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Sea Level Pressure Minimum along the Kuroshio and Its Extension*
YOUICHI TANIMOTO
Faculty of Environmental Earth Science, and Graduate School of Environmental Science, Hokkaido University, Sapporo,and Research Institute for Global Change, JAMSTEC, Yokohama, JapanTOMOHISA KANENARI
Faculty of Environmental Earth Science, and Graduate School of Environmental Science, Hokkaido University, Sapporo, JapanHIROKI TOKINAGAInternational Pacific Research Center, University of Hawaii at Manoa, Honolulu, Hawaii

SHANG-PING XIEInternational Pacific Research Center, and Department of Meteorology, University of Hawaii at Manoa, Honolulu, Hawaii(Manuscript received 9 September 2010, in final form 10 February 2011)
ABSTRACT
Atmospheric effects of sea surface temperature (SST) fronts along the Kuroshio and Kuroshio Extension(K-KE) are investigated by examining spatial characteristics of the climatological sea level pressure (SLP),
surface winds and surface heat flux (Q) fields based on an in situ observation dataset. A hydrostatic effect of
the SST front is observed during the northwesterly monsoon characterized by a westward-extending low-SLP
wedge (trough) slightly south of the peak in Q along the K-KE. Ageostrophic surface westerlies crossing SLP
isobars toward a trough center are found north of the low-SLP wedge apparently because of the eastward
acceleration by the vertical mixing effect. This feature of the ageostrophic winds is less pronounced south of it,a north–south asymmetry arising as vertical mixing accelerates (decelerates) eastward ageostrophic winds
north (south) of it. While the SLP trough near the SST front is found nearly year-round except for summer, its
meridional location seasonally migrates probably due to the change in the mean surface flow.
Regional atmospheric model experiments are conducted to examine the effect of the SST front on theoverlying atmosphere. The observed features in winter are adequately simulated when high-resolution SST isprescribed as the boundary condition. The strong Q along the K-KE and resultant SLP trough and ageostrophic surface winds are absent when the SST boundary condition is smoothed in space to weaken the SST
front. These results illustrate that the cross-frontal change in Q is of great importance to leave the SST imprint
on the overlying atmosphere via hydrostatic and vertical mixing adjustments.

1. Introduction
Atmospheric responses in the marine atmosphericboundary layer (MABL) to sea surface temperature (SST)* International Pacific Research Center Publication Number 764
and School of Ocean and Earth Science and Technology Publication Number 8105.

Corresponding author address: Dr. Youichi Tanimoto, Faculty ofEnvironmental Earth Science, Hokkaido University, N10 W5,
Kita-ku, Sapporo, Hokkaido 060-0810, Japan.
E-mail: tanimoto@ees.hokudai.ac.jp
DOI: 10.1175/2011JCLI4062.1
Ó 2011 American Meteorological Society

have been mostly investigated with a focus on the surface
wind changes across the SST fronts. Lindzen and Nigam(1987) propose a hydrostatic effect of the tropical SST on
the overlying sea level pressure (SLP) field throughchanges in air temperature in the MABL. The resultant
SLP gradient across the SST front causes wind adjustment at the sea surface, with the strongest wind response
found between the warm and cold SST. On the other
hand, Wallace et al. (1989) and Hayes et al. (1989) propose a vertical mixing effect in which warm (cold) SST
decreases (increases) the near-surface static stability
and thereby enhances (reduces) a vertical mixing in the
MABL. This effect induces the strongest wind adjustments

4420

JOURNAL OF CLIMATEon the warm flank of the SST front where the downward
transfer of large momentum from the upper MABL
accelerates the surface winds.
Small et al. (2003, 2005) examined SST effects on thesurface winds by analyzing the momentum budget ina regional atmospheric model (RAM). They found thatthe hydrostatic effect is a major factor in driving crossfrontal change in surface winds. Under the advection by
the mean southeasterly trade winds, the SLP gradient islocated downstream of the SST gradient, giving rise tostrongest winds on the downwind side of the SST front.
The analysis of buoy measurements by Cronin et al. (2003)
supports the hydrostatic effect on SLP. In situ observations
by Hashizume et al. (2002) and East Pacific Investigationof Climate Processes in the Coupled Ocean–Atmosphere
System (EPIC) campaign (Raymond et al. 2004), however,indicate that the vertical mixing effect is at work in theacceleration of the surface winds.
The study of the atmospheric responses to the SST
front has been extended to midlatitude SST fronts andmesoscale eddies (Xie 2004; Small et al. 2008). Owing tolarge temperature and moisture differences between ocean
and atmosphere, fronts anchor the local maximum of upward surface heat flux that can act to induce hydrostaticand vertical mixing effects in the MABL. Satellite measurements of the surface winds display divergence when
the winds go across the SST front and wind curl when theyblow along the fronts in the extratropics (O’Neill et al.
2003; Chelton et al. 2004; Vecchi et al. 2004; O’Neill et al.
2005; Tokinaga et al. 2005; Park et al. 2006). These windadjustments are consistent with the vertical mixing effect around the SST front. In contrast, the hydrostatic
adjustment is also suggested from analyses of the in situ
surface meteorological observations, satellite measurements, and operational and reanalysis datasets of thelower atmosphere. Over the warm flank of the Kuroshio
Extension (KE), Tokinaga et al. (2009) show a zonal
band of the local minimum in spatially high-pass-filtered
SLP anomalies in association with positive MABL temperature anomalies and vice versa over the cold flank of
it. More recently, Minobe et al. (2008, 2010) suggest thatthe hydrostatic adjustment along the Gulf Stream leads toconvective cloud formation and enhanced precipitation.
Song et al.(2006) show that, based on the RAM experiments, pressure gradient across the Gulf Stream front ismain force modifying the surface winds. Xu et al. (2010)
also detected a hydrostatic effect over the cool water pool
formed between the Pacific coast of Japan and the large
meander of the Kuroshio.
In the western North Pacific, tropospheric circulation
is highly monsoonal (Fig. 1), bringing about a distinct
seasonal cycle in surface heat flux. The present study examines how the monsoonal flow over the western boundary

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current region is modified when it crosses a midlatitude
SST front. By careful inspections of spatial characteristics in the SLP and surface wind fields during each season, we attempt to detect hydrostatic and vertical mixing
effects. These ocean imprints on the lower atmosphereare further examined using RAM experiments.
While the Japanese 25-yr Reanalysis (JRA-25) (Onogi
et al. 2007) still does not adequately resolve small-scalefeatures near ocean fronts, direct measurements of SLP
on merchant ships have been carried out along the busy
shipping lanes in the Pacific near Japanese islands (e.g.,
Fig. 2b in Tokinaga et al. 2009). The present study compares reanalysis and in situ measurements to study smallscale SLP features over the western North Pacific.
The rest of this paper is organized as follows. Section 2
introduces the datasets and model used in the present
study. Section 3 examines spatial characteristics of the
SLP and surface wind fields in the western North Pacificduring the winter and summer monsoon, respectively,
and make a comparison of those characteristics between
reanalysis and in situ measurements. Section 4 discusses
the ocean imprint near the SST front on the overlying
atmosphere based on RAM experiments. Section 5 issummary and discussion.

