Oka et al. (2007). Then we interpolated each
T–S profile on a one-dbar grid using the Akima spline(Akima 1970) and calculated potential temperature (u),su, and Q. Here, Q is defined as
Q5

f ›su
,
r ›z

(1)

where f is the Coriolis parameter, r is in situ water density, and z is the vertical coordinate (positive upward). In
Eq. (1), relative vorticity is neglected because it is comparable to planetary vorticity only in the vicinity of theKuroshio Extension and other strong currents’ axis (e.g.,
Qiu et al. 2006), whereas deep winter mixed layers are
typically formed between such currents.
MLD is defined as the shallower value of the depth atwhich su increases by 0.03 kg m23 from the sea surface andat which u decreases or increases by 0.28C from the surface,following the definition of Oka et al. (2007). Values of u,S, and su at 10 dbar were used as the surface values andwere considered to represent those of the mixed layer.

3. Structure of winter mixed layer in the CMW
formation region
Using the 6-yr float data, an MLD map in the North
Pacific is constructed for three late winter months of
February–April, during which MLDs in the CMW formation region reach their maximum (Oka et al. 2007;
Fig. 1). The MLD map indicates three regions exceeding
150 dbar, as demonstrated in previous studies (e.g., Karaet al. 2000; Suga et al. 2004; Ohno et al. 2009). The largest

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FIG. 2. (a) The u–S relationship for the mixed layer deeper than 150 dbar in the western and central midlatitude
North Pacific in February–April of 2003–08. Symbol types indicate MLD. (b) Number of points in (a), classified foreach 0.05 (S) 3 0.58 (u) bin. Thick black lines indicate properties of mode waters. In both (a) and (b), solid and dashedcontours denote su (kg m23).

one with the deepest mixed layers spreads widely in thewestern and central part of the midlatitude North Pacific.
The second one is located in the deep basin of the Bering
Sea and south of the Aleutian Islands, where the lower
part of the winter mixed layer remains in the subsurface
as the dichothermal water in the following seasons (e.g.,
Dodimead et al. 1963; Miura et al. 2002; Ueno et al. 2005).
The smallest one with the shallowest mixed layers among
the three regions lies in the eastern North Pacific Transition Region Mode Water near 308N, 1408W, which corresponds to the formation region of the North PacificEastern Subtropical Mode Water (Hautala and Roemmich
1998). Because CMW is formed in the largest region in thewestern and central midlatitude North Pacific, we willconcentrate on the winter mixed layer there hereafter.
The deep winter mixed layer in the region exhibits anear-linear, continuous u–S relation between 58 and 208C
(Fig. 2a). In this diagram, outcropping mode waters,
namely the deep winter mixed layer having common u–S
and a large horizontal extent, are expected to appear as
concentrations of points because of relatively uniform
distribution of the profiling float observations (Fig. 1). To
detect mode waters, the number of points in Fig. 2a
are counted for each bin of u 3 S (Fig. 2b) and u and S(Figs. 3a,b). A pronounced maximum is found aroundu 5 18.08C and S 5 34.80, which are the typical propertiesof the North Pacific Subtropical Mode Water (STMW;
Masuzawa 1969).

On the colder and fresher side of the maximum for
STMW, there seem to be three smaller maxima around u 5
13.08, 9.08, and 6.58C and S 5 34.45, 34.05, and 33.80 in eachof Figs. 2b and 3a,b. When we compose groups of adjacentu–S bins with three or more float observations, as indicatedby thick black lines in Fig. 2b, the properties of each groupagree with the previously reported properties of a mode
water. That is, the warmest group (u 5 16.08–21.58C, S 5
34.65–34.95, and su 5 24.2–25.6 kg m23) corresponds to
STMW; the second and third groups (10.58–16.08C, 34.30–
34.65, and 25.4–26.3 kg m23; 8.08–10.58C, 33.80–34.20,
and 26.1–26.5 kg m23) correspond to the lighter CMW(L-CMW) and denser CMW (D-CMW) observed in the
1658E repeat section (Oka and Suga 2005); and the coldest
group (5.08–7.58C, 33.60–33.90, and 26.4–26.6 kg m23)
corresponds to the recently identified transition region
mode water (TRMW; Saito et al. 2007), which is formed
and distributed in the transition region north of the subarctic front in the western North Pacific. Consequently, wedefine the four mode waters in their formation regions
using these four groups of u–S bins.
Although the four mode waters have distinct u and S,
the su of L-CMW, D-CMW, and TRMW is relatively
close to each other, particularly between the latter two. As
a result, number of float observations for each su bin exhibits only two peaks: one peak at 24.8–25.5 kg m23 for
STMW and the other peak at 25.8–26.6 kg m23 for theremaining three mode waters (Fig. 3c). This implies that

