J Oceanogr (2015) 71:469–497
DOI 10.1007/s10872-015-0283-7“Hot Spots” in the Climate System: New Developments inthe Extratropical Ocean-Atmosphere Interaction Research

SPECIAL SECTION: REVIEW

Oceanic fronts and jets around Japan: a review
Shinichiro Kida1 · Humio Mitsudera2 · Shigeru Aoki2 · Xinyu Guo3 · Shin‑ichi Ito4 ·Fumiaki Kobashi5 · Nobumasa Komori1 · Atsushi Kubokawa6 · Toru Miyama1 ·Ryosuke Morie7 · Hisashi Nakamura1,8 · Tomohiro Nakamura2 · Hideyuki Nakano9 ·Hajime Nishigaki10 · Masami Nonaka1 · Hideharu Sasaki1 · Yoshi N. Sasaki11 ·Toshio Suga12 · Shusaku Sugimoto12 · Bunmei Taguchi1 · Koutarou Takaya13 ·Tomoki Tozuka14 · Hiroyuki Tsujino9 · Norihisa Usui9

Received: 17 July 2014 / Revised: 26 February 2015 / Accepted: 9 March 2015 / Published online: 8 April 2015
© The Oceanographic Society of Japan and Springer Japan 2015

Abstract This article reviews progress in our understanding of oceanic fronts around Japan and their roles in air–sea interaction. Fronts associated with the Kuroshio and its
extension, fronts within the area of the Kuroshio-Oyashio
confluence, and the subtropical fronts are described withparticular emphasis on their structure, variability, and rolein air–sea interaction. The discussion also extends to thefronts in the coastal and marginal seas, the Seto Inland Sea
and Japan Sea. Studies on oceanic fronts have progressedsignificantly during the past decade, but many of thesestudies focus on processes at individual fronts and do notprovide a comprehensive view. Hence, one of the goals of
this article is to review the oceanic fronts around Japan bydescribing the processes based on common metrics. These
metrics focus primarily on surface properties to obtaininsights into air–sea interactions that occur along oceanic
fronts. The basic characteristics derived for each front (i.e.,
metrics) are then presented as a table. We envision thatmany of the coupled ocean-atmosphere global circulation
models in the coming decade will represent oceanic fronts
reasonably well, and it is hoped that this review along withthe table of metrics will provide a useful benchmark forevaluating these models.

* Shinichiro Kida
kidas@jamstec.go.jp

7

Graduate School of Environmental Science, Hokkaido
University, Sapporo, Japan

8

Research Center for Advanced Science and Technology,
The University of Tokyo, Tokyo, Japan

9

Meteorological Research Institute, Tsukuba, Japan
Faculty of Education and Welfare Science, Oita University,
Oita, Japan

1

Application Laboratory, Japan Agency for MarineEarth Science and Technology, 3173‑25 Showa‑machi,
Kanazawa‑ku, Yokohama 236‑0001, Japan

Keywords Oceanic fronts · Western North Pacific · Air–sea interaction · Kuroshio Extension · Kuroshio-Oyashio
confluence region · Subtropical fronts

1 Oceanic fronts and their roles in climate
Scientific studies of oceanic fronts have a long and rich history in the field of oceanography since the first half of thetwentieth century. In Japan, oceanic fronts were referred toas “shiome” (e.g., Uda 1938) and have been studied extensively because of their importance for fisheries. Kitahara

2Institute of Low Temperature Science, Hokkaido University,
Sapporo, Japan

10

3

Center for Marine Environmental Studies, Ehime University,
Matsuyama, Japan

11

Graduate School of Science, Hokkaido University, Sapporo,
Japan

4

Atmosphere and Ocean Research Institute, The Universityof Tokyo, Kashiwa, Japan

12

Department of Geophysics, Graduate School of Science,
Tohoku University, Sendai, Japan

5

Graduate School of Marine Science and Technology, Tokyo
University of Marine Science and Technology, Tokyo, Japan

13

Faculty of Science, Kyoto Sangyo University, Kyoto, Japan

14

Department of Earth and Planetary Science, Graduate Schoolof Science, The University of Tokyo, Tokyo, Japan

6

Faculty of Environmental Earth Science, Hokkaido
University, Sapporo, Japan

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(1921) proposed what is known today as Kitahara’s law:fish tend to aggregate around fronts. The importance offronts led to the initiation of operational oceanography inJapan in 1935, with oceanic conditions and fishing groundpositions analyzed routinely from oceanic temperature andfish-catch data collected from fishing boats, which werebroadcasted through the NHK (Japan Broadcasting Corp.)
radio (Kimura 1949). Oceanic fronts were called “tidalrips” or “current rips” in English. In recognition of its similarity to an atmospheric front, Cromwell and Reid (1956)
first introduced the term “oceanic front” defined as “a band
along the sea surface across which the density changes
abruptly.” Today, the term “oceanic front” is used in asomewhat broader sense as an oceanic zone with a strong
horizontal gradient in water properties such as temperature,salinity, etc.
Oceanic fronts are ubiquitous with various magnitudes
and spatial scales. On the large scale, fronts form alongoceanic jets, such as the Kuroshio Extension (KE), andthrough coastal upwelling. On a smaller scale, fronts form
through tides, freshwater input from rivers, and along continental shelf breaks. While many studies have investigatedtheir formation mechanisms, some oceanic fronts along
jets have recently gained interest as locations of subduction where the oceanic interior exchanges properties withthe surface mixed layer (Pollard and Regier 1992). Since
subduction plays a critical role in determining the dynamical properties and stratification of the ocean interior (e.g.,
Joyce and Jenkins 1993; Spall 1995), full understanding of
this mechanism has thus been considered one of the fundamental problems for understanding the global oceanic circulation (see Talley et al. 2011 and reference therein).
Today, extra-tropical oceanic fronts are recognized in
the climate system as narrow zones of vigorous air–sea
coupling that occurs through the modification of the planetary boundary layer (PBL) (Xie 2004; Small et al. 2008;
Kelly et al. 2010). Across an oceanic front, differential heatrelease into the atmosphere yields cross-frontal contrasts instratification and thus vertical mixing by turbulence within
the PBL, influencing the downward transport of wind
momentum to the surface (Wallace et al. 1989; Hayes et al.
1989). The frontal influence on the surface wind-speed
through the modulated vertical mixing depends on the winddirection relative to the frontal axis (Chelton et al. 2004). A
front can modify cross-frontal winds acting to yield surface
convergence/divergence, whereas surface wind curl canbe generated for winds along the frontal axis. The modulation of surface winds could be complicated under thestrong low-level thermal winds (Tanimoto et al. 2011). Differential heat release across an oceanic front also modifies
thermal conditions in the PBL and thereby surface pressure via the hydrostatic effect (Lindzen and Nigam 1987),
with frictional surface convergence over locally warm

13

S. Kida et al.

water (Tanimoto et al. 2011). This hydrostatic effect has
been shown to be operative ubiquitously along the major
warm western boundary currents (Shimada and Minobe2011). With such frictional near-surface wind convergenceat work, locally enhanced heat and moisture supplied by awarm western boundary current can lead to organizationof convective cloud systems along the current, as shown inrecent observational and numerical studies (Minobe et al.
2008, 2010; Kuwano-Yoshida et al. 2010; Miyama et al.
2012).
While identifying solid evidence of air–sea couplingon the large-scale atmospheric circulation has been challenging, significant advances have been made during thepast decade by uncovering the potential impacts of the sea
surface temperature (SST) fronts on the free troposphere(e.g., Nakamura et al. 2004; Minobe et al. 2008; Kwon
et al. 2010). The vigorous heat and moisture release into
the atmosphere near oceanic fronts and their sharp crossfrontal contrast (Fig. 1) are beginning to be accepted as
an essential feature of the extra-tropical climate (Hoskins
2012; Imawaki et al. 2013). In fact, decadal-scale variability of the SST in the extratropical North Pacific is known tobe concentrated on oceanic frontal zones (e.g., Nakamura
et al. 1997; Nakamura and Kazmin 2003). By modulatingheat supply into the atmosphere (Tanimoto et al. 2003), persistent SST anomalies generated along the oceanic frontal
zone in the western North Pacific with its meridional displacement have the potential to force a large-scale atmospheric anomaly recognized as the anomalous Aleutian Low(Frankignoul et al. 2011; Taguchi et al. 2012; Okajima
et al. 2014) through modulating storm-track activity.
In this article, we aim to provide a review of dynamical studies on some of the oceanic fronts that exist in thewestern North Pacific (Figs. 1, 2). We will focus mainlyon those associated with the western boundary currents,
such as the Kuroshio and Oyashio, and jets that lie in theinterior, such as the KE, those in the area of the KuroshioOyashio confluence [referred to as the Kuroshio-Oyashio
confluence (KOC) region hereafter; see Sect. 4 for details
on the terminology of this region], and subtropical fronts(STFs). Some of the coastal and marginal-sea scale features around Japan will also be discussed. While we cannot completely cover all of the fronts around Japan, many
fronts in the western North Pacific have received extensive attention because of the “Hot Spot in Climate System” Project (Nakamura 2012), and significant progresshas been made on various aspects. Since the knowledgeof fronts around Japan has not been collectively discussed
since Uda (1938), we hope that this article will be a step
toward synthesizing what has accumulated separately ateach front and updating some of the recent findings ontheir impact on the atmosphere. As one of the tools for
examining the fronts comprehensively and qualitatively,

Oceanic fronts and jets around Japan

471

Fig. 1  Global distribution of the sum of latent and sensible heat
fluxes based on the J-OFURO2 (1993–2007). The region of thefronts focused on in this article is the squared dotted lines and isshown in Fig. 2. The locations of the Gulf Stream, Agulhas Current,

Agulhas Return Current, and Subantarctic Front are shown by blacksolid lines, where they show some resemblance to the western North
Pacific as a region of enhanced heat fluxes

Fig. 2  Geographical locations of oceanic fronts over the western
North Pacific that are described in this article. The inset shows thetidal front in the Seto Inland Sea of Japan. The start and end of thefrontal lines are those in Tables 1, 2, 3, 4, and 5, and the arrows correspond to the direction of the currents along the frontal axis. As indicated with different colors, the fronts are grouped together based on

the sections in this article. Light-colored lines are drawn to illustrate
the connection of fronts. The names of the fronts are abbreviated asindicated in Tables 1, 2, 3, 4, and 5. SAB is the subarctic boundary.
The front associated with the SAC is indicated as SAF here based on
Ueno and Yasuda (2000) and Yasuda (2003)

we have also created a table of metrics that summarizes
the basic dynamical properties of each front. We believethat such a table will be helpful for discussing common
and unique aspects of oceanic fronts around Japan andserves as a benchmark for examining how well these fronts
are represented by numerical models.

This article is organized as follows. In Sect. 2, we introduce the dynamical parameters (i.e., metrics) that are commonly used to describe oceanic fronts and present a tableof metrics for the oceanic fronts in the above-mentioned
western North Pacific. In Sect. 3, we review the individual fronts in more detail and use the metric table for

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comparison, starting from those fronts associated with the
Kuroshio. Extra emphasis is placed on the KE since thisfeature is by far the most well-studied front in the region.
Fronts within the KOC region, such as the northern branchof the KE [referred to as the Kuroshio Extension northernbranch (KENB) hereafter; see Sect. 4 for details on the terminology] and those in the Subarctic Frontal Zone (SAFZ),
are reviewed in Sect. 4. Fronts in the subtropical region are
reviewed in Sect. 5. Fronts in the coastal and marginal seassurrounding Japan are reviewed in Sect. 6. The summaryand discussion are presented in Sect. 7.