2. Data and modela. In situ observations
We use a monthly product of the International Comprehensive Ocean–Atmosphere Dataset (ICOADS) release 2.4 (Worley et al. 2005) from January 1950 to
December 2007, available on a 18 grid. Individual surface meteorological measurements from ships, buoys,
and other marine platforms are gridded after careful
quality controls. Further description of the data processing is documented in Tokinaga et al. (2009). We calculatemonthly climatology of SST, surface air temperature(SAT), surface wind velocity, latent and sensible surface
heat fluxes, and SLP by averaging data from 1950 to 2007
in the western North Pacific (208–508N, 1208–1708E).
We will examine the sum of latent and sensible surface
heat fluxes Q of ICOADS as an indicator of the heating of
the MABL from the ocean surface. The latent heat flux, in
addition to the direct sensible heat flux, may contribute to
the warming of the MABL in the Kuroshio and Kuroshio
Extension (K-KE) region where the midlatitude storms
are developed, accompanied with the zonal rainband
and low-level clouds. The ocean surface heat flux dataset
based on satellite measurements [Japanese Ocean Flux
Datasets with Use of Remote Sensing Observations
version 2 (J-OFURO2); Kubota et al. 2002; Tomita et al.
2010] provides the latent, sensible, and radiative heat fluxes.
Their seasonal climatology during 1993–2006 shows that,

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FIG. 1. Monthly climatology of SST [gray contours; contour interval (CI) 5 18C], Q (colors;
W m22), and surface winds (arrows; m s21) averaged over 1428–1478E based on ICOADS.
Note that the horizontal axis begins and ends in August. Color convention and scale of arrows
are represented on the right.

in the K-KE region, Q is the predominant term in themeridional maximum of the net surface heat exchange
from autumn through winter to spring (Fig. 2). In the
aerodynamic bulk formulae, sensible and latent heat
fluxes are a function of temperature difference between
ocean and atmosphere DTOA and moisture difference,
respectively. In the present study, we employ only DTOA
as an indicator of these differences because surface
moisture in the atmosphere strongly depends on local air
temperature (Kleeman and Power 1995; Tanimoto et al.
2003).
As stated in introduction, the prevailing surface windover the western North Pacific display a pronounced
seasonal cycle. In the monthly climatology of ICOADS,
the meridional component of extratropical surface winds
averaged over 1428–1478E changes its sign twice a year,
between March and April and between September andOctober, respectively (Fig. 1). Therefore, we will examine

northwesterly and southwesterly monsoons separately.
In the northwesterly (southwesterly) monsoon, threemonth means are calculated with October–December as
autumn and January–March as winter (April–June asspring and July–September as summer), following studies in the North Pacific (Suga and Hanawa 1995; Deseret al. 1996; Schneider et al. 1999; Tourre et al. 1999;
Sugimoto and Hanawa 2007).

b. Reanalysis products (JRA-25)
We employ the monthly JRA-25 from January 1979 to
December 2004, available on 1.258 grid (Onogi et al.
2007). From 2005 to 2007, we use the monthly product of
the operational analysis provided by Japan Meteorological Agency Climate Data Assimilation System (JCDAS),
which uses the same data assimilation system as in JRA-25
and has been used for near real-time climate monitoring(Onogi et al. 2007). The variables include wind velocity,

FIG. 2. Seasonal mean sensible (red), latent (blue), Q (dashed black), and Qnet (solid black) based on J-OFURO2 for (a) autumn,(b) winter, (c) spring, and (d) summer averaged over 1428–1478E.

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latent and sensible heat fluxes, and SLP. Monthly climatology is calculated by averaging from 1979 to 2007.

c. Regional atmospheric model experiment
To identify the ocean imprint on the overlying atmosphere, we conduct experiments using the Weather Research and Forecasting (WRF) model version 3.1 with
the Advanced Research WRF (ARW) dynamic core
(Skamarock et al. 2008) in a domain of the western North
Pacific (208–508N, 1208–1708E). The model resolution is
27 km in the horizontal and 43 sigma levels in the vertical,of which 18 levels are placed below 2 km high. We chosethe following set of the schemes: the asymmetric convective model version 2 (ACM2) planetary boundary
scheme (Pleim and Chang 1992), the Betts–Miller–Janjić
cumulus parameterization (Betts and Miller 1986; Janjić1994), the double moment 6-class microphysical scheme(Hong and Lim 2006), and rapid radiative transfer modeland Dudhia schemes for longwave and shortwave radiation calculations (Dudhia 1989; Mlawer et al. 1997),
respectively.
To examine the effect of the SST front, we conduct two
experiments. One is the control (CTRL) run by prescribing
10-day mean SST observed by Advanced Microwave
Scanning Radiometer for the Earth Observing System(EOS; AMSR-E) satellite measurements (Wentz andMeissner 2000). AMSR-E data is interpolated from
0.258 interval to the model’s 27-km grid. A daily highresolution SST dataset based on the Advanced VeryHigh Resolution Radiometer (AVHRR) infrared satellite measurements (Reynolds et al. 2007) is also used
to cover the coastal regions where AMSR-E SST data
are missing. An additional run uses spatially smoothed
SST (SMTH) to reduce frontal and small-scale structures in SST. For this SMTH run, we apply a 108 runningmean filter to the observed SST. We should note that theamount of enthalpy flux into the model atmospheremight not be same between the two runs, although thiskind of experimental design has been employed in the
numerical studies (Nakamura et al. 2008; Taguchi et al.
2009; Xu et al. 2010; Iizuka 2010).
We conduct four sets of 10-day integration separately
for each experiment, for 1–10 December 2002, 1–10
January 2003, 1–10 February 2003, and 1–10 March 2003.
We avoid long-term model integrations in which we may
have difficulty in separating the persistent ocean front’s
influence from the internal model drift in mean state.
Indeed, a 1-month integration displays a weak decreasing
drift in the SLP along the K-KE region. We additionallyconducted two 1-month runs for December 2002 with andwithout the 10-daily reinitializations. The December averages of these two runs also display nearly the same
pattern of surface heat flux, SLP, and surface wind. This