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FIG. 3. Number of points in Fig. 2a, classified for each (a) 0.58 bin of u, (b) 0.05 bin of S, and (c) 0.1 kg m23
bin of su. Note that su increases downward on the ordinate in (c).

the three mode waters cannot be distinguished fromeach other by means of su only. It should be also mentioned that no deep winter mixed layer is observed on
the denser side of the 25.8–26.6 kg m23 peak; 26.6 kg m23
is the densest su class in the western and central midlatitude North Pacific, as demonstrated in previous studies (e.g., Talley 1993).
The formation regions of the four mode waters are
mapped in Fig. 4. The STMW formation region lying in therecirculation gyre south of the Kuroshio and the Kuroshio
Extension is located between 268 and 368N, with a somewhat larger horizontal extent than demonstrated in previous studies (e.g., Oka and Suga 2003; Oka 2009) becauseof the wider u–S range used for the STMW definition inthis study. To the north of the Kuroshio Extension, the
L-CMW formation region extends zonally between 1438E
and 1638W at 338–398N, shifting gradually southward toward the east. In the farther north, the formation regions of
D-CMW and TRMW mingle with each other at 398–438N,
also extending zonally from 1428E to 1608W. In this latitude band, both D-CMW and TRMW are formed between
1488 and 1788E, whereas only TRMW (D-CMW)
is formed west (east) of it. Thus, the two zonal bands of the
L-CMW and D-CMW–TRMW formation regions extend from the east coast of Japan to 1608W for roughly
5000 km.
The geopotential anomaly distribution in Fig. 4 denotesthat the subsurface geostrophic flow in these regions isgenerally eastward with a speed of ;3 cm s21, which isequivalent to ;108 in longitude per year. Therefore, ineach of the L-CMW and D-CMW–TRMW formation

regions, the mode water is expected to be advected eastward, spending about 5 yr from the western to the easternend of the region. During this journey, the propertiesof mode water are likely to be altered by various processes, such as the surface heat and freshwater fluxes,
detrainment–entrainment to and from the underlying
layer, horizontal advection, mixing, etc., not only in late
winter when deep mixed layer is formed but also in theother seasons.
In fact, the properties of each mode water change considerably eastward within the formation region (Fig. 5).
The u of L-CMW increases from ;118 to ;148C, and its S
slightly increases from ;34.4 to ;34.5, in contrast with theu and S of STMW decreasing eastward. The u and S increases of L-CMW are believed to be caused primarily by
the northward component of geostrophic flow (this is indicated by a few streamlines joining the L-CMW formation region from the south; Fig. 4) that brings warmer andsaltier water to the region, because other processes—suchas the surface heat and evaporation minus precipitation
fluxes that are both negative on the annual average in mostpart of the region (Fig. 6), the entrainment from the colder
and fresher underlying layer, and the southward Ekman
transport due to the prevailing westerlies—are all expected to decrease the u and S of L-CMW. As a result of
the eastward u increase, the su of L-CMW in the formation region decreases from ;26.3 to ;25.7 kg m23(Fig. 5c).
The u of D-CMW in the formation region increases
eastward from ;88 to ;108C, whereas its S decreases
from ;34.1 to ;33.9 (Figs. 5a,b). TRMW is much colder

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FIG. 4. Distributions of the mixed layer deeper than 150 dbar in the western and central midlatitude
North Pacific with the properties of STMW (white circles), L-CMW (black squares), D-CMW (white
triangles), and TRMW (black stars) and without the mode water properties (dots) in February–April of
2003–08. Contours indicate geopotential anomaly (m2 s22) at 150 dbar relative to 950 dbar, calculatedusing the Argo profiling float data during 2003–08 and an optimal interpolation technique (Bretherton
et al. 1976).