2 A table of metrics
2.1 The dynamical properties of oceanic fronts
What are oceanic fronts? As mentioned earlier, Cromwell
and Reid (1956) define an oceanic front as “a band along
the sea surface across which the density changes abruptly.”
Yanagi (1987) defines fronts as “where two water masses
meet, with the spatial gradients of temperature and/or salinity much stronger than the background.” Since the temperature and salinity of seawater are spatially inhomogeneousin nature, convergence, divergence, shear, and mixing often
lead to formation of “fronts” in various places and withdifferent magnitudes. As the mechanism behind the formation of fronts varies, the appropriate set of parametersthat describe the underlying dynamics of fronts is also
likely to differ for each front. Yanagi (1987) has attempted
to classify the oceanic fronts based on their primary forcing agents and geographical locations. In this article, ourfocus is placed on the dynamical features that are common to many fronts. Our focus is also placed on the fronts’
climatological surface properties since the fronts are now
recognized as regions of intensive air-sea interaction. For
example, the cross-frontal SST gradient is in direct contact
with the atmosphere and therefore is emphasized here morethan the corresponding gradient in subsurface temperature.
The metrics estimated for individual fronts may be classified into four property types as follows. (1) Geometricproperties: the location and length of the oceanic front. The
latter metric provides the length scale of the front, which
will be important for considering the spatial scale of its
dynamics as well as the atmospheric processes it might
influence. (2) Surface water mass properties: SST, sea
surface salinity (SSS), and their gradients. These metrics
are the surface properties of what is traditionally used for
defining oceanic fronts and show the differences between
the two water masses in contact at the frontal axis. The SST
and its gradient are also important parameters on which theheat and moisture exchange with the atmosphere depends.
(3) Dynamical properties: sea surface height (SSH)13

S. Kida et al.

gradient, along-front flow speed, volume transport, frontal
depth from the surface, and eddy kinetic energy (EKE) at
the surface. These metrics provide the information on the
momentum field associated with fronts. The first four ofthe metrics of dynamical properties describe the mean flow
field, while the last one gives a measure of eddy activity.
Through the thermal wind balance, cross-frontal density
gradients determined from the temperature and salinity distribution are linked to vertical shear in velocity with flow
speed typically stronger toward the surface. From the geostrophic balance, the SSH gradient is a surface parameterthat reflects the geostrophic velocity at the surface, and the
SSH contours are often used for locating the frontal axis.
(4) Surface fluxes: latent and sensible heat fluxes. As these
are involved in air–sea interaction in the vicinity of an oceanic front, the turbulent heat fluxes depend on both oceanicand atmospheric conditions. The metrics listed above arelimited and by no means complete. Although inclusion ofsome other metrics may lead to a more complete description of frontal characteristics, we focus on the metrics listed
above for conciseness. Specific descriptions for individual
fronts are given in Sects. 3–6.
2.2 The metric table
Tables 1, 2, 3, 4, and 5 show various metrics estimated at
the oceanic fronts. Tables 1, 3, and 4 show the KE, Kuroshio south of Japan (K-SOJ), Kuroshio along the shelf
break of the East China Sea (K-ECS), KENB, which isalso referred to as the Kuroshio Bifurcation front or the
Kuroshio northern branch (see Sect. 4 for details on theterminology), Subarctic Current (SAC), Isoguchi Jets (J1and J2), SAFZ, and northern, southern, and eastern STFs
(Fig. 2). These fronts are associated with the westernboundary currents and their extensions or those that formin the interior of the ocean. In addition, the corresponding
metrics of fronts in coastal and marginal seas around Japanare presented, such as the tidal fronts of Seto Inland Sea
and the Japan Sea subpolar front (JS-SPF) (Table 5). Some
fronts that are not located in the western North Pacific arealso discussed, mainly for comparison. Specifically, major
fronts along the Gulf Stream (GS), Agulhas Current (AC),Agulhas Return Current (ARC), and Antarctic Circumpolar Current (ACC) are compared to the KE (Table 2). Likewise, the Hawaiian Lee Countercurrent (HLCC) is compared to the western Pacific STFs (Table 4).
An overview of the data sets and methods used to estimate the metrics for the aforementioned fronts are given
here. More details are described in the Appendix. We madean effort to use similar observational data sets and periods
unless noted otherwise (Table 6 in the Appendix), and thesedata sets are chosen primarily because of the high spatial
resolution necessary for representing realistic values for[0.19] (Sep)[0.10] (May)61 ± 11 [71 ± 13] (Ann)SST gradient [°C (100 km)−1]

Salinity gradient [(100 km)−1]

Dynamics
SSH gradient [cm (100 km)−1]
Flow speed (m s−1)

110 ± 13 [140 ± 18] (Jan)0 ± 3 [10 ± 5] (Aug)Net flux
440 ± 40 [500 ± 50] (Jan)80 ± 20 [90 ± 20] (July)

52 ± 12 [61 ± 16] (Ann)1.0 ± 0.3 (10 m) (Ann)140 ± 16 (Jan)5 ± 3 (July)

290 ± 41 (Dec)
55 ± 12 (June)

1000
0.05–0.2

42

0.70–1.580 ± 2.9 (Ann)

0.1–0.2 (Sep)
0.02–0.05 (Mar)28.5 ± 0.26 (Aug)19 ± 0.39 (Feb)
5 ± 0.80 (Dec)
2.2 ± 0.55 (Aug)34.7–34.8 (Mar)34–34.1 (Sep)

30°N, 131°E
32–34.3°N, 140°E
1000–1500

K-SOJ

South of Japan106 ± 12 (Jan)2 ± 2 (July)

278 ± 34 (Dec)
75 ± 13 (June)

700
0.1

25

1

50 ± 4 (Ann)1.12 ± 0.95 (July)
0.04 ± 0.11 (Jan)29.10 ± 0.76 (July)
22.62 ± 0.58 (Jan)3.63 ± 1.67 (Jan)0.67 ± 0.97 (July)
34.70 ± 0.11 (Jan)34.12 ± 0.32 (July)

26°N, 124°E
31°N, 127°E
600

K-ECS

East China Sea

The metrics for KE are provided from observations (left) and FRA-JCOPE2 (right). The acronyms of the fronts are written in the first row. The metrics are categorized into four property types,
which are written in the first column. Those with two rows show the maximum monthly mean above the minimum monthly mean with the months written in brackets. Annual means are indicated as Ann. The magnitude of the interannual variability is shown with the ± signs. Metrics for KE show values from the whole frontal axis and those from the upstream axis in [squaredbrackets]. Metrics from past studies are written in italics and those from numerical models are underlined. Detailed methods about how each metric is estimated are provided in the Appendix

Sensible flux (W m−2)

260 ± 37 [290 ± 45] (Dec)30 ± 12 [30 ± 15] (July)

500–1500
0.09

Depth (m)EKE (m2 s−2)

Heat flux
Latent flux (W m−2)

140

Transport (Sv)

1.4

26.3 ± 0.6 (Sep) [26.7 ± 0.7 (Aug)]
15.8 ± 1.0 (Mar) [16.3 ± 1.1 (Mar)]2.4 ± 0.6 (Feb) [2.9 ± 0.8 (Jan)]0.8 ± 0.2 (Aug) [1.0 ± 0.4 (Aug)]34.7 ± 0.05 (Apr) [34.7 ± 0.05 (Apr)]
34.3 ± 0.05 (Sep) [34.2 ± 0.06 (Aug)]

26.8 ± 0.7 [27.2 ± 0.7] (Aug)
16.3 ± 0.6 [16.6 ± 0.8] (Mar)1.8 ± 0.7 (Dec) [2.1 ± 0.7] (Dec)
0.7 ± 0.2 (Aug) [0.9 ± 0.3] (Aug)[34.6] (Apr)[34.2] (Aug)0.17 ± 0.04 (Oct) [0.24 ± 0.09 (Oct)]0.10 ± 0.04 (July) [0.13 ± 0.07(July)]

35°N, 141°E
34°N, 165°E
3000–4000

35°N, 141°E
34°N, 165°E
3000–4000

SalinityKE

Extension (FRA-JCOPE2)

KE

Acronym
LocationStart (lat, lon)
End (lat, lon)
Length (km)Surface water
Temperature (°C)

Extension (observations)

Kuroshio

Table 1  Metrics for the fronts associated with main branch of the Kuroshio: KE, Kuroshio south of Japan, and Kuroshio along the shelf break of the East China Sea

Oceanic fronts and jets around Japan
473

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S. Kida et al.

Table 2  Metrics for major fronts compared to the KE: the Gulf Stream, Agulhas Current, Agulhas Return Current, and Antarctic Circumpolar
Current-Subantarctic front
Gulf Stream

Agulhas Current

Agulhas Return Current

Antarctic Circumpolar
Current-Subantarctic front

GS

AC

ARC

ACC-SAF

35°N, 75°W
40°N, 50°W
3500

27°S, 33°E
40°S, 19°E
2000

40°S, 19°E
45°S, 66°E
3900

41°S (north) 0–360°E
60°S (south) 0–360°E
31,00027 ± 0.8 (Aug)12 ± 2.5 (Mar)4 ± 1.5 (Mar)28 ± 2 (Feb)
19 ± 1 (Aug)
1.8 ± 0.7 (Ann)15 ± 2 (Feb)
10 ± 2 (Aug)
4.1 ± 1.2 (Ann)35–36 (Ann)
0.5 (Ann)35.5 (Ann)
0.1 (Ann)34.5 (Ann)
0.4 (Ann)

8.0 ± 0.7 (Feb)
4.8 ± 0.5 (Sep)0.9 ± 0.1 (Mar)0.8 ± 0.1 (Aug)
34 (Ann)
0.05 (Ann)80 ± 30 (Ann)70 ± 20 (Ann)50 ± 20 (Ann)

20–70

0.8

0.4

Transport (Sv)

140

70

54

30–60

Depth (m)EKE (m2 s−2)3500
0.2 ± 0.1 (Ann)2400
0.1 ± 0.1 (Ann)1500
0.2 ± 0.2 (Ann)

2500
0.4

350 ± 75 (Jan)100 ± 25 (July)

327 ± 67 (Aug)191 ± 30 (Jan)251 ± 31 (Aug)181 ± 34 (Jan)65 ± 7 (Sep)−39 ± 4 (Jan)Acronym
LocationStart (lat, lon)
End (lat, lon)
Length (km)Surface water
Temperature (°C)
SST gradient [°C (100 km)−1]
Salinity
Salinity gradient [(100 km)−1]
Dynamics
SSH gradient [cm (100 km)−1]
Flow speed (m s−1)

Heat flux
Latent flux (W m−2)
Sensible flux (W m−2)

1.5

150 ± 45 (Jan)10 ± 5 (July)

1.2

111 ± 34 (July)
56 ± 11 (Jan)71 ± 21 (July)
38 ± 11 (Feb)

3.0 ± 3.5 (Sep)−7.5 ± 2.0 (Feb)

See the footnote of Table 1 for descriptions

sharp gradients of oceanic variables and thus for defining
fronts. For the surface water mass properties in Tables 1, 2,
3, 4, and 5, monthly means of AMSR-E SST (2003–2008;
Gentemann et al. 2010), MGDSST (2003–2008; Kurihara
et al. 2000), and OISST (2003–2008; Reynolds et al. 2002,2007) are used for estimating the SST and its gradient. The
monthly climatology of World Ocean Atlas 2005 (Antonov
et al. 2006), 2009 (Antonov et al. 2010), and 2013 (Zweng
et al. 2013) is used for estimating the SSS and its gradient.
For the dynamical properties, the monthly mean of AVISO(1993–2007; http://www.aviso.oceanobs.com/duacs/) isused for estimating the SSH gradient and EKE. The EKE
is based on geostrophic velocity anomalies. For the surface fluxes, monthly means of latent heat and sensible heat
fluxes are obtained from Version 2 of Japanese Ocean Flux
data sets using Remote Sensing Observations (J-OFURO2;
1993–2007; Tomita et al. 2010). The J-OFURO2 is chosen because of its best estimate for the KE region (Tomita
et al. 2010). Standard deviation of the interannual variability is provided for SST, SSH, their gradients, EKE, and theheat fluxes so that they provide a measure of the variability
observed at the fronts. However, SSS and its gradient are13

only presented with their monthly means because only theclimatological monthly means are available and interannual
variability cannot be estimated. Other metrics, such as location, along-front flow speed, volume transport, and frontal
depth, are mostly based on past studies. Metrics are also
estimated from Japan Coastal Ocean Predictability Experiment-2 (JCOPE-2; Miyazawa et al. 2009; see Appendix)
for the KE (Table 1). This is to compare how well current state-of-the-art data assimilation models capture theobserved frontal signatures.
The estimates of each metric are based on monthly mean
fields unless noted otherwise. This is because we aim topresent a climatological view of oceanic fronts and many ofthe past studies are based on monthly means. Many fronts
are also highly variable in time, and we find it reasonable
to focus on the monthly mean time scale to distinguish therole of fronts from eddies with much smaller time scales.
Spatial gradients are per 100 km instead of estimating theirlocal maximum. We chose to use this spatial scale because
it can roughly present the difference in the values acrossa front and is the spatial scale at which the atmosphere is
likely to be affected. A higher spatial scale can capture the475

Oceanic fronts and jets around Japan

Table 3  Metrics for fronts within the Kuroshio-Oyashio confluence region: the Kuroshio Extension northern branch, Subarctic Current, Isoguchi Jets (J1 and J2), and Subarctic Frontal Zone
Kuroshio Extension
northern branch