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result indicates that the 10-daily reinitializations neither
affect the certain patterns in those fields nor give a longterm drift in a continuous integration. We have conducted additional WRF experiments forced by daily SST
and by monthly mean SST for 1–31 December 2002. The
results are very similar in the surface heat flux, SLP, andsurface wind. This is because SST does not change thatmuch in observations over 10 days. Therefore, we will examine the WRF response to the 10-day mean SST during
each of months to identify the MABL adjustment to thepersistent SST feature. The 2002/03 winter is chosen because the surface heat exchange was very large in the K-KE
region on that winter. As for initial and lateral boundary
conditions, we employ the National Centers for Environmental Prediction (NCEP) Final Global Analyses on
a 18 3 18 grid. To adjust initial and boundary conditions,the model is initialized at 0000 UTC on one day before thefirst day of the 10-day integrations. The results in the experiments did not change even if we have the spinup period for 2 or 3 days. The lateral condition is updated atevery six hours. For these RAM experiments, wintermeans averaged over the four sets of 10-day integrations(a total of 40 days) are analyzed.

3. Observed low pressure wedge over the Kuroshio
and its extension
a. Autumn–winter season in ICOADSAs expected from the northwesterly monsoon during
autumn–winter (Fig. 1), the large-scale feature of the climatological SLP in the western North Pacific (contours in
Figs. 3a and 3b) presents a gradual westward increase.
While the large-scale northerlies or northwesterlies advect the cold and dry air mass from the continent in mostof the western North Pacific and its marginal seas, thelarge heat release (Q . 200 W m22) is found only along
the K-KE region from Taiwan through Japan to the east(color shading in Figs. 3a and 3b). Figure 1 displays thatthe steep meridional gradient of SST along the K-KE isfound at 358–428N, and that the Q maximum is located on
the warm flank of the SST front year-round. This is because the frontal structure in SST along the K-KE maintains large DTOA as observed in the 1428–1478E longitude
band (Figs. 3c and 3d) together with large moisture differences, even with scalar wind speed nearly uniform in themeridional direction. As found in Fig. 2, the Q maximum
accounts for the net heat exchange (Qnet) in the K-KE
region. The Q values decrease in the interior region to
the east as the frontal structure in SST weakens (Figs. 3eand 3f).
A closer look at the SLP distribution in the K-KE
region unveils a westward-extending SLP trough (white15 AUGUST 2011

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FIG. 3. (a),(c),(e) Autumn and (b),(d),(f) winter climatologies based on ICOADS: (a),(b) Seasonal-mean SLP (black contours; CI 5
1 hPa) and Q (colors; W m22). Supplemental dashed contours at 0.5-hPa intervals are plotted. White dashed line indicates the SLP troughin the K-KE region. Coloring convention is given to the right of each panel. (c),(d) Climatological SST (red; 8C), SAT (blue; 8C), surface
wind speed (green; m s21), Q (dotted black; 10 W m22), and DTOA (solid black; 8C) averaged over 1428–1478E. (e),(f) As in (c),(d), but forclimatologies averaged over 1608–1658E. Note that the axis for DTOA is at the top of the panels, but the axis for the other variables is at thebottom of the panels.

dashed line in Figs. 3a and 3b) where the wedge-likeshapes pointing to the west are represented by the region between the isobars of 1015.5–1016.5 hPa during
autumn and 1012–1016 hPa during winter. The center
axis of this westward-extending low-SLP wedge is displaced south of the meridional maximum in Q by about
18 latitude. The meridional phase relation of the lowSLP wedge and Q maximum is probably due to thethermal advection effect discussed later in the section.
While Tokinaga et al. (2009) have shown negative valuesof spatially filtered anomalies during the (December–
February) SLP climatology over the KE region basedon ICOADS (Tokinaga et al. 2009), we find that the

westward-extending low-SLP wedge in the actual (unfiltered) isobars can be observed throughout the northwesterly monsoon. This observational finding indicatesthe hydrostatic effect, in which the heat release from theocean acts to warm the overlying atmosphere, results inthe visible minimum of SLP via hydrostatic adjustment.
While surface northeasterlies are expected from thegeostrophic balance north of the low-SLP wedge, theobserved surface winds (black arrows in Fig. 4a) are stillwesterly across the isobars toward the SLP trough.
Ageostrophic flow (red arrows in Fig. 4a), defined as
residuals by removing the geostrophic wind from theobserved wind, tend to show a large westerly component

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FIG. 4. (a) Winter climatology of surface wind (black arrows; m s21), and ageostrophic wind (red arrows; m s21),
superimposed on SLP (gray contours; CI 5 0.5 hPa), and Q (gray shades; W m22) based on ICOADS. Ageostrophic
winds less than 3.5 m s21 are not represented. Shading convention and scale of arrows are represented to the right.
(b) Climatological meridional plots of winter means in ageostrophic zonal winds (solid red; m s21) and Q (dashedblack; W m22) averaged over 1428–1478E based on ICOADS. Note that the axis for Q is at the top of the panel, but
the axis for the ageostrophic zonal winds is at the bottom of the panel.

north of the low-SLP wedge where Q is strongest (Fig. 4b),
suggesting that enhanced vertical mixing brings thewesterly components aloft to the surface. While the vertical resolution of the JRA-25 is inadequate to represent
the MABL adjustment, we note that its climatologicalwinds at the 850-hPa level are westerly (not shownhere). In contrast, the ageostrophic winds south of the
low-SLP wedge are very weak (Fig. 4) even though Q
still exceeds 300 W m22. O’Neill et al. (2010) examinedthe effect of surface wind direction associated withmesoscale SST variability having spatial scale of 100–
1000 km and noted a clockwise rotation by 108 of thesurface wind direction over the positive high-passfiltered SST anomalies along the Kuroshio. This finding
is consistent with our result about northwesterly ageostrophic winds north of the low-SLP wedge on the warm
flank of the SST front.
Let us consider the momentum budget north and southof the low-SLP wedge separately as in the schematic viewof Fig. 5. It is a four-way balance among pressure gradient(white arrow in Fig. 5), Coriolis forces (blue arrow),
friction (yellow arrow) and vertical momentum mixing(red arrow). North of the low-SLP wedge, eastward acceleration by vertical mixing (red arrow) acts to turn theactual surface wind farther to the left of the geostrophic
winds and slightly increases wind speed (Fig. 5c). Thus, thesurface friction and vertical mixing are working together,leading to strong ageostrophic westerlies (green arrow).
In contrast, south of the low-SLP wedge, the surface friction and vertical mixing counteract each other with a small
ageostrophic component (Fig. 5d). Therefore, the observed meridional change in ageostrophic wind acrossthe K-KE front may be the result of both hydrostatic andvertical mixing effects working constructively north of
the front and destructively to the south.