(;68C) and fresher (;33.7) than D-CMW at the westernend of the formation region, but its u and S gradually
increase and approach those of D-CMW toward the eastand seem to merge into them near 1808. This structure isconsistent with a recent analysis of Argo float data during
2001–05 (Saito 2006), which demonstrated that TRMW is
modified to warmer and saltier D-CMW within approximately a year from its formation in winter. This occurs
because after spring geostrophic flow with a vertical shear
from the south brings the warmer, saltier, and equallydense water to the surface layer of the TRMW formation
region, leading to double-diffusive salt-finger convection
between the surface water and the underlying TRMW
that increases the u and S of TRMW (Saito 2006). Suchmodification would provide a precondition favorable to theformation of more D-CMW than TRMW to the downstream, possibly leading to the termination of TRMW
formation near the date line.
The downstream u increase of D-CMW and TRMW is
also likely to be due to the northward geostrophic advection into the formation region of these waters (Fig. 4),
whereas the S decrease of D-CMW cannot be explainedby the same mechanism. In the D-CMW formation region,
the salinity input by the geostrophic advection is probably
exceeded by the surface freshwater flux that is muchstronger than in the L-CMW formation region (Fig. 6b)
and the southward Ekman transport that brings much
fresher water of the transition region. The TRMW to
D-CMW modification explained above would also dilute
D-CMW.
The su of D-CMW in the formation region decreases
eastward from ;26.5 to ;26.1 kg m23 because of the u
increase and S decrease (Fig. 5). TRMW is slightly denser
than D-CMW, and it also decreases its su from ;26.6 to;26.4 kg m23 because of the u increase. As a result of thelarge zonal su changes of L-CMW and D-CMW, the
D-CMW in the easternmost part of the formation region
is comparably as dense as or even lighter than the
L-CMW in the westernmost part of the formation region, although both u and S are significantly different
between these waters. Such property distribution is reflected in the single peak for L-CMW, D-CMW, andTRMW in the su-based observation number histogram(Fig. 3c) and the three individual peaks for the threewaters in the u- and S-based histograms (Figs. 3a,b).
The late winter MLD also exhibits significant zonal
variation within each mode water formation region (Fig. 7).
The MLD in the STMW formation region is at a maximum between 1408 and 1508E, where the southern recirculation of the Kuroshio Extension is strongest (e.g.,
Qiu and Chen 2005). The MLD in the L-CMW formation
region is largest farther east near 1708E, decreasing to both
the west and the east, except for some large values west of
1558E that might be related to anticyclonic mesoscaleeddies. This means that the MLD in the L-CMW formation region increases to the downstream in the westernpart of the region and decreases in the eastern part, possibly because of intense winter cooling in the western partand milder cooling in the eastern part (Fig. 6a). The MLD
in the D-CMW–TRMW formation region presents a similar structure to the L-CMW formation region; it slightly
increases to the downstream in the western part andgradually decreases in the eastern part, reaching the maximum near 1708E (Fig. 7c). In both the L-CMW and
D-CMW–TRMW formation regions, the eastward MLD
decrease in the eastern part is gradual without any sharp
MLD front, as in the climatological MLD field (Sugaet al. 2004, 2008).

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FIG. 5. Plots of (a) u, (b) S, and (c) su against longitude for the mixed layer deeper than 150 dbar in thewestern and central midlatitude North Pacific in February–April of 2003–08. The symbols follow Fig. 4.
Note that su increases downward on the ordinate in (c).