Subarctic Current

KENB

Isoguchi Jets

Subarctic Frontal Zone

J1

J2

SAC

J1

J2

SAFZ

38°N, 153°E
40°N, 170°E
2000

41°N, 150°E
44°N, 155°E
500

40°N, 150°E
43°N, 155°E
500

40°N, 165°E
43°N, 170°E
500

40°N, 145°E
45°N, 180°6000

22.3 ± 0.7 (Sep)
11.2 ± 0.2 (Mar)1.9 ± 0.4 (June)
0.7 ± 1.0 (Nov)34.4 (May)34.2 (Aug)18.1 ± 0.7 (Aug)
4.0 ± 0.5 (Apr)
3.3 ± 0.4 (Apr)
2.0 ± 1.5 (Oct)33.6 (Mar)33.1 (Sep)

19.9 ± 0.8 (Aug)7.0 ± 0.4 (Mar)3.6 ± 0.8 (Jan)0.6 ± 0.9 (July)
33.9 (Mar)33.5 (Dec)

18.8 ± 1.2 (Sep)
7.8 ± 0.4 (Mar)3.6 ± 0.6 (Feb)
2.0 ± 0.4 (Sep)
33.9 (June)
33.6 (Aug)17 ± 1.0 (Sep)
5.6 ± 0.2 (Mar)0.11 (Ann)0.54 (Sep)
0.35 (Apr, Dec)

0.50 (Dec)
0.34 (Mar)0.47 (Nov)0.32 (Feb)

Dynamics
SSH gradient [cm(100 km)−1]

5.2 ± 13.0 (Ann)12.9 ± 6.9 (Ann)19.0 ± 4.7 (Ann)14.2 ± 4.8 (Ann)

Flow speed (m s−1)
Transport (Sv)0.05 ± 0.14 (Ann)
15* (0–1500 m)

0.13 ± 0.07 (Ann)

0.19 ± 0.05 (Ann)
10* (0–1500 m)

0.14 ± 0.05 (Ann)
10* (0–1500 m)

Depth (m)EKE (m2 s−2)
Heat flux
Latent flux (W m−2)300
0.021 ± 0.007 (Ann)

300
300
300
0.008 ± 0.004 (Ann) 0.013 ± 0.006 (Ann) 0.009 ± 0.005 (Ann)209 ± 29 (Dec)
2 ± 12 (July)

119 ± 20 (Dec)−21 ± 11 (July)

Acronym
LocationStart (lat, lon)
End (lat, lon)
Length (km)Surface water
Temperature (°C)
SST gradient [°C(100 km)−1]
Salinity
Salinity gradient[(100 km)−1]

Sensible flux (W m−2)

81 ± 10 (Jan)−1 ± 3 (July)10

54 ± 16 (Dec)−12 ± 5 (May)163 ± 26 (Dec)−12 ± 15 (July)

89 ± 12 (Jan)−3 ± 7 (May)156 ± 25 (Dec)−12 ± 11 (July)

49 ± 16 (Dec)−6 ± 3 (July)

33.0–33.9 (Ann)
0.2 (Ann)136 ± 20 (Dec)−17 ± 9 (July)

49 ± 15 (Dec)−9 ± 3 (May)Metrics noted with asterisks are our rough estimates (see Appendix). See the footnote of Table 1 for further descriptions

abrupt changes at the front better, but is likely to result inlarge uncertainties as well. We also tried to avoid the differences that arise from products with a different spatial
resolution because high spatial resolution products tend toresolve gradients better and thus yield stronger local maxima of gradient intensity. The metrics are provided withmonthly maximum and minimum values to present the climatological seasonal cycle, and the variability on a shorteror longer time scale is described in Sects. 3–6 where
necessary.

3 Kuroshio fronts
3.1 Kuroshio Extension front (KE front)
The Kuroshio separates from the Japanese coast at the Boso
Peninsula at around 35°N, forming an eastward jet known as
the KE (Fig. 2). This jet is accompanied by two prominent
meanders between the Izu Ridge and the Shatsky Rise (e.g.,

Fig. 2 of Niiler et al. 2003a). The jet is also accompanied by a
sharp subsurface density front, known as the KE front, characterized by a steep upward slope of the main pycnocline tilting
northward. Since the subtropical main pycnocline is controlled
by the temperature stratification, the KE front is essentially atemperature front (Fig. 2 of Nonaka et al. 2006). The positionof the KE front is known to differ from the subtropical-subpolar boundary that is predicted from the Sverdrup streamfunction in the interior. The northern recirculation gyre to the northof the KE (Qiu et al. 2008; Nakano et al. 2008) and effects
of eddies (e.g., Waterman et al. 2011) may contribute to thisdiscrepancy between the Sverdrup streamfunction and theobserved KE front. Despite the many studies carried out thus
far, we have not reached a full understanding of how the KE
front forms and how its mean position is determined.
3.1.1 Metrics used to locate the KE front
The importance of the KE for fisheries led to the early
establishment of metrics that have been used operationally

13

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S. Kida et al.

Table 4  Metrics for fronts located in the subtropical gyre of the North Pacific: Northern, Southern, and Eastern subtropical fronts and the
Hawaiian Lee Counter Current
Subtropical front

Acronym
LocationStart (lat, lon)
End (lat, lon)
Length (km)Surface water
Temperature (°C)
SST gradient [°C (100 km)−1]
Salinity
Salinity gradient [(100 km)−1]
Dynamics
SSH gradient [cm (100 km)−1]
Flow speed (m s−1)
Transport (Sv)
Depth (m)EKE (m2s−2)Heat fluxes
Latent flux (W m−2)
Sensible flux (W m−2)

Hawaiian Lee Countercurrent

North

South

East

NSTF

SSTF

ESTF

HLCC

22–25°N, 130°E
22–25°N, 180°
5100

18–20°N, 130°E
21–22°N, 180°
5200

25–27°N, 180°
25–27°N, 140°W
4000

18°N, 165°E
19–20°N, 160°W
7000

29 ± 0.4 (July–Sep)
24 ± 0.3 (Feb–Mar)
0.7 ± 0.1 (Feb–Mar)
0.2 ± 0.1 (Aug– Oct)35.1 (Jan–July)
34.9 (Sep)29 ± 0.3 (July–Sep)
27 ± 0.3 (Feb–Mar)
0.3 ± 0.1(Jan–Feb)
0.1 ± 0.04 (July–Aug)
34.9 (Jan–July)
34.7 (Sep)27 ± 0.3 (Aug–Sep)
22 ± 0.6 (Feb–Mar)
0.7 ± 0.1 (Feb–Mar)
0.2 ± 0.1 (July–Sep)
35.4 (July–Aug)
35.2 (Mar)28 ± 0.2 (Sep)25 ± 0.2 (Feb)
34.9 (Sep–Jan)
34.7 (May)0.03 (Ann)0.06 (Ann)0.03 (Ann)0.06 (Ann)4 ± 4 (May)
0.4 ± 3 (Sep–Nov)

3 ± 4 (Mar–Apr)
0.4 ± 2 (Nov–Dec)

4 ± 3 (Ann)3.3 ± 2.9 (Ann)0.05 ± 0.05 (Ann)

0.11 + 0.05 (Oct)0.05 ± 0.04 (Mar)0.08 ± 0.06 (May)0.01 ± 0.06 (Sep–Nov)

0.05 ± 0.08 (Mar–Apr)
0.01 ± 0.05 (Nov–Dev)

0.017 ± 0.005
0.035 ± 0.009
0.031 ± 0.006(Apr–June)(May)(May–June)
0.017 ± 0.003 (Dec–Jan) 0.018 ± 0.005 (Dec–Jan) 0.012 ± 0.003 (Nov–Jan)
160 ± 20 (Nov–Jan)
90 ± 20 (May–June)
20 ± 4 (Jan–Feb)
2 ± 2 (Aug–Sep)

180 ± 20 (Nov–Dec)
110 ± 10 (June–Aug)
10 ± 4 (Jan–Feb)
3 ± 2 (Apr–Nov)

160 ± 20 (Nov–Jan)
100 ± 20 (May–June)
12 ± 5 (Nov–Mar)
4 ± 2 (June–Aug)

4.4
1800.022 ± 0.036 (Ann)

195 ± 23 (Dec)
124 ± 21 (May)13 ± 4(Mar)4 ± 1(July)

See the footnote of Table 1 for descriptions

for detecting the KE axis. While the position of the KE
axis is ideally defined from the maximum velocity at thesurface, it was traditionally based on subsurface temperature. This is because temperature data are more commonly available, and there is a close relationship between
the subsurface density/temperature front and the velocity
structure of the KE. Kawai (1969, 1972) suggested that theposition of the maximum surface velocity across the KE
is well correlated with the positions of the following isotherms: 13.9 and 13.7 °C at a depth of 200 m in the longitudinal sectors from 143.5°E to 145.5°E and from 148.0°E
to 151.5°E, respectively (Fig. 3); 9.0 and 8.0 °C at a depthof 400 m are also used for the two sectors. These metrics
are based on a statistical analysis of quasi-concurrent synoptic measurements of subsurface temperature and surface
velocity. The 14.0 °C isotherm at 200 m has been widelyused as a metric for locating the KE axis but, at times,

13

beyond the domains examined by Kawai (e.g., Bingham
1992). This is likely because the availability of temperature data is higher at 200 m depth than at 400 m depth
although the 9 °C isotherm at 400 m depth corresponds
better to the KE axis. Another widely used metric was that
proposed by Mizuno and White (1983), who found that thelocation of the 12 °C isotherm at 300 m depth as a metricis well suited for locating the KE axis based on 1° latitude/
longitude gridded seasonal-mean temperature fields. Thereappears to be a high correlation between 300 m temperature and dynamic thickness between 100 and 1000 db in
the mid-latitude western North Pacific at least as far east
as 170°E (Bernstein and White 1981) in all seasons. The
12 °C isotherm at 300 m depth is also widely accepted asa metric for the KE axis (e.g., Suga et al. 1997). Murakami(1993) proposed a metric for the KE axis with an indicative temperature at 100 m depth. The depth, however, may477

Oceanic fronts and jets around Japan
Table 5  Metrics for fronts located in the coastal and marginal seasaround Japan: Seto Inland Sea tidal front and Japan Sea subpolar
front
Seto Inland Sea
tidal front
Acronym
LocationStart (lat, lon)
End (lat, lon)
Length (km)Surface water
Temperature (°C)

Japan Sea subpolar
front
JS-SPF33.24°N, 131.85°E 40°N, 133°E
33.38°N, 132.08°E 40°N,138°E
40
400
23

23.7 ± 1.1 (Aug)6 ± 1.1 (Mar)4.2 ± 0.6 (Jan)
1.3 ± 0.3 (Aug)34.1 (Jan–May)
33.6 (Oct–Nov)

SST gradient [°C(100 km)−1]
Salinity

4 °C/10 km

Salinity gradient[(100 km)−1]

0.6/10 km

0.18 (Oct)−0.13 (June)

1.0 cm/10 km

11.3 ± 1.5 (Ann)

Dynamics
SSH gradient [cm(100 km)−1]34

Flow speed (m s−1)

0.1–0.2

0.3

Transport (Sv)

0.01

1.2

Depth (m)20
0.01–0.04200

120 (Nov)
40 (June–July)

169 ± 26 (Dec)−6 ± 11 (July)

EKE (m2 s−2)
Heat fluxes
Latent flux (W m−2)
Sensible flux (W m−2)

20 (Dec)
0 (July–Aug)

107 ± 31 (Dec)−4 ± 4 (May)The gradients for the Seto Inland Sea are per 10 km since its spatial
scale is <100 km. See the footnote of Table 1 for descriptions

be too shallow to apply throughout the year since the KE
axis shows distinct seasonality in its best correspondence
to a particular isotherm, which varies from 16 °C in February to 19 °C in October.
Representation of the mean state and variability of theKE has been revolutionized by the advent of satellite altimetry (e.g., Qiu et al. 1991). With gridded products of SSH,
one can define the KE path by tracing an appropriate SSH
contour that is located along or in the immediate vicinity of
local SSH gradient maxima (e.g., Qiu and Chen 2005; Sugimoto and Hanawa 2012). As an example, Qiu and Chen
(2005) offered a comprehensive description of the positionand length of the KE path, cross-frontal SSH gradient, andEKE within the KE region, in addition to their decadal variability (Table 1). Note that the absolute values of SSH used
as an indicator of the KE path depend on a particular SSH
product because the global mean SSH values are different
between the SSH products.