b. Spring–summer season in ICOADS
The prevailing winds in the western North Pacificchange the direction from northwesterly to southwesterly during April (Fig. 1) when the subtropical high begins to form with a center at 308N, 1708E (Fig. 6a). This
subtropical high is well developed during summer (Fig.
6b). As expected from geostrophic balance, the surface
winds are mostly south–southwesterlies and their speeds
gradually weaken over the K-KE region from spring tosummer, indicating the weak warm advection over thesurface.
In the horizontal map of Q (Fig. 6a), we find the strong
surface heat flux regions over 75 W m22 along the K-KE
region from Taiwan through Japan into the KE. This ismuch like the northwesterly monsoon (Figs. 3a and 3b),
except with reduced amplitudes of smaller DTOA andslower surface wind speed in spring than in autumn andwinter (Fig. 6c). Unlike in the northwesterly monsoon,

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FIG. 5. Schematic figures of the surface momentum balance along the low-SLP wedge. (a),(b) Illustrated is theconventional friction effect on the geostrophic surface winds, which acts to push the air across the isobars from high to
low pressure both (a) north and (b) south of the low-SLP wedge. (c) An additional eastward acceleration by thevertical mixing north of the low-SLP wedge acts to turn the pressure gradient to its left in the surface momentum
balance. As a result, the actual surface wind will adjust by turning further left. (d) In contrast, south of the low-SLP
wedge, the eastward acceleration acts on the pressure gradient to its right in the surface momentum balance. The
wind will adjust by turning back toward the geostrophic wind.

however, the SLP field displays a low-SLP (1011.5–
1013.5 hPa) wedge that extends eastward from Taiwan
through Japan (white dashed line in Fig. 6a). As in thenorthwesterly monsoon, this eastward-extending wedge
corresponds to an SLP trough over the large heat flux region. Ageostrophic northeasterlies are pronounced southof the eastward low-SLP wedge where DTOA , 0:98C(Fig. 7). Because of this near-neutral condition of surface
static stability and associated weak surface heat flux, thevertical mixing is not as strong as during winter–spring
seasons. Ageostrophic winds indicate large surface dragunder near-neutral condition compared to unstableone. In the region under the unstable condition with
DTOA . 0:98C, the pressure gradient toward the west
is counteracting the possible eastward accelerationinduced by the vertical mixing. As a result, the ageostrophic winds weaken along the center axis of the lowSLP wedge. A distinct peak of Q is not found during

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FIG. 6. As in Fig. 3, but for (a),(c),(e) spring, and (b),(d),(f) summer climatologies based on ICOADS. (a),(b) Seasonal-mean SLP (black
contours; CI 5 1 hPa) and Q (colors; W m22). Supplemental dashed contours at 0.5-hPa intervals are plotted. White dashed line indicatesthe SLP trough in the K-KE region. Coloring convention is given to the right of each panel. (c),(d) Climatological SST (red; 8C), SAT(blue;
8C), surface wind speed (green; m s21), Q (dotted black; 10 W m22), and DTOA (solid black; 8C) averaged over 1428–1478E. (e),(f) As in(c),(d), but for climatologies averaged over 1608–1658E. Note that the axis for DTOA is at the top of the panels, but the axis for the othervariables is at the bottom of the panels.

summer (Fig. 6b), probably because strong solar radiation during summer weakens the frontal structure in SST
along the K-KE region (Fig. 6d). As observed during
autumn–winter seasons, no distinct peak of Q is present in
the interior region (Figs. 6e and 6f). The lack of a peak in
Q is not conducive to forming a local feature of the SLP.
Figure 8 shows the seasonal cycle of high-pass-filtered
Q and SLP anomalies, superimposed on the climatological surface winds averaged over 1428–1478E. For thehigh-pass filtering, we removed a spatial running meanof 58 latitude over the ocean. It illustrates the phase relation of the meridional maximum in Q and the SLP
minimum along the KE region. The Q maximum isanchored year-round at 35.58N by the large DTOA along
the KE, as in the horizontal maps (Figs. 3 and 6). In thenortherly or northwesterly monsoon during November–March, the SLP minimum is displaced southward by 18–28
latitude by temperature advection of these climatological winds. Because of weak advection by the weaknorthwesterlies during October and southerlies during
April–June, the SLP minimum is collocated with the Q
maximum.

c. JRA-25
Most in situ ship reports in ICOADS were instantaneously transmitted to the national weather services such as

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FIG. 7. Spring (April–June) climatology of surface wind (blackarrows; m s21), and ageostrophic wind (red arrows; m s21), superimposed on SLP (gray contours at 0.5 hPa interval), and DTOA
(gray shades; 8C) based on ICOADS. Ageostrophic winds less than
3.5 m s21 are not represented. Shading convention and scale ofarrows are represented to the right.

the Japan Meteorological Agency (JMA) and the National Ocean and Atmosphere Administration (NOAA),
and then incorporated into marine weather forecast.
Atmospheric reanalysis assimilates all archived reports,
and therefore is expected to represent the observed features, specifically over busy shipping lanes off the Pacificcoast of the Japanese Islands. However a westwardextending low-SLP wedge over the K-KE region is foundbut much less pronounced in JRA-25 (Fig. 9) than in
ICOADS (Fig. 3b). The intense heat flux over the K-KE
is adequately represented in JRA-25 regarding its distribution and amplitude. While we will not pursue the exact
cause of JRA-25 differences from ICOADS, we found, by
analyzing other reanalysis datasets, that this problem isgeneral, indicating problems of current assimilating
AGCMs in representation of physics, especially those for
MABL (Song et al. 2009).

4427

FIG. 8. Climatological seasonal cycle of high-pass-filtered SLP(contours; CI 5 0.1 hPa) and Q (gray shades; W m22) and surface
wind vectors (arrows; m s21) averaged over 1428–1478E, based on
ICOADS. Negative values are dashed. Note that the horizontal
axis begins and ends in August. Shading convention and scale ofarrows are given to the right.