What subduction scheme can be inferred for L-CMW
and D-CMW–TRMW from the observed mixed layer
and current structures in their formation regions (Figs. 4,5, 7)? In the western part of the formation regions westof 1708E where the winter mixed layer becomes deeper
and lighter to the downstream, it is expected that thelower portion of the winter mixed layer remains in thesubsurface as low-Q water in spring to summer, and it isthen reentrained into the mixed layer in the following
winter about 108 in longitude to the east. On the other
hand, in the eastern part of the formation regions east of
1708E where the winter mixed layer becomes shallower
and lighter to the downstream, the lower portion of thewinter mixed layer is unlikely to be reentrained into themixed layer in the farther east in the following winters,
thereby being subducted into the permanent pycnocline.
If this occurs, the mixed layer density distribution in Fig.
5c implies that the L-CMW with su 5 25.7–26.2 kg m23
and the D-CMW–TRMW with 26.1–26.4 kg m23 are
subducted from the eastern part of the formation regions between 1708E and 1608W, whereas the L-CMW
with 26.3 kg m23 and the D-CMW–TRMW with 26.5–
26.6 kg m23 that exist only in the western part of theformation region are not subducted to the permanent
pycnocline. In the next section, we will demonstrate that
this subduction scheme is basically correct, by examining
the subsurface CMW circulation in the subtropical gyre.
The above scheme implies that in the western part of theformation region the CMW left in the subsurface during
the warm season is reentrained into the mixed layer having
slightly lower density in the following winter. This scenario
might seem surprising, but it is confirmed by Lagrangian
observations performed by several Argo profiling floats.
One example is obtained by a float that drifts eastward

around 418N in the D-CMW–TRMW formation region at
a speed3 of about 108 in longitude per year (Fig. 8).
In summer 2007, low Q (,1.5 3 10210 m21 s21)
of D-CMW with su 5 26.5 kg m23, which is presumably formed in late winter 2007 in the farther west, isseen at depths of 150–250 dbar below high Q (.5 3
10 210 m21 s21) of the seasonal pycnocline. In October,the depth and su of the mixed layer start increasing andfinally reach the maximum of 292 dbar and 26.43 kg m23
on 10 March 2008. During this period, the depth and su of
the seasonal pycnocline beneath the mixed layer also increase while its Q decreases. In association with these
deepenings, the low Q of D-CMW gradually weakens at
su 5 26.5 kg m23, presumably because of reentrainment to
the overlying seasonal pycnocline enhanced by mechanicalmixing, and it is finally replaced by relatively high Q (.3 310210 m21 s21) just beneath the developed mixed layer
in February. After the mixed layer shallows in March andApril, low Q of newly formed D-CMW is left in the subsurface at su 5 26.4 kg m23, while Q at su 5 26.5 kg m23
is not low anymore. Thus, through the winter season, the
D-CMW lying in the subsurface in the western part of theformation region is reentrained into the slightly lighter
winter mixed layer, which is then left in the subsurface as
newly formed D-CMW.

4. CMW circulation in the subtropical gyre
In most of the previous studies, the CMW circulation in
the subtropical gyre has been examined using climatological3This speed happens to coincide with the eastward advection
speed at 150 dbar in the CMW formation region estimated from
Fig. 4, although the float drifts at a parking pressure of 1000 dbar.

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FIG. 6. Annual-mean (a) net surface heat flux (W m22) and (b) evaporation minus precipitation flux (cm yr21), calculated using monthly climatologies of surface heat and freshwater
fluxes and wind stress from da Silva et al. (1994). The fluxes are positive when the ocean gains
heat and loses freshwater. Negative values are shaded. Dots (circles) denote the formation
regions of L-CMW (D-CMW–TRMW).

data, which do not properly represent the CMW structurearound the formation region (section 1). The alternativeuse of in situ observation data from Argo profiling floats is
expected to provide a better view, but simply averaging Q
values from the float data at each location could also mar
the low-Q structure of CMW around the formation region,
as in the climatological cases. We therefore calculated thefrequency of low Q at each location on the isopycnals
between su 5 25.7 and 26.6 kg m23 for the warm seasonof 2003–08 (Fig. 9), based on which the CMW subduction
scheme proposed in the previous section is verified. It
should be noted that the following results do not dependon the choice of critical Q value for each isopycnal.
On the isopycnal of su 5 26.6 kg m23, the frequency of
low Q is high around 408–428N, 1528–1638E, which lies inthe D-CMW–TRMW formation region (Fig. 4). This low-Q
patch is considered to represent TRMW that is formed at
1458–1608E in winter (Fig. 5c) and then advected eastward
to the observed location. The low-Q patch, however, does
not extend to the downstream anymore; the frequency
drops to less than 10% east of 1708E (Fig. 9). This discontinuity implies that the TRMW lying at 1528–1638E in
spring to summer is reentrained into the mixed layer in thefollowing winter in the farther east and modified to lighter
TRMW or D-CMW, as demonstrated by the Lagrangian