Figure 3 shows the locations of the KE axis as defined by
the isotherms of 14 °C at 200 m depth (red), 12 °C at 300 m
depth (blue), and 9 °C at 400 m depth (green), in addition to
the axis defined as the contour line of SSH (=0.05 m) (black),
all of which are based on the 1993–2007 annual climatology
derived from the FRA-JCOPE2 reanalysis (Miyazawa et al.
2009). It is evident that all the metrics correspond well to thecore of the surface velocity (shading) in the FRA-JCOPE2
reanalysis, especially near Japan, although their discrepancy increases rapidly downstream. Miyazawa et al. (2009)
pointed out that variations of the KE as defined by tracking
grid points of maximum kinetic energy at 100 m depth of
the JCOPE2 reanalysis compare well with the observationalcounterpart, including decadal variability of the KE. Consistency among the definitions of the KE path can also be confirmed in climatological maps based on observations (Fig. 4)
where the two isotherms of 14 °C at 200 m depth and 12 °C
at 300 m depth roughly correspond to the large SSH gradientassociated with the KE at 141°E–165°E. While more rigorouscomparison between the indicative isotherms and their SSH
counterpart may be needed for more in-depth discussions, allof the aforementioned metrics used for determining the KE
axis seem appropriate.
3.1.2 The KE front and its impact on the atmosphere
The flow field associated with the KE can be divided into
up- and downstream regions by the 153°E meridian according to the satellite observations (e.g., Qiu and Chen 2005).
The KE in the upstream region (west of 153°E) is characterized by the presence of two quasi-stationary meanders,
whereas the KE in the downstream region (153°E–165°E)
bifurcates into the (main) KE and the northern branch(KENB). The magnitude of the SST gradient associatedwith the KE exhibits a clear seasonality with the maximum in winter [about 2 °C (100 km)−1] and minimum insummer [about 1 °C (100 km)−1] (Table 1; Fig. 5). In the
upstream region, the strengths of the SST and SSH gradients, in addition to the flow speed, exhibit seasonal (Tatebe
and Yasuda 2001) and interannual to decadal variability(Qiu and Chen 2005, 2010a); therefore, the “typical” values
listed in Table 1 include some uncertainty. It is argued (e.g.,
Howe et al. 2009) that the KE may be better captured in
the stream-coordinate system, in which no lateral smoothing of the jet structure can occur that arises from the Eulerian mean of the highly variable KE axis. Our estimates
for the SST, SSH, and their gradients are therefore basedon the stream-coordinate system. As evident in Fig. 5, theSST gradient associated with the KE in its upstream region
based on FRA-JCOPE2 reanalysis exhibits large interannual variations, whose strength is comparable to that of theclimatological seasonal cycle.

13

478

S. Kida et al.

Fig. 3  Comparison of the locations of the KE calculated fromthe annual mean 1993–2007 climatology of the FRA-JCOPE2
reanalysis. Isocline contours
are 14 °C at 200 m depth (red),
12 °C at 300 m depth (blue),and 9 °C at 400 m depth (green)of the potential temperaturesand 0.05 m of the sea surface
height (black). Gray shadingshows the surface absolute
velocity at 10 m depth

While the KE is associated with a thermocline front, ittransports warm water and causes a strong SST front just to
the north of its axis, especially in winter (Fig. 4). Along thisSST front, turbulent surface heat fluxes also exhibit a strong
frontal gradient (Fig. 4e, f), especially in the KE upstream
region (Figs. 2 and 3 of Taguchi et al. 2009). At 145°E,
for example, the meridional SST gradient exceeds 6 °C(100 km)−1 in winter, while the corresponding gradientin the turbulent heat flux (sum of latent and sensible heat
fluxes) reaches as high as 100 W m−2 (100 km)−1. Alongthe warm KE, latent heat flux is locally enhanced particularly in winter, supplying moisture to the atmosphere (e.g.,
Iizuka 2010), which is numerically found to be importantfor cyclone intensification (Iizuka et al. 2013). Observations and numerical models both show that sensible heat
flux is also enhanced along the KE axis in winter, whichleads to the climatological formation of a near-surface pressure trough slightly south of the KE (Tanimoto et al. 2011)
and associated local maxima of cloudiness (Tokinaga et al.
2009; Masunaga et al. 2014). Observations also revealchanges in the SST and turbulent heat flux (sum of latentand sensible heat fluxes), which reflect the changes in the
KE path. The KE front shows a decadal-scale north-south
movement ranging approximately 200 km in latitude (Qiu
and Chen 2005, 2010a), which yields large SST anomalies persistently between 34° and 36°N exerting significant
synoptic-scale changes in PBL (Joyce et al. 2009). Additionally, the KE front alternates decadally between a stablestate with two quasi-stationary meanders and an unstablestate with a convoluted path (Qiu and Chen 2005; Seo et al.
2014a); in the unstable state, warm eddies detach northward from the KE front (Itoh and Yasuda 2010), resulting
in increased SST and upward heat release in winter (Sugimoto and Hanawa 2011).

13

In summer, by contrast, meridional gradients of SST andturbulent heat flux at 145°E drop to about 2 °C (100 km)−1
and 20 W m−2 (100 km)−1, respectively (Fig. 4d), although
the flow speed as derived from the SSH gradient exhibits
no significant seasonality (Table 1). In situ observations by
Tanimoto et al. (2009) revealed distinct impacts of the KE
on summertime low-level clouds. Fog formation is likely
over the cooler water with the warm, moist southerly winds
across the KE, whereas the formation of convective stratocumulus is most likely over the warm KE under the cool
northerlies. This tendency of low-level clouds is confirmedby another in situ observation by Kawai et al. (2014), who
also show that the depth of the PBL is modulated in thepresence of the KE especially in its upstream region. Theynevertheless observed a local minimum of relative humidity over the warm KE, along which turbulent activity within
the PBL and the consequent entrainment of free-tropospheric air into the PBL were both locally enhanced under
the cool northerlies.
3.2 Kuroshio south of JapanThe Kuroshio is observed to flow into the deep Shikoku
Basin from the shallow East China Sea through the Tokara
Strait (Fig. 2). This portion of the Kuroshio within the Shikoku Basin, which we refer to as the K-SOJ, can take different paths: the “straight path” and “large meander path”(Kawabe 1995). The path length of the Kuroshio axis
between 131°E and 140°E is about 1,000 and 1,500 km forthe straight and large-meander paths, respectively (Table 1),
which are estimated by tracking the strongest part of thesurface geostrophic current (Ambe et al. 2004). These
two states persist from a year to a decade (Kawabe 1987).
Recent modeling studies indicate that the large meander479

Oceanic fronts and jets around Japan

(a)

(b)

(c)

(d)

(e)

(f)

Fig. 4  a Annual climatology (1993–2007) of SSH (contoured every
10 cm) and its horizontal gradient [cm (100 km)−1; shaded] based on
AVISO (Rio et al. 2011). b The corresponding climatological SSH(contoured every 10 cm) and the following associated fronts and jets:(red line) KE defined by the 14 °C isotherm at 200 m depth (Kawai
1972); (magenta line) KE defined by the 12 °C isotherm at 300 m
depth (Mizuno and White 1983); (blue lines) KENB defined by the 6
and 8 °C isotherms at 300 m depth (Mizuno and White 1983); (greenline) subarctic boundary defined by the salinity of 34.0 at 100 m
depth (Favorite et al. 1976); (purple line) SAC traced by the 4 °C iso-therm, which is the southern boundary of the “pure” subarctic water(Favorite et al. 1976; Belkin et al. 2002). Temperature and salinity arebased on WOCE Global Hydrographic Climatology (Gouretski and
Koltermann 2004). c The horizontal gradient of SST in January [°C(100 km)−1; shaded] based on the January climatology (September
2002–August 2011) of AMSR-E SST. Annual climatology of SSH(same as a) is contoured every 10 cm. d Same as c but for July. e
January climatology of latent heat flux (W m−2; shaded) based on the
J-OFURO2 (1993–2007). Annual climatology of SSH (same as a) is
contoured every 10 cm. f Same as e but for sensible heat flux

13

480

Fig. 5  Seasonal variations of SST gradient [°C (100 km)−1; dots]
across the KE between 141°E and 165°E based on the monthly climatology (1993–2007) of the FRA-JCOPE2 reanalysis data. Maximum
and minimum values are those presented in Table 1. The interannual
standard deviation of the SST gradient is indicated as a bar for eachcalendar monthtends to occur and persist when the volume transport is low(Tsujino et al. 2013; Usui et al. 2013).
The velocity of the Kuroshio has a subsurface maximum
and extends downward to at least 1,000 m (Table 1, e.g.,
Nakano et al. 1994; Book et al. 2002). Due to its deep structure, the Kuroshio is accompanied by a large SSH gradientacross its axis [80 cm (100 km)−1] (Fig. 4a). The annual
mean volume transport of the Kuroshio is 42 Sv (Imawaki
et al. 2001), which almost corresponds to the wind-driven
Sverdrup transport. The seasonal variation of the transport(about 10 Sv) is, however, much less than that expectedfrom the Sverdrup transport because of the existence of the
Izu-Ogasawara Ridge, which, based on a numerical model,
filters the seasonal signal propagating westward from theinterior (Isobe and Imawaki 2002).
In contrast to the velocity-related variables, surface
temperature and salinity exhibit significant seasonalchanges (Table 1). In SST fields, the Kuroshio can be recognized as a warm tongue (e.g., Shimada et al. 2005). The
particular feature emerges distinctly in winter and diminishes in summer. As a result, a strong SST front [5 °C(100 km)−1] is formed in winter between the warm tongue
and relatively cold coastal waters to the north of theKuroshio (Fig. 4c), while it weakens in summer [2.2 °C(100 km)−1] (Fig. 4d). It is well known that the 15 °C isotherm at 200 m depth is a good indicator for the Kuroshio
axis (Kawai 1969), but it is usually difficult to define a
specific value of SST that corresponds to the Kuroshio axis
because of the strong seasonal change in SST. A salinity

13

S. Kida et al.

front is also formed between offshore high-salinity waterof the Kuroshio and nearshore low-salinity water becauseof river discharge (Imasato and Qiu 1987). Unlike the SST
front, the magnitude of the salinity front is at its maximum[0.1–0.2 (100 km)−1] in summer and minimum in winter[0.02–0.05 (100 km)−1].
Surface latent and sensible heat fluxes in the region
south of Japan also undergo substantial seasonal variations (Kubota et al. 2002). Under the winter monsoon,
outbreaks of cold and dry continental air over the persistently warm surface water of the Kuroshio induce a large
amount of heat release from the ocean to the atmosphere(Kwon et al. 2010). In winter, the latent and sensible
heat fluxes over the Kuroshio reach 290 and 140 W m−2(Fig. 4e, f), respectively, while they diminish in summer
and the sensible heat flux even becomes negative at times.
A recent observational study by Nakamura et al. (2012)
showed that the huge heat release and bimodal path fluctuations of the Kuroshio have an impact on the development and tracks of wintertime extratropical cyclones passing south of Japan.
3.3 Kuroshio along the shelf break of the East China
Sea
The path of the K-ECS is almost fixed and thus much morestable than that to the south of Japan (Fig. 2). The front
along the Kuroshio is one of the ten thermal fronts in the
East China, Yellow, and Bohai Seas identified from the satellite SST data (HicKox et al. 2000). The Kuroshio front
forms between the Kuroshio main current and continental
shelf water, extending from the northeast of Taiwan (26°N,
124°E) to the southwest of Kyushu (31°N, 127°E) (HicKox
et al. 2000). The Kuroshio surface water is 29 °C in summer and remains warm (22 °C) in winter, while the continental shelf water is 28 °C in summer but cools down to
15 °C in winter under the cold monsoonal airflow becauseof the shallow bathymetry. Consequently, the Kuroshio
front is much stronger in winter than in summer. The maximum water temperature difference across the Kuroshio
front is approximately 4 °C in winter but diminishes insummer (HicKox et al. 2000) (Table 1; Fig. 4c, d).
It is difficult to identify a salinity front between the
Kuroshio main current and shelf water. An apparent salinity front exists just offshore of the Changjiang River mouth
throughout the year (Chen 2009). The low-salinity shelfwater and high-salinity Kuroshio surface water maintainthe salinity gradient across the Kuroshio main current. The
salinity of the Kuroshio surface water is 34.8 in winter anddecreases to 34 in summer. Low-salinity water from theriver mouth spreading more in summer toward the northern continental shelf enhances the salinity gradient (Chen2009).