Both CTRL and SMTH experiments using WRF(Figs. 10a and 10b) successfully reproduce the following
large-scale features of the wintertime SLP field over thewestern North Pacific as observed in the JRA-25 (Fig. 10c).
The SLP fields show a gradual westward increase northof 308N associated with the northwesterly monsoon anda zonal ridge extending from the Yellow Sea at 358N,
1208E toward 258N, 1708E. Besides the large-scale features in SLP, the CTRL experiment is successful in capturing the westward-extending low-SLP wedge within the
1011–1016-hPa isobars over the region from Taiwan
through Japan to the KE (Fig. 10a) where the strong Q
region with values over 500 W m22 is simulated ina narrow (;500 km) zonal band. The low-SLP wedge in
the CTRL experiment is located slightly south of the
maximum in Q as observed. The magnitude of Q in the
CTRL experiment is 1.5–2 times larger than in the JRA-25
owing to the stronger SST front in AMSR-E, while thewestward-extending low-SLP wedge is similar between

4. RAM experimenta. Low-SLP wedge
The previous section showed the westward-extending
low-SLP wedge displaced slightly south of the local maximum of Q and ageostrophic westerlies right over the Q
maximum. While these spatial relationships strongly indicate the oceanic influence on the overlying atmosphere
via the hydrostatic and vertical mixing adjustments, theobservational diagnosis alone cannot conclusively distinguish these ocean-to-atmosphere effects from atmosphere-to-ocean effects and midlatitude weather noise.
We use the WRF to isolate SST frontal effects by comparisons of the CTRL and SMTH experiments. The
experimental design was stated in section 2c.

FIG. 9. Winter climatology of SLP (contours; CI 5 1 hPa) and Q(colors; W m22) based on JRA-25. Shading convention and scale ofarrows are given to the right of each panel.

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FIG. 10. Seasonal mean SLP (contours; CI 5 1 hPa),
Q (colors; W m22) during the 2002/03 winter in the (a)control and (b) smoothed-SST experiments using
WRF. (c) As in (a), but for JRA-25. Black dotted line
indicates the SLP ridge.

the CTRL experiment and JRA-25 but weaker than thatin ICOADS. This result indicates that the MABL response to the intense surface heat flux in the K-KE regionis still imperfect in the models (Tanimoto et al. 2009).
In the CTRL experiment, ageostrophic surface westerlies are significant only along the K-KE where surface
heat flux is intense, as observed (Fig. 11a). We note thatthe simulated winds at the 850-hPa level are mostly
westerly as in the JRA-25 (not shown here). Because of
the KE meander, the meridional maximum of Q in 1478–
1488E is located at 348N, slightly more southward than in
the surrounding region. Ageostrophic surface westerlies
tend to follow this meander. Following Takatama et al.
(2011, manuscript submitted to Geophys Res. Lett.), we
defined the vertical momentum transfer from aloft as
residuals by removing the surface stress from the vertically integrated vertical mixing term in the lowest 100 m.
The difference in the vertical momentum transfer between the CTRL and SMTH experiments displays westerly acceleration over the region of intense surface heat
flux along the K-KE region, indicative of enhanced vertical mixing (Fig. 11c). As observed, the ageostrophic
westerlies are strong only north of the low-SLP wedge.
The meridional plots averaged over 1428–1478E showthat the ageostrophic component is meridionally asymmetric against the meridionally symmetric change in Q
with a peak at 358–368N (Fig. 11b). This result from the
WRF experiments supports our hypothesis of four-way

balance (Fig. 5), in which the westerly acceleration byvertical mixing and surface friction cooperates in forming
the ageostrophic component north of the low-SLP wedge
but counteracts against each other to its south.1
Figure 11d displays the differences in winter mean
ageostrophic winds between the CTRL and SMTH experiments, superimposed on SLP and Q in the SMTH
experiment. The low-SLP wedge is less pronounced andageostrophic westerlies are much weaker off the Pacificcoast of Japan in the SMTH experiment than in the CTRL
experiment because the narrow zonal band of strong surface heat flux disappears from the K-KE region. In themeridional plots averaged over 1428–1478E (Fig. 11e), no
significant peak is found either in the ageostrophic components or in Q. This drastic change in the local maximumof Q is due to the reduced DTOA and moisture differenceat the surface in the same region where the SST frontal
structure is much reduced in the boundary condition for
the SMTH experiment. Instead, a band of weak Q maximum is found in the SMTH experiment in a region very far
off the Pacific coast of Japan from 288N, 1308E through
308N, 1428E to 368N, 1478E. Over and/or south of the band,

1
In the surface momentum budget in the WRF experiments, thepressure gradient, Coriolis force, friction, and vertical mixing terms
are dominant in the K-KE region. Other terms such as horizontal
and vertical advection are small.

15 AUGUST 2011

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FIG. 11. Winter mean of WRF experiments: (a) ageostrophic wind (black arrows; m s21), superimposed on SLP (gray contours CI 5 0.5 hPa) and Q (gray shades; W m22) in the control experiment.
Ageostrophic winds less than 3.0 m s21 are not represented. (b) Meridional plots of ageostrophic zonal
wind (solid; m s21) and Q (dashed; W m22) averaged over 1428–1478E in the control experiment. Notethat the axis for Q is at the top of the panel, but the axis for the ageostrophic zonal winds is at the bottomof the panel. (c) Vertical momentum transfers from aloft (black arrows; m s22) between the control
and smoothed experiments, superimposed on SLP (gray contours; CI 5 0.5 hPa) and Q (gray shades;
W m22) in the control experiment. Vertical momentum transfers less than 0.05 m s22 are not represented. (d) Differences in ageostrophic winds (black arrows; m s21) between the control and smoothedexperiments, superimposed SLP (gray contours; CI 5 0.5 hPa), and Q (gray shades; W m22) in thesmoothed experiment. Wind differences less than 1.0 m s21 are not represented. (e) As in (b), but forthe smoothed experiment. Shading convention and scale of arrows are given on the right.

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FIG. 12. (a, top) Meridional vertical cross section of winter meanuV; contours; CI 5 2 K) and its high-pass-filtered spatial anomalies(colors; K) averaged over 1428–1478E in the control experiment. (a,bottom): Meridional plots of winter-mean Q (black; W m22) andhigh-pass-filtered spatial SLP anomalies (red; hPa). Note that theaxis for Q is at the left of the panel, but the axis for the SLP
anomalies is at the right of the panel. (b) As in (a), but for thesmoothed experiment. (c) Vertical profiles of winter mean uV averaged over 34.58–358N (red) and 378–37.58N (blue) in the 1428–
1478E section. The solid (dashed) lines denote the profile in thecontrol (smoothed) experiment.

a weak westward-extending low-SLP wedge and anomalous easterlies in ageostrophic surface winds (enhanced
ageostrophic surface westerlies in the SMTH experiment)
are found in Fig. 11d. These results show that the SST