float observations (Fig. 8). The same applies to D-CMW–
TRMW with su 5 26.5 kg m23, which is formed at 1458–
1658E in winter (Fig. 5c) and advected to 1568–1748E in
spring to summer but disappears in the farther east (Fig. 9).
Thus, the D-CMW–TRMW with su 5 26.5–26.6 kg m23
is not subducted to the permanent pycnocline, in spite ofits abundance in the formation region (Fig. 3c).
The low-Q frequency on su 5 26.4 kg m23 exhibits aclearly different distribution from the denser isopycnals;
the low Q extends anticyclonically from 418N, 1708E to
188N, 1358E while decreasing its intensity (Fig. 9). This
indicates that D-CMW–TRMW with su 5 26.4 kg m23 ispermanently subducted east of 1708E, and subsequently
advected in the subtropical gyre toward the westernboundary. On su 5 26.3 kg m23, the low Q of D-CMW
also extends anticyclonically from 408N, 1808 to thedownstream. However, low Q is hardly seen in the upstream region west of 1708E, although the L-CMW with
26.3 kg m23 is formed at 1438–1558E in winter (Fig. 5c).
Because the L-CMW in the westernmost part of the formation region is relatively thin (Fig. 7b), its low Q might bedissipated quickly after spring.
On su 5 26.2 kg m23, low Q departs from 418N,
1728W in the D-CMW–TRMW formation region andfrom 378N, 1768E in the L-CMW formation region, both

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121

FIG. 7. Plots of MLD against longitude for the mixed layer in the western and central midlatitude North Pacific with the properties of (a) STMW, (b) L-CMW, and (c) D-CMW–TRMW(defined in Fig. 2b) in February–April of 2003–08.

extending anticyclonically to the downstream. This isopycnal is ventilated by both varieties of CMW, which havethe same density but significantly different u–S (Fig. 5).
The isopycnal of su 5 26.1 kg m23 is also ventilated byboth D-CMW and L-CMW, whose low Q extends from
388N, 1688W and 368N, 1788W, respectively (Fig. 9). On
the lighter isopycnals of su 5 25.7–26.0 kg m23, low Q
arises only from the L-CMW formation region at 338–
368N, at a farther downstream location with decreasing su.
These features of the CMW circulation are summarized
as follows. D-CMW–TRMW (L-CMW) with su 5 26.1–
26.4 (25.7–26.2) kg m23 is subducted to the permanent
pycnocline around 408–428N (338–378N) east of 1708E andthen extends anticyclonically in the subtropical gyre. The
subduction point for each variety of CMW shifts eastward
with decreasing su. On the other hand, D-CMW–TRMW
with su 5 26.5–26.6 kg m23, whose low Q is confined toa relatively small area in the western part of the formation
region, is not subducted to the permanent pycnocline but
reentrained into the winter mixed layer and modified to

the lighter water. The low Q on su 5 26.5–26.6 kg m23 in
the downstream region, which reappears near 308N, 1608W
and extends southwestward, is possibly supplied from the
CMW on the lighter isopycnals through vertical diffusion.
Thus, the subsurface CMW circulation represented by the
low-Q frequency agrees well with that expected from thewinter mixed layer and current structures. The most noticeable feature is the discontinuity of low Q on su 5 26.5–
26.6 kg m23, which is not seen in the climatological Q
maps (e.g., Suga et al. 2004). It should be emphasized that
the CMW subduction actually occurs only in the easternpart of the formation region east of 1708E; the western partof the formation region preconditions the subduction in
the eastern part, as speculated by Ladd and Thompson(2000) and Mecking and Warner (2001).
The confirmed CMW subduction scheme implies thatduring 2003–08 large-scale subduction to the permanent
pycnocline occurs up to 26.4 kg m23 in the open North
Pacific, whereas the winter mixed layer density reaches
the maximum of 26.6 kg m23. This should be reflected in122

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FIG. 8. (a) Time series of longitude (black curve) and latitude (red curve) obtained by an Argo profiling
float with World Meteorological Organization identifier (WMO ID) 2900691 from 3 Aug 2007 through26 Oct 2008. Time–pressure sections of (b) su (kg m23) and (c) Q (10210 m21 s21) and (d) time–su
section of Q (10210 m21 s21) obtained by the same float as in (a). Contours are drawn at 0.5, 1, 1.5, 2, 3, 4,
and 5 3 10210 m21 s21 in (c) and (d). White curves in (b) and (c) denote MLD. Tick marks on the top ofeach panel indicate the float observations.