Oceanic fronts and jets around JapanThe surface current is strong in summer and weak in
winter (Kawabe 1988; Oka and Kawabe 1998). However,the volume transport of the Kuroshio in the East China Sea
exhibits only a small range of seasonal variation (4 Sv inFujiwara 1981; 3.7 Sv in Hinata 1996; 1.6 Sv in Andres
et al. 2008) compared to the Sverdrup transport calculated
from the basin-scale wind fields (Kagimoto and Yamagata
1997). Observations of the Kuroshio Transport across thePN transect show high-frequency variations as large as
10 Sv, which may be caused by frontal eddies (Ichikawa
and Beardsley 1993).
The SST of the Kuroshio exerts influences on the overlying atmosphere (Fig. 4e, f). The winter thermal front
across the Kuroshio pathway is observed to weaken thesurface winds (Kasamo et al. 2014), and the frontal SST
gradient along the shelf break contributes positively tocyclone development (Xie et al. 2002). The high SST
along the Kuroshio is also found, from observations andnumerical models, to play a critical role in organizing an
early summer convective rainband (Miyama et al. 2012),
thus forming a local rainfall maximum in June (Sasaki
et al. 2012a).
3.4 Comparison to the Gulf Stream, Agulhas Current,and Antarctic Circumpolar Current
To further shed light on the characteristics of oceanic fronts
associated with the Kuroshio, it is useful to make a comparison with those associated with the GS, AC, ARC, andACC. As will be seen in this subsection, these oceanic
fronts have similarities in many different aspects with theoceanic fronts associated with the Kuroshio.
The GS is the counterpart of the Kuroshio and the KE
in the North Atlantic subtropical gyre. In particular, after
the separation at Cape Hatteras, the GS has many similarities to the KE as a free eastward jet associated with strong
eddy activity. The GS transports a large amount of warm
water poleward, especially in this downstream region (from60 to 140 Sv between 55°W and 75°W; Johns et al. 1995),
and yields a large amount of heat/moisture release into theatmosphere (500 W m−2) (Fig. 1; Table 2). A sharp SST
front also extends along its northern flank after the separation at Cape Hatteras [4 °C (100 km)−1 in winter], whichis comparable to that of the KE front (Table 1). With thisSST gradient and strong land-sea thermal contrasts with
the North American continent, the GS exerts a strong influence on the overlying atmosphere (e.g., Minobe et al. 2008,2010; Kuwano-Yoshida et al. 2010; Nakamura et al. 2004,2010), as is found for the KE. The large heat release by the
GS results in surface wind convergence aloft and plays animportant role in the enhanced precipitation over the GS(Minobe et al. 2008; Kuwano-Yoshida et al. 2010).

481

The AC (Gordon 1985) and ARC (Lutjeharms andAnsorge 2001; Nonaka et al. 2009; Tozuka and Cronin2014) are the counterparts of Kuroshio and KE in the South
Indian Ocean subtropical gyre (Beal et al. 2011) (Fig. 1;
Table 2). Large amounts of heat, which are comparable to the Kuroshio and KE, are released into the atmosphere with the maximum latent and sensible fluxes of 327
and 111 W m−2, respectively. We note that these values
obtained from the J-OFURO2 data (Tomita et al. 2010)
are much larger than 206 and 61 W m−2, respectively, ofthe OAFlux data (Yu and Weller 2007). The AC and ARC
accompany the SST (SSS) gradient of 1.8 °C (100 km)−1[0.1 (100 km)−1] and 4.1 °C (100 km)−1 [0.4 (100 km)−1],
respectively, which are as strong as their counterpart in the
KE region. The strong SST gradient is due to large transport by these two currents (Tozuka and Cronin 2014); themean volume transport of the AC and ARC is estimatedto be 70 Sv (Bryden et al. 2005) and 54 Sv (Lutjeharms
and Ansorge 2001), respectively. Compared to the K-SOJ,
the volume transport of the AC is about 30 Sv stronger. An
active ocean-atmosphere interaction occurs over the ARC
region (Liu et al. 2007; Nonaka et al. 2009; Tozuka andCronin 2014), and the associated storm track plays a role in
the poleward extension of the summertime subtropical high(Miyasaka and Nakamura 2010).
The Subantarctic front (ACC-SAF) is the most equatorward feature of the ACC, which is continuous across the
major oceanic basins (Fig. 1; Table 2) and about 30,000 kmlong (Rintoul and Naveira Garabato 2013; Orsi et al. 1995).
The location of the ACC-SAF ranges from roughly equal(41°S) to much higher latitudes (60°S) than that (40–45°N)of the SAFZ in the Northern Hemisphere, and the SST and
its gradient are partly similar to those of the SAFZ. The
baroclinic transport associated with the ACC-SAF is estimated to be 53 Sv through the Drake Passage (Cunningham et al. 2003) and 105 Sv south of Australia (Sokolov
and Rintoul 2007), while the total transport at each section is 137 Sv and 147 Sv, respectively. Hence, the ACCSAF transport is comparable or larger than the transport of
K-SOJ. On its encounter with southward-flowing westernboundary currents of subtropical gyres, the ACC systemenhances the SST gradient around the Brazil-Malvinas
confluence and ARC (Table 2). The presence of the high
SST gradients is considered to be essential in anchoring
the storm track and polar front jet of the overlying atmosphere (Nakamura et al. 2004; Nakamura and Shimpo 2004;
Hotta and Nakamura 2011; Ogawa et al. 2012), as could beseen in the KE region. The air–sea coupling associated withthe storm tracks in turn can contribute to the maintenanceof the fronts in both the ACC and KE regions through the
downward transfer of westerly momentum and driving the
ACC and oceanic gyres (Nakamura et al. 2004).

13

482

4 Kuroshio–Oyashio confluence region
In the region north of the KE, eddies, filaments, and jets,
separated or bifurcated from the main streams of the Kuroshio and the Oyashio, merge and re-diverge frequently. We
will refer to the region where waters that originate fromeach current mix, from the SAC to the KE, as the KOC
region (Fig. 2).1 While there are various terminologies used
to describe the region, we will make an effort to follow thenomenclature provided by Favorite et al. (1976). Footnotes
are added where different names have been used to indicate
the same hydrographic feature discussed in this section.
Readers are referred to Yasuda (2003) for a detailed reviewof the hydrographic structure in this area.
Though hydrographic features in this region are complicated, there are several distinct fronts and jets (Figs. 4, 6;
Table 3). The broad region of large meridional gradients of
SST [about 1.5 °C (100 km)−1 around 40°N (WOA09), seealso Fig. 4b; Table 3] and SSS [about 0.2 (100 km)−1, see
Yuan and Talley 1996) in the northern part of the KOC
region is called the SAFZ.2 In synoptic sections, the northern (southern) boundary of the SAFZ is marked by the outcrop of the 33.0 (33.8) isohaline, which forms the top (bottom) of the subarctic permanent halocline (Yuan and Talley
1996; see also Fig. 6). The SAFZ corresponds to the boundary between the wind-driven gyres of the North Pacific(e.g., Hurlburt et al. 1996). The SAFZ contains a couple ofsharp SST and SSS fronts (Fig. 4b) with the SST and SSS
gradients compensating each other in terms of densitychange. The SST gradient in the SAFZ is shown to haveappreciable influence on the large-scale atmospheric circulation extending into the upper troposphere in the timemean (Nakamura et al. 2004; Taguchi et al. 2009) as wellas time-varying states (Frankignoul et al. 2011; Taguchi
et al. 2012; Smirnov et al. 2015). The near-surface baroclinicity maintained by differential heat exchanges acrossthe oceanic front anchors the storm track (Nakamura et al.
2008; Sampe et al. 2010; Hotta and Nakamura 2011), andits modulation associated with latitudinal shifts of the oceanic fronts exerts eddy feedback forcing on the mean
atmospheric circulation (e.g., Nakamura et al. 2004, 2008;
Taguchi et al. 2009, 2012; O’Reilly and Czaja 2014).

1This name is also used by Sugimoto and Hanawa (2011), but manyother names have been used to indicate this region, for example, thePerturbed Area (e.g., Kawai 1972), the Mixed Water Region (e.g.,
Talley et al. 1995), the Kuroshio-Oyashio interfrontal zone (e.g., Yasuda et al. 1996), and the Kuroshio-Oyashio Extension region (e.g.,Schneider et al. 2002), etc.
2
The SAFZ almost coincides with the region conventionally termed
the Transition Domain in the oceanographic community (Favorite
et al. 1976), which is bounded by the SAC and the subarctic boundary, which are described later in this section (see also Fig. 6).

13

S. Kida et al.

The SAC or extension of the Oyashio and the northernbranch of the KE or KENB3 are well known as the persistent jets in the KOC region. The SSH gradient suggests frequent bifurcation and confluence of jets (Fig. 4a). Nevertheless, the pathways of these two currents may be traced fromthe east coast of Japan as far east as 170°E by using subsurface temperature proxies. In the northern part of the KOC
region, the 4 °C isotherm at 100 m depth (Favorite et al.
1976) traces the southern boundary of the SAC well in thewestern and central portions of the North Pacific. This isotherm is the formal proxy of the southern limit of the “pure”
subarctic stratification with a subsurface temperature minimum underlain by a temperature maximum (Belkin et al.
2002), which is termed as the Subarctic Front (SAF) by
Ueno and Yasuda (2000) and Yasuda (2003).4 Further, recent
studies based on satellite observations have identified two
remarkable quasi-stationary jets in the SAFZ immediately
south of the SAC [Isoguchi et al. 2006; J1 and J2 in Figs. 4aand 6; see also Wagawa et al. (2014) for a recent hydrographic observation]. Across these quasi-stationary jets
(Fig. 4b) the SST gradients exceed 4–5 °C (100 km)−1,
which marks the strongest gradients within the SAFZ andtends to be stronger than that associated with the KE(Nonaka et al. 2006).
In the southern part of the KOC region, the KENB originates from the KE, which can be identified by the 6–8 °C
isotherms at 300 m depth (Mizuno and White 1983). Fromaround 160°E, the KENB merges with another large SSH
gradient, which extends eastward from about 150°E, 39°N
(Fig. 4a). The latter roughly coincides with the subarctic
boundary,5 which is defined by the near vertical 34.0 isohaline in the surface layer within 38°–40°N (Favorite et al.
1976). Zhang and Hanawa (1993) showed that the KENB
broadly corresponds to the subarctic boundary east of
170°E. Note that the subarctic boundary approximates thesouthern boundary of the SAFZ (Fig. 6b). Salinity increases
with depth north of the subarctic boundary as in the subarctic North Pacific. The area between SAC and the subarctic
boundary, which is called the Transition Domain (Favorite
et al. 1976), has unique water properties much warmer andsaltier than subarctic water.
Within the SAFZ, the SST fronts are observed throughout the year even in summer (Fig. 4c, d; see also Nakamura and Yamagata 1999). They become most prominent
in winter and early spring. In winter, turbulent latent andsensible heat fluxes are locally maximized on the southern
3

The northern branch of the KE (Mizuno and White 1983) is also
called the Kuroshio Bifurcation front (Yasuda 2003) or the Kuroshio
northern branch (Hurlburt and Metzger 1998).
4
Roden et al. (1982) and Belkin et al. (2002) termed this front “thePolar Front.”
5
Roden et al. (1982) termed this boundary “the Subarctic Front.”

483

Oceanic fronts and jets around Japan
Fig. 6  Latitude-depth sections
along 165°E of a potential
temperature (°C), b salinity, and c potential density σθ
(kg m−3) based on WOCE
Global Hydrographic Climatology (Gouretski and Koltermann2004). Proxies for the locationsof fronts used in Fig. 4b are
also indicated by marks. The
crosses mark the position of KE
defined by the 14 °C isotherm
at 200 m depth and the 12 °C
isotherm at 300 m depth. The
circles mark the position of
KENB defined by the 6 and
8 °C isotherms at 300 m depth.
The sharp marks the position of
the subarctic boundary definedby the salinity of 34.0 at 100 m
depth. The asterisk marks the
position of the SAC defined by
the 4 °C isotherm. The outcropof isohaline of 33.8 (33.0) in b
defines the southern (northern)boundary of SAFZ

(a)

(b)
SAB
SAFZ

(c)

KE

flank of the quasi-stationary jets (Tomita et al. 2011; seealso Fig. 4e, f), sharply declining northward across thejets (J1 and J2 in Fig. 4a). The surface baroclinic jets canbe identified throughout the year, but positions of SAC
and KENB vary significantly on short time scales owing
to the effect of mesoscale eddies. In relation to the hugeamount of wintertime heat release into the atmospherein the vicinities of these fronts (Fig. 4e, f), several kindsof pycnostad, or mode water, are identified in the KOC
region. Readers are referred to Oka and Qiu (2012) for a

KENB

SAC

review. Reproducibility of mode waters in this region is agood metric for evaluating the fidelity of model simulations. Mode waters in model simulations may be evaluated by comparing the distribution of the isopycnic potential vorticity (PV) (Fig. 5 of Suga et al. 2004), subduction
rate (Fig. 5 of Suga et al. 2008), and θ–S relationship forthe deep winter mixed layer in the mode water formation
region (Fig. 2 of Oka et al. 2011) with observations.
On interannual to decadal time scales, the latitudinal
positions of SAC and KENB vary by about ±2° (Belkin

13

484

et al. 2002; Nakamura and Kazmin 2003; Nonaka et al.
2006). The dual jets identified by Isoguchi et al. (2006) are
quasi-stationary in their flow directions, but the western jet(J1) varies on decadal scale synchronously with the meridional migration of KE (Wagawa et al. 2014). Isoguchi et al.
(2006) suggested a possibility of bottom topography effects
on the steady direction of the current, although no obvious
interaction of the jet with the bottom topography was foundin a snapshot observation (Wagawa et al. 2014).
The westernmost portion of the front that traces the southern boundary of the SAC is specifically called the Oyashio
front, which marks the southern edge of the western subarctic water where subsurface salinity is <33.6. Kawai (1972)
thus suggested the 33.7 isohaline at 100 m depth as anindex of the Oyashio front. In terms of temperature, Kawai(1972) suggests a 7 °C isotherm in January, 6 °C in February,
5 °C from March to May, 6 °C from June to August, 7 °C
from September to October, and 8 °C from November toDecember at 100 m depth. As discussed by Yasuda (2003),
the Oyashio front is known to have two branches intruding
southward adjacent to the east coast of Japan, exhibiting
large seasonal and interannual variations. We will not pursue
metrics for the Oyashio front beyond these indices owing to
its overwhelmingly complex behavior.