VOLUME 24

frontal structure is necessary for the formation of the
westward-extending low-SLP wedge and ageostrophic
westerlies along the K-KE region.
In the meridional vertical section of the 1428–1478E
band, a local MABL response to the strong heat flux
from the ocean is found across the front near 358N in the
CTRL experiment (Fig. 12a). High-pass-filtered spatial
anomalies of virtual potential temperature uV in the CTRL
experiment are positive (negative) up to the 850-hPa
level from the surface in 338–358N (368–388N). The positive
anomalies are located slightly south of the peak in Q,
rather than over it. This displacement results from theadvection by the strong northerly winds in climatology,
as in the tropical simulations of Small et al. (2003, 2005).
No significant high-pass-filtered uV anomaly is presented
in the SMTH experiment (Fig. 12b) due to the lack of
the significant peak in Q in the meridional direction. The
vertical profiles of uV in Fig. 12c shows that the verticalgradient in the MABL south of 358N is much weaker in
the CTRL experiment than in the SMTH experiment,
but the gradient difference between the two experiments is not so large north of 378N. This cross-frontal
change in the MABL response indicates that the significant peak in Q is key to developing the warm, wellmixed, and thick MABL that eventually forms the SLP
minimum on the same location.
The only difference in the model setting between the
CTRL and SMTH experiments is the surface boundary
condition with and without the SST frontal structure,
respectively. The results from the experiments show that
the SST frontal structure affects the overlying atmosphere by the strong heat flux south of the front at 34.58N(Fig. 12a). The atmosphere adjusts hydrostatically to thesurface heat flux and lowers the SLP by 0.2 hPa (Fig.
12a) in the same region. Under the hydrostatic balance,
this value of the SLP decrease is mostly explained by thehigh-pass-filtered positive temperature anomalies fromthe surface up to 850 hPa. In addition, enhanced verticalmixing in the warmed MABL along with the surface
friction induces significant ageostrophic westerlies northof the low-SLP wedge (Fig. 11a).

b. Possible effect of the SST front on storms
Previous AGCM experiments reveal that strong sensible and latent heat fluxes along the extratropical SST
front maintain the baroclinicity (Brayshaw et al. 2008;
Nakamura et al. 2008; Nonaka et al. 2009). In addition,
RAM experiments show that the associated moisturesupplies can help energize midlatitude synoptic disturbances and thereby increase precipitation along the SST
front (Taguchi et al. 2009; Iizuka 2010). To examine theseeffects of the SST front onto the midlatitude storm, weanalyzed the precipitation field in the experiment. In the15 AUGUST 2011

TANIMOTO ET AL.

4431

the rainfall intensity is different between them. A coherent narrow (18–28 latitude length) band of increased
rainfall over 80 mm month21 is found north of the lowSLP wedge in the CTRL experiment (Fig. 13b). In addition, the intraseasonal variance is larger in the CTRL
than in the SMTH experiment (Fig. 13c), indicatingsome enhanced storms passing by along the rainband.
The results indicate that intense surface heat flux along
the SST front could influence the storm activity passing
eastward over the K-KE. In the climatology, however,intensity of the storm activity is generally strong at theentrance of the Pacific jet stream in the midlatitudes (308–408N). Therefore, it is still unclear whether the collocationof the strong Q across the SST front and active storm-track
activity is by chance or the consequence of the atmospheric
response to the SST front. The present study indicatesthat the SST front intensifies the overlying rainband eventhough the rainband itself may be originally formed by anyother reason. Adachi and Kimura (2007) pointed out thatthe cyclogenesis is intense along the Pacific coast of Japan.
Lower SLP and weaker stability of the MABL along the
SST front, compared to its surrounding areas, can be favorable for the development of cyclones that involve theupper trough. If it is the case, we will investigate, in the
future, how the growth of storms along the SST front
contributes to the climatological low-SLP wedge.

5. Summary and discussion

FIG. 13. Winter mean SLP (contours; CI 5 1 hPa) and (a) accumulated rainfall (colors; mm month21) in the control experiment.
(b) Differences in the simulated winter mean rainfall between thecontrol and smoothed experiments (colors; mm month21), and (c)differences in their standard deviations (mm day21).

winter mean of the CTRL experiment (Fig. 13a), a zonal
band of precipitation, with a meridional length of about
58 latitude, extends east–northeastward from the Pacificcoast of Japan into 408N, 1708E. While the similar rainband is simulated in the SMTH experiment (not shown),

We have investigated the spatial characteristics of theclimatological SLP and surface wind fields in the western
North Pacific where the monsoonal flow is pronounced. In
our analyses, in situ marine observations of ICOADS are
employed along with the JRA-25 reanalysis and the RAM
simulations. Previous in situ atmospheric soundings acrossthe SST front along the Kuroshio and its extension (K-KE)
revealed transitions of the MABL structure including its
height, temperature, moisture, wind, and cloud regime(Tokinaga et al. 2006; Tanimoto et al. 2009; Tokinaga
et al. 2009), suggesting both hydrostatic and verticalmixing effects in the MABL. Therefore, our focus is how
these frontal effects modify the monsoons.
In the winter climatology, a westward-extending lowSLP wedge is evident in SLP contours, displaced slightly
south of the zonal band of strong surface heat flux Q
along the K-KE (Fig. 3b). The steep SST front maintains
a distinct peak of temperature difference between theocean and atmosphere DTOA . The low-SLP wedge in theunfiltered field strongly indicates the hydrostatic effectalong the SST front as found in high-pass-filtered SLP fieldby Tokinaga et al. (2009) and Xu et al. (2010) along the
K-KE region and Minobe et al. (2008, 2010) along the Gulf
Stream. The southward shift of the low-SLP wedge from