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the distribution of chemical properties in the subtropicalgyre. As the most comprehensive example, we here examine the vertical structure of apparent oxygen utilization (AOU) in the CMW distribution region, using
conductivity–temperature–depth–oxygen (CTDO2) datafrom the World Ocean Circulation Experiment Hydrographic Program (WHP) P02 section along 308N in June–
August 2004, downloaded from the WHP office Web site(see online at http://whpo.ucsd.edu). The P02 section exhibits two significant low-Q areas of mode waters embedded in high Q of the seasonal and permanent pycnoclines:
one of STMW at su 5 ;25.4 kg m23 west of 1808 and theother of CMW at 25.7–26.5 kg m23 east of 1758E (Fig. 10).
The AOU in the STMW is ;0.5 ml l21, which is low
enough to indicate that the observed STMW is formed in
the late winter of 2004 (Suga et al. 1989; Oka and Suga
2005).
The AOU in the CMW is also expected to be significantly lower than the underlying layers. To examine thisrelation, both Q and AOU were averaged at 33 CTDO2
stations between 1758 and 1558W where thick CMW isobserved (Fig. 11). The average Q increases almost linearly upward at su . 26.6 kg m23 in the unventilated layer,
and it shows negative anomalies at 25.7–26.6 kg m23 in
the CMW density range. The average AOU decreases
upward above the AOU maximum at 27.3 kg m23. In theunventilated layer, ›AOU/›su is large and almost constant at 26.7–27.0 kg m23, but it decreases upward at su #26.6 kg m23 in the ventilated layer. The ventilated layer
can be further divided by su 5 26.3 kg m23 into theshallower layer with relatively low AOU and the deeper
layer where AOU is high and greatly increases downward.
This probably reflects primarily the difference in the
subducted water volume between the two layers: that is,the difference that the shallower layer is heavily ventilated
through the mode water formation, whereas the deeper
layer is lightly ventilated because the formed mode waters
are mostly reentrained into the mixed layer within a year.

5. Summary and discussion
Temperature and salinity data from Argo profiling
floats in the North Pacific during 2003–08 have been analyzed to study the structure of winter mixed layer and thesubsurface distribution of potential vorticity in relation to
the formation and subduction of the central mode water(CMW). In late winter, two zonal bands of deep mixed
layer extend north of the Kuroshio Extension at 338–398
and 398–438N from the east coast of Japan to 1608W forroughly 5000 km. The mixed layer properties indicatethat the southern band corresponds to the formation region of the lighter variety of CMW (L-CMW) and that thenorthern band corresponds to that of the denser variety of123

CMW (D-CMW) and the recently identified transition
region mode water (TRMW). The subsurface geostrophic
flow in these regions is generally eastward, with a small
northward component.
Within the L-CMW (D-CWM–TRMW) formation region, the mixed layer density decreases eastward: that is, tothe downstream, from 26.3 (26.6) to 25.7 (26.1) kg m23.
The mixed layer depth in both regions increases to thedownstream in the western part and gradually decreases inthe eastern part, reaching the maximum at 1708E. As
a result, the mode waters formed in the western part of theformation regions west of 1708E are reentrained into themixed layer in the farther east in the following winter andmodified to the lighter water, thus being unable to be
subducted to the permanent pycnocline. In the easternpart of the formation regions between 1708E and 1608W,
on the other hand, the L-CMW (D-CWM–TRMW) withpotential density (su) of 25.7–26.2 (26.1–26.4) kg m23 is
subducted to the permanent pycnocline, and is then advected anticyclonically in the subtropical gyre. The subduction in the eastern part of the formation regions ispreconditioned by the processes in the western part, as
speculated by Ladd and Thompson (2000) and Mecking
and Warner (2001).
These results imply that during the analysis period
large-scale subduction to the permanent pycnocline occurs
up to su 5 26.4 kg m23 in the open North Pacific, whereas
the mixed layer density reaches the maximum of 26.6
kg m23. This is largely supported by the vertical distribution of apparent oxygen utilization in a hydrographic section along 308N. Earlier studies used maps of winter sea
surface density and the Sverdrup streamfunction to demonstrate that the large-scale subtropical subduction is limited to su lower than 26.2 or 26.3 kg m23 (Talley 1985),
which could be modulated by the winter mixed layer
structure (Woods 1985). In this study, we have clarified
how this works in detail based on the float observations.
Although both D-CMW and TRMW are formed in thenorthern formation region, the TRMW formation is highly
limited in the eastern part of the region east of 1708E
where the subduction to the permanent pycnocline occurs(Figs. 4, 5). Therefore, little TRMW is believed to be
subducted to the permanent pycnocline in spite of its
abundance in the western part of the region. This probably
explains partly the discrepancy in the volume distribution
on the potential temperature–salinity (u–S) diagram between the winter mixed layer and the lower ventilated
pycnocline in the subtropical gyre: in other words, thereason why the water in the TRMW u–S range does notproduce a volumetric mode in the permanent pycnocline
in the subtropical gyre, as argued by Suga et al. (2008).
In the repeat hydrographic section along 1658E, the
L-CMW and D-CMW–TRMW formation regions are