5 Fronts in the subtropical North Pacific5.1 Northern, Southern, and Eastern subtropical fronts
An STF is a subsurface temperature and density front at
depths of about 50–300 m, accompanying a steep northward shoaling of the upper pycnocline (Fig. 7a). The slopeof the upper pycnocline gives rise to the eastward current
shear of the Subtropical Countercurrent (STCC) near thesurface in opposition to the barotropic westward flow predicted by the Sverdrup theory at these locations. The STCC
is a shallow current and is usually found above the core
depth of the STF (Fig. 7b).
There are three STFs in the North Pacific (Fig. 2), andthey are basically maintained by different mode waters in
the ventilated thermocline (Kubokawa 1999; Aoki et al.
2002; Kobashi et al. 2006; Kobashi and Kubokawa 2012).
The STCCs associated with the Northern and Southern
STFs show a distinct seasonal cycle in velocity with amaximum in spring and minimum in autumn, whereas the
STCC associated with the Eastern STF exhibits no significant seasonal cycle (Table 4). Each of the STFs forms a vertically sheared current system between the STCCs and anunderlying westward flow, which favors high eddy activity.
The EKE undergoes significant seasonal modulation witha maximum in May and minimum in December through13

S. Kida et al.

January, reflecting the seasonal change in the dynamical
stability of the sheared current system (Qiu 1999; Kobashi
and Kawamura 2002), which is modified by the presenceof the SST front in the mixed layer (Chang and Oey 2014).
The EKE also changes considerably on interannual anddecadal timescales (Qiu and Chen 2010b, 2013) because of
the large variability of the STCCs (Table 4).
In some literature, STFs refer to SST fronts (e.g., Roden1980). Different from the subsurface STFs, SST fronts are
relatively broad and form in response to surface Ekman
convergence and/or inhomogeneity in surface heat fluxes(Kazmin and Rienecker 1996; Dinniman and Rienecker
1999; Qiu and Kawamura 2012). Satellite data reveal a
local peak of the SST gradient embedded in a broad surface
frontal zone along the Northern and Eastern STFs. The surface front exhibits the strongest SST gradient in February
through March and the weakest in August through September. Probably anchored by the subsurface STFs, the SST
front enhances the overlying atmospheric baroclinicity andthereby influences surface wind stress curl and precipitation (Kobashi et al. 2008). The Northern and Eastern STFs
also accompany a weak but notable local peak of upward
sensible heat flux in spring along the northern front and in
winter and spring along the eastern front (Fig. 8). Somecoupled general circulation model (GCM) studies indicate that the thermal advection of STCCs may yield SST
anomalies leaving imprints on surface heat flux and surface
winds (Xie et al. 2011).
Salinity is considered to make a minor contribution to
STFs because of relatively warm water in the subtropics.
At the surface, the Northern and Eastern STFs are roughly
located along the maximum of high-salinity Tropical Water,whereas the Southern STF is along the southern flank ofTropical Water. The salinity gradient does not exhibit significant seasonal changes at the surface (Table 4). At subsurface depths, the Northern and Eastern STFs are observed
along the northern boundary of Tropical Water (Uda andHasunuma 1969; Kobashi and Kubokawa 2012), where thesalinity front could weaken the density front, because temperature and salinity have opposing effects on density.
5.2 Hawaiian Lee Countercurrent
The HLCC (Flament et al. 1998) is a narrow surface eastward current, extending far westward from Hawaii against
the easterly trade winds and broad westward North Equatorial Current (Fig. 2). The western end of the current has
significant interannual variations. Climatologically, it isfound around 170°W–175°W from drifting buoy observations (Qiu et al. 1997; Yu et al. 2003; Lumpkin and
Flament 2013; Abe et al. 2013) and a bit further westward over the international dateline to about 165°E from

485

Oceanic fronts and jets around Japan(b) Ug(a) Sigma theta
0

0

0

−12

26

400
26

500

27

600

−4

500

−4

−4

400

0

−8

−4

−4

300

−4

300

0

25

−4

200

−4

Depth (m)

25

−4

24

26

−8

100

23

200

−16

100

600
5

10

15

Lat

20

25

30

5

10

15

Lat

20

25

30

1.5
2.0
2.5
Meridional gradient of density (10−6 kg m−4)
Fig. 7  Mean sections of a potential density and b zonal geostrophic
velocity (referenced to 1000 dbar) along 137°E, calculated from CTD
observations every winter and summer season during 2000–2009,
conducted by the Japan Meteorological Agency. Shade indicates themeridional gradient of potential density. Stations along the sections
are located approximately 1° apart in latitude. The locations of theNorthern and Southern STFs (STCCs) are denoted by arrows. Modified from Kobashi and Kubokawa (2012)

altimeter observations (Sasaki et al. 2010; Abe et al. 2013)and hydrographic observations (Kobashi et al. 2006). The
current is accompanied by subsurface temperature and density fronts, but the SST along the current is warm comparedto that of surrounding water because of the warm water
advection from the west (Fig. 8b).
Although the HLCC flows at almost the same latitudes
as the STCC, which is associated with the Southern STF(Fig. 2), the generation mechanism is quite different. The
HLCC is a basically wind-driven flow and is generated by
the wind wake near Hawaii (Xie et al. 2001). The Rossby
waves, triggered by the dipole pattern of the surface windcurl near Hawaii, propagate westward to induce the front
with surface eastward current. The extent of more than a
few thousand kilometers is the longest among the currents(with fronts) triggered by the orographic wake on earth.
The speed of the HLCC varies seasonally, remaining strong
from summer through winter and weakened in spring (Kobashi and Kawamura 2002), induced by meridional migration of the trade winds (Sasaki et al. 2010). Interannual
variations of the current are attributable to variations of
the northward extent of low-PV water (Sasaki et al. 2012b)
and/or to westward Rossby wave propagations from theeast of Hawaii triggered by wind anomalies (Abe et al.
2013). However, further studies are needed to reveal thedominant mechanism of the variations.

The SST maximum associated with the HLCC creates
upward sensible and latent heat fluxes (Table 4; Fig. 8b)
and induces surface wind convergence and high cloudwater over the HLCC (Xie et al. 2001). The magnitudeof the heat fluxes is comparable to that along the STFs
(Table 4). The wind convergence, together with the
Coriolis force, enhances (weakens) the easterly tradewinds to the north (south), which may shift the SST
maximum southward via the thermal feedback of latent
heat release (Xie et al. 2001). A comparison of coupled
GCM experiments with and without the SST maximum
further suggests that the HLCC is accelerated, extended
further westward, and shifted southward by the dynamical feedback of Ekman suction that is associated withthe positive wind stress curl along the HLCC via thethermal wind balance (Sasaki et al. 2013). Such dynamical feedback associated with the increase of HLCC
can potentially intensify the seasonal and interannual
variations.
The HLCC, as well as STCCs, forms a verticallysheared current system with underlying westward flow.
Reflecting seasonal variations of the baroclinicity of thesheared current, the EKE tends to be high from spring toearly summer and low in winter (Kobashi and Kawamura
2002; Yoshida et al. 2011). The EKE also varies on interannual and decadal time scales. A high EKE was observed

13

486
Fig. 8  Mean SST (contours) as
well as high-pass filtered sensible heat flux (color; W m−2) ina April and b October. SST is
contoured every 1 °C with thindashed contours of 0.5 °C intervals over 27 °C. The high passfiltering is done by subtracting a
10° moving average in latitude
from the original data to remove
the large-scale background
distribution. Thick dashed lines
indicate the climatological location of the STFs and HLCC

S. Kida et al.

(a) Apr mean SST and high−pass filtered SHF
35˚N
30˚N

20

20

20

25˚N

25

25

20˚N

25

15˚N

25

10˚N
120˚E

140˚E

160˚E

180˚

160˚W

140˚W

120˚W

100˚W

120˚W

100˚W

(b) Oct
35˚N

25

25

30˚N

25

25˚N
20˚N

25

15˚N
10˚N
120˚E

140˚E

160˚E

in 1993–1998 and 2002–2006, while the EKE was low
in 1999–2001 and 2007–2009. During the positive PDO
phase, cold SST anomalies north of the current, induced
by enhanced heat loss to the atmosphere, augment thevertical shear and thereby enhances the EKE (Yoshida
et al. 2011).

6 Fronts in coastal and marginal seas
6.1 Tidal fronts of the Seto Inland Sea
There are two types of thermal fronts in the Seto Inland Sea.
One is a tidal front in summer, and the other is a thermohaline front in winter (Yanagi and Koike 1987). As reflected
in its name, a tidal front is associated with tidal currents. It
is produced by surface warming in coastal seas, where spatial variations in tidal currents are strong. Water is vertically
mixed where tidal currents are strong, whereas water isstratified where tidal currents are weak. Between these tworegions, a tidal front thus forms with a pronounced watertemperature gradient (Simpson and Hunter 1974).
A number of tidal fronts are observed in the Seto Inland
Sea (Takeoka 2002). As a typical example, a front aroundHayasui Strait (Yanagi and Koike 1987) is included in

13

180˚

160˚W

140˚W

Table 5 (see also the inset panel of Fig. 2). The particular
front can be identified from late April to early September.
The front has a length of about 40 km and width of about
5 km across which the SST changes by 2 °C and the SSS
by 0.3. The SST gradient undergoes a fortnightly variation:
being stronger in the spring tide than in the neap tide (Yanagi and Koike 1987), whose dynamical explanation is given
in Sun and Isobe (2006).
Among thermal fronts with different spatial scales (Yanagi 1987), tidal fronts are one of the smallest types of oceanic fronts that can influence the overlying PBL structure(Shi et al. 2011). The surface wind speed reduces when the
air travels from the warm to cooler water and vice versa,
which is mainly because of the perturbation pressure gradient force with the SST gradient and not the stratification
control on the turbulent vertical mixing of wind momentum(Shi et al. 2011).
6.2 Japan Sea subpolar front
The JS-SPF (Fig. 2; Table 5) is the boundary between thewarm and salty water mass entering from the Tsushima
Strait and the cold and fresh water mass formed within
the Japan Sea (Isoda et al. 1991; Talley et al. 2006). This
front is sometimes referred to as the polar front (e.g.,487