4432

JOURNAL OF CLIMATEthe Q peak is probably due to thermal advection by themean northwesterlies, as shown in the equatorial region by
Small et al. (2003, 2005). While weak easterlies are expected from geostrophy north of the low-SLP wedge, theobserved surface winds are northwesterlies with a crossbaric component toward the center axis of the low-SLP
wedge. This result indicates an eastward acceleration due toenhanced vertical mixing of the midlatitude westerlies fromaloft along the band of strong Q, leading to a northwesterly
ageostrophic component in the presence of surface friction(Fig. 5c). In contrast, the observed winds are mostly geostrophic south of the low-SLP wedge where vertical mixing and friction work against each other (Fig. 5d). Albeit
slightly weaker in magnitude, the autumn climatology
displays similar spatial features. These observational
findings indicate that the intense surface heat flux heats
the MABL over the K-KE region, lowering the local
SLP by a hydrostatic adjustment.
To examine the effect of the SST front on the overlying MABL, we conduct regional atmospheric model(RAM) experiments using nonhydrostatic and highresolution WRF model for winter 2002/03. In the control(CTRL) experiment with the prescribed high-resolution
SST on a 0.258 grid, the westward-extending low-SLP
wedge and ageostrophic westerlies north of it are wellsimulated along the K-KE. The intense surface heat flux
is simulated over the SST front, albeit its amplitude is
1.5–2 times larger than the observed. In contrast, in thesmoothed (SMTH) experiment that weakens the SST
front, the band of intense heat flux along the K-KE andlow-SLP wedge are not represented. These results support the ideas that the SST frontal structure along the
K-KE is important for the strong heat flux, and that theheat flux heats the MABL, lowers the SLP locally, andproduces the ageostrophic westerlies through enhanced
vertical mixing.
Despite assimilating in situ observations, the low-SLP
wedge in the JRA-25 climatology is weak compared to
the ICOADS climatology. The MABL response acrossthe SST front in AGCMs used for the assimilation should
be carefully examined. An important implication fromthe low-SLP wedge in the reanalysis data is that not allobservational datasets contain this low-SLP wedge pattern except in the climatology from ICOADS and hence
our understanding on their interannual characteristics isstill highly limited. Longer-term numerical model simulations are necessary to assess the local and larger-scale
impact of this regional SLP pattern arising from the fronts
on the atmosphere.
In spring climatology, the surface heat flux is still
strong along the K-KE. Because the temperature in theocean and the atmosphere is not still completely in equilibrium, the resultant low-SLP wedge extending eastward

VOLUME 24

and ageostrophic winds are found even under the southwesterly monsoonal flow. In the seasonal cycle along theKE, the meridional maximum of Q is anchored at 35.58N
where the KE front maintains large DTOA year round.
The SLP minimum (center of the low-SLP wedge) showsa meridional migration, displaced south of the front
during the northwesterly monsoon and collocated with
the KE front during the southwesterly monsoon. This spatial relation of SLP and Q is due to thermal advection by themean winds. Thus, the low-SLP wedge and ageostrophic
winds are observed not only off the Pacific coast but also
along the KE region that is far (800–1000 km) from theJapanese islands from autumn to spring regardless of thechange in monsoonal wind. The RAM successfully reproduces those atmospheric adjustments in the CTRL
experiment but fails in the SMTH experiment. If theorographic effect were a major factor for the pressure
minimum, the SMTH experiment would simulate thelow-SLP wedge along the Pacific coast of Japan. As in
section 4, it is not the case in our experiments (Fig. 10b).
Ageostrophic winds in the CTRL experiment are simulated all the way from 1308 to 1508E along the K-KEand follow the local maximum of Q along the KE meander (Fig. 11a). These results indicate that the atmospheric adjustments are caused by intense heat flux
along the K-KE, rather than by the orographic effect ofhigh mountains on Japanese islands.
Acknowledgments. The authors wish to thank Drs.
H. Nakamura, B. Taguchi, M. Nonaka, M. Kimoto,H. Mukougawa, M. Watanabe, S. Adachi, and K.
Takatama for useful discussions, and M. Xu and H. Xu
for providing the WRF codes for momentum budget
analysis. We are also grateful to three anonymous reviewers for their constructive comments, which helped
improve the manuscript. This work was supported inpart by the Global Environment Research Fund (S-5) of
the Ministry of the Environment, Japan, Grand-In-Aid
for Scientific Research defrayed by the Ministry of Education, Culture, Sports, Science and Technology ofJapan (18204044, 22340132, 22340135, 22106007), NOAA,
NASA, and JAMSTEC.
REFERENCES
Adachi, S., and F. Kimura, 2007: A 36-year climatology of surface
cyclogenesis in East Asia using high-resolution reanalysis
data. SOLA, 3, 113–116.
Betts, A. K., and M. J. Miller, 1986: A new convective adjustment
scheme. Part II: Single column tests using GATE wave, BOMEX,ATEX and arctic air-mass data sets. Quart. J. Roy. Meteor. Soc.,
112, 693–709.
Brayshaw, D. J., B. Hoskins, and M. Blackburn, 2008: The stormtrack response to idealized SST perturbations in an aquaplanet GCM. J. Atmos. Sci., 65, 2842–2860.

15 AUGUST 2011

TANIMOTO ET AL.

Chelton, D. B., M. G. Schlax, M. H. Freilich, and R. F. Milliff, 2004:
Satellite measurements reveal persistent small-scale featuresin ocean winds. Science, 303, 978–983.
Cronin, M. F., S.-P. Xie, and H. Hashizume, 2003: Barometricpressure variations associated with eastern Pacific tropicalinstability waves. J. Climate, 16, 3050–3057.
Deser, C., M. A. Alexander, and M. S. Timlin, 1996: Upper-ocean
thermal variations in the North Pacific during 1970–1991.
J. Climate, 9, 1840–1855.
Dudhia, J., 1989: Numerical study of convection observed during the winter monsoon experiment using a mesoscale twodimensional model. J. Atmos. Sci., 46, 3077–3107.
Hashizume, H., S.-P. Xie, M. Fujiwara, M. Shiotani, T. Watanabe,
Y. Tanimoto, W. T. Liu, and K. Takeuchi, 2002: Direct observations of atmospheric boundary layer response to SST
variations associated with tropical instability waves over theeastern equatorial Pacific. J. Climate, 15, 3379–3393.
Hayes, S. P., M. J. McPhaden, and J. M. Wallace, 1989: The influence of sea surface temperature on surface wind in theeastern equatorial Pacific: Weekly to monthly variability.
J. Climate, 2, 1500–1506.
Hong, S.-Y., and J.-O. J. Lim, 2006: The WRF single-moment 6class microphysics scheme (WSM6). J. Korean Meteor. Soc.,
42, 129–151.
Iizuka, S., 2010: Simulations of wintertime precipitation in the vicinity of Japan: Sensitivity to fine-scale distributions of SST.
J. Geophys. Res., 117, D10107, doi:10.1029/2009JD012576.
Janjić, Z. I., 1994: The step-mountain eta coordinate model: Further developments of the convection, viscous sublayer, andturbulence closure schemes. Mon. Wea. Rev., 122, 927–945.
Kleeman, R., and S. B. Power, 1995: A simple atmospheric modelof surface heat flux for use in ocean modeling studies. J. Phys.
Oceanogr., 25, 92–105.
Kubota, M., N. Iwasaka, S. Kizu, M. Konda, and K. Kutsuwada,
2002: Japanese ocean flux data sets with use of remote sensing
observations (J-OFURO). J. Oceanogr., 58, 213–225.
Lindzen, R. S., and S. Nigam, 1987: On the role of sea surface
temperature gradients in forcing low-level winds and convergence in the Tropics. J. Atmos. Sci., 44, 2418–2436.
Minobe, S., A. Kuwano-Yoshida, N. Komori, S.-P. Xie, and R. J.
Small, 2008: Influence of the Gulf Stream on the troposphere.
Nature, 452, 206–209.
——, M. Miyashita, A. Kuwano-Yoshida, H. Tokinaga, and S.-P.
Xie, 2010: Atmospheric response to the Gulf Stream: Seasonalvariations. J. Climate, 23, 3699–3719.
Mlawer, E. J., S. J. Taubman, P. D. Brown, M. J. Iacono, and S. A.
Clough, 1997: Radiative transfer for inhomogeneous atmospheres: RRTM, a validated correlated-k model for the
longwave. J. Geophys. Res., 102, 16 663–16 682.
Nakamura, H., T. Sampe, A. Goto, W. Ohfuchi, and S.-P. Xie,
2008: On the importance of midlatitude oceanic frontal zones
for the mean state and dominant variability in the troposphericcirculation. Geophys. Res. Lett., 35, L15709, doi:10.1029/
2008gl034010.
Nonaka, M., H. Nakamura, B. Taguchi, N. Komori, A. KuwanoYoshida, and K. Takaya, 2009: Air–sea heat exchanges characteristic of a prominent midlatitude oceanic front in thesouth Indian Ocean as simulated in a high-resolution coupled
GCM. J. Climate, 22, 6515–6535.
O’Neill, L. W., D. B. Chelton, and S. K. Esbensen, 2003: Observations of SST-induced perturbations of the wind stress fieldover the Southern Ocean on seasonal timescales. J. Climate,
16, 2340–2354.