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FIG. 9. Frequency of Q lower than a critical value on each isopycnal between su 5 25.7 and 26.6 kg m23(the su value is given in the top left corner of each panel), based on Argo profiling float data in
June–October of 2003–08. The frequency for each 18 3 18 grid box is calculated using the float data in a
38 3 38 grid box centered by the 18 3 18 grid box, only if 10 or more data exist in the 38 3 38 grid box. Criticalvalues of 2.0, 1.75, and 1.5 3 10210 m21 s21 are used for the isopycnals of su 5 25.7–26.0, 26.1–26.2,
and 26.3–26.6 kg m23, respectively. Thick gray curve denotes the outcrop of each isopycnal in March, andblack contours indicate annual-mean acceleration potential relative to 2000 dbar (the contour interval is
0.5 m2 s22) on each isopycnal south of the March outcrop, both based on the World Ocean Atlas 2001
climatology (Conkright et al. 2002).

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125

FIG. 9. (Continued)

separated by a thermohaline front associated with thenorthern bifurcation of the Kuroshio Extension (Okaand Suga 2005). However, our analysis has demonstrated
that both formation regions extend zonally beyond thelongitude range of the Kuroshio bifurcation (approximately between 1508E and 1708W; Levine and White 1983;
Mizuno and White 1983; Sainz-Trapaga et al. 2001) to both
the west and the east. This might mean that the two zonal

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FIG. 10. Distributions of (a) Q (10210 m21 s21) and (b) AOU (ml l21) with respect to su at WHP P02
section along 308N in June–August 2004. Contours are drawn at 0.5, 1, 1.5, 2, 3, 4, and 5 3 10210 m21 s21,
and values ,1 3 10210 m21 s21 (of 1–2 3 10210 m21 s21) are heavily (lightly) shaded in (a). Negativevalues are shaded in (b). Tick marks on the top of each panel indicate the CTDO2 stations.

low-Q bands of L-CMW and D-CMW–TRMW in theirformation regions are not separated by the Kuroshio bifurcation, as Oka and Suga (2005) speculated, but rather
enhance the bifurcation. If, for example, an L-CMW
pycnostad with su of 26.0 kg m23 and a D-CMW pycnostad with 26.4 kg m23 are aligned at similar depths
in a meridional section, the isopycnals of su 5 26.1–
26.3 kg m23, which passes below the L-CMW pycnostad
to the south and above the D-CMW pycnostad to thenorth, should shallow sharply northward between the two
pycnostads, leading to the generation of an eastward flow.
In fact, the longitude range of the Kuroshio bifurcation

corresponds to that where both L-CMW and D-CMW are
thickest (Figs. 7b,c).
Although the formation regions of STMW and the twovarieties of CMW are all zonal bands of deep mixed layer
with considerable zonal variations of properties (Figs. 4,5), the subduction to the permanent pycnocline occurs ina quite different manner between these regions. STMW isformed in the recirculation gyre south of the Kuroshio
and the Kuroshio Extension, which consists of several
anticyclonic circulations (Oka 2009). The STMW formedin each anticyclonic circulation in winter tends to remainin the respective circulation after spring and is mostly

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OKA ET AL.