Oceanic fronts and jets around Japan

Isoda 1994), and the presence of the JS-SPF has long
been recognized (Uda 1938), owing possibly to its prominent frontal signal extending zonally across the Japan
Sea at 40°N throughout the year. Satellite observations
show bifurcations at its eastern and western ends, which
are near the coasts and where variability is stronger thanin the interior (Park et al. 2004). Due to the depth of theJapan Sea, the JS-SPF shares many common featureswith its counterpart in the open ocean. For example, thefront is considered a region of intense subduction andthus intermediate water formation. This has led to intense
studies on the mechanism of frontal subduction during thelast decade or so (Thomas and Lee 2005; Yoshikawa et al.
2001; Lee et al. 2006).
Although relaxed substantially in summer, the SST
gradient across the subpolar front is strong particularly
in winter, locally observed to reach as high as 5 °C over
5 km at times (Lee et al. 2006), with climatological values comparable to or even stronger than that across theKE (Table 1). Likewise, upward turbulent heat fluxes
at the surface are strong in winter, especially to thesouth of the subpolar front (Fig. 4e, f), where direct
observations show the heat fluxes frequently exceeding 300 W m−2 (Lee et al. 2006). Adjacent to the Asian
continent, the large amount of heat loss from the Japan
Sea occurs under the monsoonal cold air outbreaks.
In addition to this strong surface cooling of the warm
water, strong cold Ekman advection across the sharp
SST gradient results in the destabilization of the nearsurface stratification in the ocean and contributes to thefrontal genesis (Thomas and Lee 2005; Lee et al. 2006).
As the response of the atmospheric PBL to the frontal
SST gradient and associated cross-frontal differential
heat release, the surface wind speed is observed to bestronger over the warmer water, contributing to intense
air–sea coupling in the region (Shimada and Kawamura
2006, 2008). Furthermore, Shimada and Kawamura(2006) showed an atmospheric response to the SST
front induced by the Tsushima Current located to thesouth of the JS-SPF.
The JS-SPF is also associated with an eastward jetobserved to be about 0.30 m s−1 at the surface (Isobe
and Isoda 1997). Observational analysis shows anannual mean transport of about 1.1–1.2 Sv (Chu et al.
2001), which is embedded in the general northeastward
flow of the Tsushima Current. The JS-SPF corresponds
to the boundary between this eastward flow to its south
and the cyclonic gyre to its north. Climatological EKE
does not exhibit a maximum along this front but rather
an increasing tendency toward its south (Jacobs et al.
1999). The JS-SPF is observed to vary interannuallyand decadally, which is possibly induced by instabilityof the jet, the anomalous in/outflows through the straits,

and basin-scale atmospheric variability (e.g., Isoda 1994;
Minobe et al. 2004). Recent studies suggest that the SST
variability in the Japan Sea may, in return, affect largescale atmospheric circulation (Yamamoto and Hirose2011), although fine-structure SST anomalies may not bequite relevant in forcing large-scale atmospheric anomalies (Seo et al. 2014b).

7 Summary and concluding remarks
This article reviewed the progress in our understandingof oceanic fronts around Japan that has accumulated in
the recent years and summarized the frontal properties inTables 1, 2, 3, 4, and 5. In addition to simply serving as areference list of the metrics, Tables 1, 2, 3, 4, and 5 canbe used for visualizing the characteristics of fronts. For
example, a scatter plot of the metrics using the SST gradient and flow speed shows how the frontal structure may
compare with other fronts (Fig. 9a). From the thermal
wind balance, the surface flow speed should be stronger
as the density gradient increases or thickness of the front
increases, assuming the flow is weaker at depth. Therefore, if density is primarily a function of temperature, theupper part of the plot shows the fronts where thickness
is likely playing a large role, such as the major westernboundary currents. Fronts such as the JS-SPF show astrong SST gradient but weak flow speed likely because of
their shallow structure. Note, however, that salinity likely
plays a large role for the fronts in the KOC region andthat latitudinal differences need to be considered for moredetailed analysis. A similar scatter plot can be createdusing the metrics for the SST gradient and sensible heat
flux (Fig. 9b). Enhanced heat fluxes are expected in thevicinity of an ocean front when air parcels travel acrossthe front, since their temperature requires additional time
to adjust to the underlying SST (see also Fig. 15 of Small
et al. 2008). The stronger the SST gradients become, thelarger the difference between the SST and surface atmospheric temperature (SAT) become, which then results inenhanced sensible heat flux. The scatter plot appears tosuggest such a trend, although there are also significant
differences among the fronts. The difference between the
SST and SAT may well be present on the large scale due
to factors such as large-scale wind speed and the distancefrom land, likely making the actual magnitude of the heat
fluxes differ depending on the front.
One of the metrics that is missing from Tables 1, 2,
3, 4, and 5 is the frontal variability on time scales longer
than the interannual scales. The KE, for instance, is known
to exhibit pronounced decadal-scale variability in frontal location and intensity. How much the metrics differ
between the varying states would be an interesting aspect

13

488

S. Kida et al.

Flow speed(a)

SST gradient (max month)

Sensible Heat Flux
(max month)(b)

SST gradient (max month)Fig. 9  a Scatter plot of the SST gradient (maximum month) andflow speed based on Tables 1, 2, 3, 4, and 5. The bars indicate themagnitude of interannual variability. Blues are the fronts in the North
Pacific. Greens are the fronts located elsewhere. b Same as a except
for the SST gradient (maximum month) and sensible heat flux (maximum month)

to compare in the future. While we hope the metric table
will be useful, two aspects need to be kept in mind. First,
redefinition of the metrics may be necessary when evaluating numerical models because model outputs often contain biases. For example, an eddy-resolving ocean GCM(OFES; Masumoto et al. 2004; Sasaki et al. 2008) successfully reproduces the KE and KOC region including threeseparated surface and subsurface temperature and/or SSH
fronts (e.g., Nonaka et al. 2006). However, the location of
the KE and KENB was better defined by the location at
300 m depth of the 14 and 10 °C isotherms rather than 12
and 8 °C (see Sects. 3.1.1 and 4 for the metrics of KE and
KENB), respectively, since the model had subsurface warm

13

bias in this region. Second, we need to be cautious not tooversimplify the metrics, although they are useful. Someof the conventional metrics are also based on temperature
because of its data availability and that they may have beenoriginally defined by other variables. For example, Kawai(1972) proposed the metric for describing the location of
the Oyashio front by the 33.7 isohaline at 100 m depth. To
interpolate this isohaline from the temperature field, different isotherms needed to be chosen for each calendar month.
In some literature, however, the definition has been oversimplified by the use of the 5 °C isotherm at 100 m depth
throughout the year. This is clearly no longer the same as
the original definition proposed by Kawai (1972).
Oceanic fronts are now reasonably well resolved inrecent high-resolution ocean GCMs and just beginning to
be resolved by high-resolution global air-sea coupled models. We envision that some of the coupled GCMs contributing to the next Assessment Report of the Intergovernmental
Panel on Climate Change (IPCC) will represent oceanic
fronts reasonably well. We need to assess what significant
impacts can be exerted in the mean state and variability ofthe model-simulated atmospheric and oceanic circulationby the improved representation of oceanic fronts. We hopethat the metric tables (Tables 1, 2, 3, 4, and 5) can provide a
useful tool for evaluating the fronts captured in the state-ofthe-art numerical models.
Acknowledgments We thank the editor, Dr. Shoshiro Minobe, andthree anonymous reviewers for their constructive comments on the
manuscript. The altimeter products are produced by Ssalto/Duacs anddistributed by AVISO with support from CNES (http://www.aviso.
oceanobs.com/duacs/). APDRC at the University of Hawaii at Manoa(http://apdrc.soest.hawaii.edu/) was also used to obtain some of thedata sets. This work is supported by the Japan Society for Promotionof Science through a Grant-in-Aid for Scientific Research on Innovative Areas 2205.

Appendix
The methods and data sets used for estimating the metrics shown in Tables 1, 2, 3, 4, and 5 are described here.
Common data sets used for the SST and its gradient aremonthly means of AMSR-E SST (Gentemann et al. 2010),
MGDSST (Kurihara et al. 2000), and OISST (Reynolds
et al. 2007) from 2003 to 2008. Monthly climatology of
WOA05 (Antonov et al. 2006), WOA09 (Antonov et al.
2010), and WOA13 (Zweng et al. 2013) is used for the
SSS and its gradient. Monthly means of AVISO (http://
www.aviso.oceanobs.com/duacs/) from 1993 to 2007
are used for the SSH gradient and EKE. The J-OFURO2(Tomita et al. 2010) from 1993 to 2007 is used for the surface heat fluxes. Specific data sets used for each front are
shown in Table 6, and detailed differences are described

489

Oceanic fronts and jets around Japan

Table 6  Data sets and references used for estimating the metrics associated with the SST, SSS, and their gradients, SSH gradient, and heat
fluxes in Tables 1, 2, 3, 4, and 5
KuroshioGS

AC and ARC

ACC-SAF

KENB, SAC, J1,
J2, and SAFZAMSR-E SST
WOA05
AVISO
J-OFURO2

Reynolds SST
WOA09Sallee et al. (2008)
J-OFURO2

AMSR-E SST
WOA13
AVISO
J-OFURO2

KE

K-SOJ

K-ECS

MGDSSTWOA13
AVISO
J-OFURO2MGDSST
WOA09
AVISO
J-OFURO2

PN line (JMA)PN line (JMA)AVISO
J-OFURO2

AMSR-E SST
WOA09
AVISO
J-OFURO2

NSTF, SSTF, and ESTF

HLCC

Seto Inland Sea tidal front

JS-SPF

SST, ΔSST
SSS, ΔSSS
ΔSSH

Reynolds SST
WOA09
AVISO

Reynolds SST
WOA05
AVISO

Yanagi and Koike (1987)
Yanagi and Koike (1987)
Takeoka (2002)

AMSR-E SST
WOA13
AVISO

Heat flux

J-OFURO2

J-OFURO2

Chang et al. (2009)

J-OFURO2

SST, ΔSST
SSS, ΔSSS
ΔSSHHeat flux

below. Other metrics are primarily based on past observational studies.
Kuroshio Extension
The location of the KE is based on the regions defined in
Qiu and Chen (2005). Since the behavior of the frontal axis
is significantly different between the upstream and downstream, two metrics are provided. One is the average over
the whole frontal axis and the other is that limited to the
upstream. The frontal axis is determined from the maximum flow speed based on the methods described in Ambe
et al. (2004). FRA-JCOPE2 utilizes the flow speed at 10 m
depth. SST and its gradient are estimated from MGDSST(1993–2007). SST is estimated by averaging the SST
50 km north and south of the frontal axis, and the SST gradient is estimated by taking the difference between the two.
However, we exclude the months when the KE axis is obviously obscured by the presence of a meso-scale eddy or the
method detects the KENB instead. SSS and its gradient are
estimated from WOA13 using similar methods as the SST
and its gradient, respectively. However, only the upstream
values are presented since WOA13 is a climatology andthus unable to capture the transient variability of the frontal axis downstream. Only the upstream values, where
variability is less, are presented. The SSH gradient is estimated from the monthly means of AVISO, but the annual
mean is presented since the magnitude of the SSH gradientdoes not change significantly with seasons. Other dynamical properties are based on observational studies. Transport and flow speed are both based on Howe et al. (2009),
which are observations from 146°E. Note that the transport
decreases significantly to about 75 Sv at 148.5°E. Depth isdetermined from WOCE hydrography (http://www-pord.
ucsd.edu/whp_atlas/pacific_index.html). The EKE is basedon the downstream and upstream values estimated fromsatellite altimetry in Qiu and Chen (2005). Heat fluxes are
estimated from the J-OFURO2 by taking a spatial average
between 141°–163°E and 34°–38°N to represent the whole
KE region and 141°–153°E and 34°–38°N to represent the
upstream KE region.
Kuroshio south of JapanThe location of the Kuroshio axis is estimated from themonthly mean SSH derived from AVISO. Using the SSH
gradient, the latitudinal position of the Kuroshio axis
is determined at each longitude with an interval of 0.25°
between 131° and 140°E. The SSH gradient is then calculated by taking a zonal average of the gradient at this Kuroshio axis. SST and its gradient are estimated from monthly
means of MGDSST using a similar method. Since the SST
front associated with the Kuroshio is not closely matchedwith the current axis as mentioned in Sect. 3.2, the positionof the SST front is determined from the maximum SST gradient within the region between 1° south and north of thecurrent axis defined from the SSH. The SSS and its gradient are estimated from monthly climatology of WOA09 andare determined as the maxima around the mean Kuroshio
axis. The length of the K-SOJ is estimated between 131°
and 140°E using the Kuroshio axis data set of Ambe et al.
(2004) from 1993 to 2007, which includes periods of bothstraight and large meander paths. Flow speed and transport
are from observations by Ambe et al. (2004) and Imawaki
et al. (2001), respectively. While the estimates of the Kuroshio transport tend to be contaminated by local phenomena
such as the Kuroshio path variations and mesoscale eddies,
the transport value presented in Imawaki et al. (2001) was
observed off-shore of Shikoku where the Kuroshio is relatively stable. The depth of the Kuroshio is based on observations by Book et al. (2002). The EKE is estimated fromthe monthly mean geostrophic velocity anomaly of AVISO.

13

490

Heat fluxes are estimated from the J-OFURO2 by taking a spatial average within 131°–140°E and 30°–35°N.
The dynamical properties, such as the SSH gradient, flow
speed, transport, and EKE, do not change significantly withseasons and thus are presented by their annual means.
Kuroshio along the shelf break of the East China Sea
The location of the K-ECS follows the current field given
by Qiu et al. (1990) with the exclusion of the Kuroshio
northeast of Taiwan because there is no apparent thermal
front there (HicKox et al. 2000). SST, SSS, and their gradients are based on quarterly hydrographic data along
the PN line in January, April, July, and October from theperiod of 1980–2012, provided by the Japan Meteorological Agency (http://www.data.jma.go.jp/kaiyou/db/vessel_obs/data-report/html/ship/ship.php). The SSH gradientis estimated from the monthly means of AVISO with the
Kuroshio axis defined by the location of the maximum SSH
gradient between 124° and 127°E. Heat fluxes are also estimated along this Kuroshio axis from the monthly means of
the J-OFURO2. Flow speed, transport, and depth are basedon observations by Oka and Kawabe (1998). The EKE isbased on numerical simulations by Miyazawa et al. (2004).
Gulf Stream
All values are based on the GS after the separation at Cape
Hatteras, where the location of the GS frontal axis is basedon Sasaki and Schneider (2011): The jet latitude is defined
as a contour of −10 cm from absolute dynamic topography
data by Niiler et al. (2003b). SST and its gradient are estimated from the monthly means of AMSR-E SST, where thegradient is based on values 2° north of the frontal axis averaged from 60°W to 70°W, and only the maximum monthis presented. SSS and its gradient are estimated from theannual mean WOA09 averaged along the frontal axis andalong 2° north of the frontal axis, respectively, from 60°W
to 70°W. The SSH gradient and EKE are estimated fromAVISO and are also annual means. The SSH gradient is at
70°W, and the EKE is estimated averaged along the frontal axis from 60°W to 70°W. The transport is based on the
maximum value from in situ observations near 55°W by
Hogg (1992). The depth is based on observations by Johns
et al. (1995). Heat fluxes are estimated from the J-OFURO2
averaged along the frontal axis from 60°W to 70°W.
Agulhas Current and Agulhas Return Current
The location of the AC and ARC is based on Gordon (1985)
and Lutjeharms and Ansorge (2001), respectively. Metrics
related to SST, SSS, and SSH are estimated from AMSR-E
SST, WOA05, and AVISO, respectively. SST and SSS are

13

S. Kida et al.

those found along the frontal axis and are estimated fromthe monthly means for SST, but the annual mean for SSS.
The SST, SSS, and SSH gradients are maxima found along
the frontal axis in the annual means. Heat fluxes are estimated from the J-OFURO2 along the frontal axis for the
AC but maximum values within 15°E–60°E and 45°S–30°S
for the ARC. Flow speed, transport, and depth are all basedon observational studies by Bryden et al. (2005) for the AC
and Lutjeharms and Ansorge (2001) for the ARC. The EKE
is estimated from the monthly mean geostrophic velocity
anomaly of AVISO.
Antarctic Circumpolar Current‑Subantarctic front(ACC‑SAF)
The location of the front is based on Orsi et al. (1995). SST
and its gradients are estimated from the monthly mean
OISST data from 1982 to 2011, with a spatial resolutionof 1° by 1° (Reynolds et al. 2002). SSS and its gradient are
estimated from WOA09 and are annual means. These estimates are the circumpolar averages along the fixed frontal
location of Orsi et al. (1995). The dynamical properties are
based on observational studies; the SSH gradient is from Sallee et al. (2008), flow speed is from Hofmann (1985), transport is from Rintoul and Sokolov (2001) and Cunningham
et al. (2003), depth is from Tomczak and Godfrey (2003),
and EKE is from Patterson (1985). Heat fluxes are estimatedfrom the J-OFURO2 and are the circumpolar averages along
the frontal axis.
Kuroshio‑Oyashio confluence region
Metrics for KENB, SAC, J1, and J2 are estimated byaveraging the respective parameters along the frontal
axes associated with the location metrics in Table 3. The
KENB axis is indexed by a straight line that connects its
location metrics, which lies between 6 and 8 °C at a depthof 300 m, so as to be consistent with Mizuno and White(1983). The metrics are estimated within 153°–170°E,
where the axis is distinctly separated from KE (Fig. 4a).
The SAC axis is defined by 4 °C isotherms at 100 m depth(Favorite et al. 1976; Belkin et al. 2002). The metrics of
SAC are estimated within 150°–155°E along the 4 °C isotherm as a typical value where the SST gradient is maximum [typically >3 °C (100 km)−1 in winter], although
the front extends farther eastward. J1 and J2 are indexed
by straight lines that connect their location metrics; theselines are located immediately southward of the maximum
SST gradients [typically >3 °C (100 km)−1 in winter] in
the SAFZ north of 40°N (Fig. 4c). The SST and its gradient for KENB, SAC, J1, and J2 are estimated from themonthly means of AMSR-E SST (Figs. 4c, 3d). SSS and
its gradient for KENB, SAC, J1, and J2 are estimated from

491

Oceanic fronts and jets around Japanthe monthly climatology of WOA13. The SSH gradientsand surface currents are estimated from the absolute SSH
of AVISO (e.g., Rio et al. 2011). The EKE is based on
the monthly mean surface geostrophic anomaly. Depths
for KENB, J1, and J2 are based on Isoguchi et al. (2006).
Transport for SAC is based on Ohtani (1970). Estimatesof transport for other fronts are estimated from WOA13
based on geostrophy referenced to 1,500 m depth. Eastward flows that are parallel to the frontal axis are used.
We are currently not aware of past studies to compare
these values with, but the geostrophic velocity estimatesnear the surface match reasonably well with those of Isoguchi et al. (2006), as well as with those calculated fromthe AVISO SSH, and thus consider the transports reasonable to the first order. The difference between SAC andJ1 is unclear in observational data sets where spatialsmoothing is applied. Heat fluxes are estimated from the
J-OFURO2.
For the metrics of SAFZ, SST is estimated along the linethat connects (145°E, 40°N) and (180°, 45°N). SSS and
its gradient are based on WOA13, and the heat fluxes are
based on the J-OFURO2 along the same line. The SST gradient within the SAFZ is characterized by those of J1 andJ2 and thus is not presented here.

calculated from the monthly mean each year between 145°E
and 165°E for the N- and S-STFs and 175°W–155°W for
the ESTF, and then the monthly climatology is estimated.
We extracted a meridional peak of the velocity and its associated SSH gradient for each month. Since the ESTF showsno significant seasonal cycle, its metrics are estimated fromthe annual mean. The transport and depth are left blankbecause these metrics need information about the verticalextent of the fronts, which may not necessarily be easily
accessible from observations. Synoptic ship observations(e.g., Aoki et al. 2002) appear to be inconsistent with altimeter-derived climatology when compared at the surface (seea review paper by Kobashi and Kubokawa 2012), probably
because of observation sparsity and eddy contamination.
The EKE is calculated from the monthly means of the geostrophic velocity anomaly. We first computed the monthly
spatial averages of the EKE each year within 23°–25°N and
145°–165°E for the NSTF, 18°–20°N and 145°–165°E for
the SSTF, and 25°–27°N and 175°–155°W for the ESTF,
and then estimated the monthly climatology. Heat fluxes
are assessed in the same way as the EKE using the monthly
means of the J-OFURO2.

Northern, Southern, and Eastern Subtropical fronts(NSTF, SSTF, and ESTF)

The location of the HLCC front is determined by examining the surface zonal geostrophic velocity relative to 400 db
in climatological hydrographic observations (Kobashi et al.
2006). The western end changes interannually from the eastof the international dateline to about 165°E from altimeter
observations (Sasaki et al. 2010; Abe et al. 2013). When the
HLCC and SSTF influence each other, the western extent
of the HLCC cannot be determined clearly. SST is estimated from the monthly means of the OISST data set from2003 to 2008 (Reynolds et al. 2007). The SST gradient is
not provided since the SST shows a maximum along the
HLCC. The SSS and its gradient are based on the monthly
climatology of WOA05. The SSH gradient and surface
flow speed are estimated from AVISO reference series ofdelayed-time MADT from 1993 to 2007 assuming geostrophy between 180 and 160°W at 19.5°N. The depth andtransport are determined between 180° and 160°W with thetransport estimated from the numerical simulations of the
OFES Quikscat run (2001–2008) (Masumoto et al. 2004;
Sasaki et al. 2008). The EKE is determined from Yoshida
et al. (2011) between 170.0°E–160.0°W and 17.0°–21.7°N,
which is based on AVISO SSH. Annual mean is presented with its variability estimated based on 1 year running weakly means. The seasonal EKE level peaks inJune (0.025 m2 s−2) and drops to a minimum in January(0.015 m2 s−2). Heat fluxes are estimated from the monthly
J-OFURO2 from 2002 to 2007.

The locations of the STFs are determined from the AVISO
reference series of the delayed-time Maps of Absolute
Dynamic Topography (MADT) product from 1993 to 2007.
STFs are defined as a continuous band of eastward surface geostrophic velocity in the monthly climatology. The
climatology is still a bit patchy, probably because of higheddy activity, but the positions of the STFs show no clear
seasonal variations. The eastern end of the SSTF is determined based on the climatology map of subsurface temperature fronts (Kobashi et al. 2006) because of the presenceof the HLCC there. SST is estimated from the monthly
means from an optimally interpolated SST data set produced from the blend of infrared and microwave satellite
observations and in situ measurements from 2003 to 2008
(Reynolds et al. 2007). SSS is estimated from the monthly
climatology of WOA09. SST, SSS, and their gradient are
estimated at the current axis of the STCCs after calculating
the zonal averages.
The flow speed is estimated from the surface eastward
geostrophic velocity since the meridional velocity blurs thecharacteristic of the STFs. The meridional velocity is generally related to the basin-scale subtropical gyre circulationand has the same order as the zonal one at STFs. To estimate the metrics, zonal averages of zonal velocity are first

Hawaiian Lee Countercurrent

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492

Seto Inland Sea tidal front
The location, SST, SSS, and their gradients are based on
observations during April by Yanagi and Koike (1987). The
SSH gradient, flow speed, transport, depth, and EKE arebased on the general knowledge of tidal fronts found in the
Seto Inland Sea from observations and numerical models(e.g., Takeoka 2002; Chang et al. 2009). Heat fluxes are
based on numerical simulations (Chang et al. 2009; Shiet al. 2011) for a limited period; thus, their interannual variability is not provided.
Japan Sea subpolar front
The location of the front is based on Park et al. (2004)
where the JS-SPF does not include major bifurcations. The
frontal axis is determined from the maximum SST gradient near 40°N from the monthly means of AMSR-E SST.
SST and its gradient are the averages along this frontal
axis. SSS and its gradient are estimated from the monthly
climatology of WOA13 based on the monthly climatological position of the frontal axis. The SSS gradient is presented with a sign since the sign changes during the year.
Positive values represent higher salinity toward the north.
The SSH gradient is also estimated along the frontal axis
using the monthly means of AVISO (2003–2008). Flow
speed is based on direct observations by Isobe and Isoda(1997), and transport is based on inverse calculation using
the observational data set by Chu et al. (2001). Depth isbased on the observational analysis of Isoda (1994) and
Minobe et al. (2004). The EKE does not show a clear
maximum along the front (Jacobs et al. 1999) and thus is
not presented. Larger values are found towards the southof the front, where the Tsushima Current exists, and the
subpolar front appears to be much of a transition area fromhigh (south) to low (north) EKE areas. Heat fluxes are estimated along the frontal axis using the monthly means of
the J-OFURO2 (2003–2008).
FRA‑JCOPE2
The Japanese Fishery Research Agency (FRA)-Japan
Coastal Ocean Prediction Experiment (JCOPE) (FRAJCOPE2, Miyazawa et al. 2009) uses the JCOPE2 ocean
model, which is based on the Princeton Ocean Model witha generalized coordinate of sigma (Mellor et al. 2002) andprovides daily mean ocean data covering the western North
Pacific (10.5–62°N, 108–180°E) with a horizontal resolution of 1°/12°. The model assimilates the remote-sensing
data of altimetry and surface temperature and in situ dataof temperature and salinity profiles including the data fromthe FRA using the 3D-VAR method. Surface momentum
and heat fluxes are calculated using the bulk formulae

13

S. Kida et al.

(Kagimoto et al. 2008) with atmospheric variables obtained
from the National Centers for Environmental Prediction/National Center for Atmospheric Research (NCEP/
NCAR) reanalysis (Kalnay et al. 1996). SSS is relaxed tomonthly climatological data (Conkright et al. 2002). The
FRA-JCOPE2 data set is provided by the Japan Agency for
Marine-Earth Science and Technology (http://www.jamstec.go.jp/jcope/).

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