FIG. 11. Vertical profiles of Q (thin solid curve), AOU (thicksolid curve), and ›AOU/›su (dashed curve) with respect to su,
averaged at 33 CTDO2 stations between 1758 and 1558W at the
WHP P02 section.

reentrained into the mixed layer in the following winter,
while some portion seeps into the southern region, mainly
because of mesoscale activities (Nishikawa et al. 2010),
thereby being subducted to the permanent pycnocline. On
the other hand, each variety of CMW is advected eastward
within the formation region, being ‘‘matured’’ for several
years, until it is finally subducted to the permanent pycnocline from the eastern part of the formation region.
Such a difference in the ‘‘ripening period’’ suggests thatthe year-to-year variation of the CMW properties is muchsmaller than that of the STMW properties, as we assumed
in section 1. In fact, the late winter sea surface temperature
anomaly in the CMW formation region is highly correlated with that one year later (Hanawa and Sugimoto
2004; Sugimoto and Hanawa 2005).
Nevertheless, decadal and longer time-scale variationof the CMW properties can be large because of changes
in the surface heat flux or to meridional shifts of the current
system around the CMW formation region (Nonaka et al.
2006). In the subtropical and subarctic North Pacific, several studies reported long-term AOU changes in the subsurface centered at su 5 26.6 kg m23 (e.g., Watanabe et al.
2001; Emerson et al. 2004; Mecking et al. 2008), which seemto be characterized by a bidecadal oscillation (Andreev andKusakabe 2001; Ono et al. 2001). The long-term variations

127of the CMW subduction scheme and the resultant subduction limit (su 5 26.4 kg m23 for 2003–08) might be acause of such AOU variation. With the accumulated Argo
float data for nearly 10 yr, detailed decadal variation of the
CMW subduction and circulation, its mechanism, and theassociated propagation of temperature and AOU anomalies in the permanent pycnocline shall be the next interesting theme.
In addition to the large-scale, alongfront CMW subduction investigated in this study, eddy-induced, mesoscale
CMW subduction is also thought to be important (e.g., Qu
et al. 2002). Oka et al. (2009) recently reported a subsurface mesoscale eddy observed at 27.58N, 1458E in fall 2008,
which contains a CMW pycnostad with low AOU indicating that the pycnostad is formed only half a year ago.
They speculated that the pycnostad is subducted to thepermanent pycnocline in association with the formation of
the subsurface mesoscale eddy and its migration across theKuroshio Extension by the Spall (1995) mechanism, although the reason why the eddy reaches as far south as
27.58N so quickly is unclear. Such mesoscale subduction
can alter the CMW distribution in the subtropical gyre
determined by the large-scale subduction, particularly inthe su range of 26.5–26.6 kg m23 in which CMW cannot be
subducted to the permanent pycnocline on the large scale,at least for the period of 2003–08. The relative importanceof the two types of CMW subduction and its influence on
the CMW distribution in the subtropical gyre need to be
examined in detail in a future study, in cooperation withhigh-resolution ocean general circulation models.
Acknowledgments. The authors thank Tsuyoshi Ohirafor help in preparing the Argo float data and Sabine
Mecking, Hisashi Nakamura, Kensuke Takeuchi, Lynne
Talley, Youichi Tanimoto, and an anonymous reviewer forconstructive and helpful comments. This work was supported in part by the Japan Society for Promotion of Science [KAKENHI, Grant-in-Aid for Scientific Research(B) 21340133] and the Ministry of Education, Culture,
Sports, Science and Technology (MEXT), Japan [Grantin-Aid for Scientific Research in Priority Areas ‘‘Western
Pacific Air-Sea Interaction Study (W-PASS)’’ underGrant 21014004] and by funds from cooperative program(No. 131 in 2007, No. 116 in 2008, and No. 142 in 2009)
provided by Ocean Research Institute, The University ofTokyo and the ‘‘Studies on Prediction and Application ofFish Species Alteration (SUPRFISH)’’ sponsored by the
Agriculture, Forestry and Fisheries Research Council(AFFRC), Japan. The Argo float data used in this study
were collected and made freely available by the International Argo Project and the national programs that
contribute to it (available online at http://www.argo.ucsd.
edu; http://argo.jcommops.org).

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