4433——, ——, ——, and F. J. Wentz, 2005: High-resolution satellite
measurements of the atmospheric boundary layer response to
SST variations along the Agulhas Return Current. J. Climate,18, 2706–2723.
——, ——, and ——, 2010: The effects of SST-induced surface
wind speed and direction gradients on midlatitude surface
vorticity and divergence. J. Climate, 23, 255–281.
Onogi, K., and Coauthors, 2007: The JRA-25 Reanalysis. J. Meteor.
Soc. Japan, 85, 369–432.
Park, K.-A., P. Cornillon, and D. L. Codiga, 2006: Modification of
surface winds near ocean fronts: Effects of Gulf Stream rings
on scatterometer (QuikSCAT, NSCAT) wind observations.
J. Geophys. Res., 111, C03021, doi:10.1029/2005jc003016.
Pleim, J. E., and J. S. Chang, 1992: A non-local closure model for
vertical mixing in the convective boundary layer. Atmos. Environ., 26A, 965–981.
Raymond, D. J., and Coauthors, 2004: EPIC2001 and the coupled
ocean–atmosphere system of the tropical east Pacific. Bull.
Amer. Meteor. Soc., 85, 1341–1354.
Reynolds, R. W., T. M. Smith, C. Liu, D. B. Chelton, K. S. Casey,
and M. G. Schlax, 2007: Daily high-resolution-blended analyses for sea surface temperature. J. Climate, 20, 5473–5496.
Schneider, N., A. J. Miller, M. A. Alexander, and C. Deser, 1999:
Subduction of decadal North Pacific temperature anomalies:
Observations and dynamics. J. Phys. Oceanogr., 29, 1056–
1070.
Skamarock, W. C., J. B. Klemp, J. Dudhia, D. M. Barker, M. G.
Duda, X.-Y. Huang, W. Wang, and J. G. Powers, 2008: A
description of the Advanced Research WRF version 3. NCAR
Tech. Note NCAR/TN-4751STR, 113 pp.
Small, R. J., S.-P. Xie, and Y. Wang, 2003: Numerical simulationof atmospheric response to Pacific tropical instability waves.
J. Climate, 16, 3723–3741.
——, ——, and J. Hafner, 2005: Satellite observations of mesoscaleocean features and copropagating atmospheric surface fields
in the tropical belt. J. Geophys. Res., 110, C02021, doi:10.1029/
2004jc002598.
——, and Coauthors, 2008: Air–sea interaction over ocean fronts
and eddies. Dyn. Atmos. Oceans, 45, 274–319.
Song, Q., P. Cornillon, and T. Hara, 2006: Surface wind response to
oceanic fronts. J. Geophys. Res., 111, C12006, doi:10.1029/
2006jc003680.
——, D. B. Chelton, S. K. Esbensen, N. Thum, and L. W. O’Neill,
2009: Coupling between sea surface temperature and lowlevel winds in mesoscale numerical models. J. Climate, 22,
146–164.
Suga, T., and K. Hanawa, 1995: The subtropical mode water
circulation in the North Pacific. J. Phys. Oceanogr., 25,
958–970.
Sugimoto, S., and K. Hanawa, 2007: Impact of remote reemergenceof the subtropical mode water on winter SST variation in thecentral North Pacific. J. Climate, 20, 173–186.
Taguchi, B., H. Nakamura, M. Nonaka, and S.-P. Xie, 2009: Influences of the Kuroshio–Oyashio Extensions on air–sea heatexchanges and storm-track activity as revealed in regional
atmospheric model simulations for the 2003/04 cold season.
J. Climate, 22, 6536–6560.
Tanimoto, Y., H. Nakamura, T. Kagimoto, and S. Yamane, 2003:
An active role of extratropical sea surface temperature
anomalies in determining anomalous turbulent heat flux.
J. Geophys. Res., 108, 3304, doi:10.1029/2002jc001750.
——, S.-P. Xie, K. Kai, H. Okajima, H. Tokinaga, T. Murayama,
M. Nonaka, and H. Nakamura, 2009: Observations of marine

4434

JOURNAL OF CLIMATE

atmospheric boundary layer transitions across the summer
Kuroshio Extension. J. Climate, 22, 1360–1374.
Tokinaga, H., Y. Tanimoto, and S.-P. Xie, 2005: SST-induced
surface wind variations over the Brazil–Malvinas confluence: Satellite and in situ observations. J. Climate, 18, 3470–
3482.
—— , and Coauthors, 2006: Atmospheric sounding over the winterKuroshio Extension: Effect of surface stability on atmosphericboundary layer structure. Geophys. Res. Lett., 33, L04703,
doi:10.1029/2005gl025102.
——, Y. Tanimoto, S.-P. Xie, T. Sampe, H. Tomita, and H. Ichikawa,
2009: Ocean frontal effects on the certical development of clouds
over the western North Pacific: In situ and satellite observations.
J. Climate, 22, 4241–4260.
Tomita, H., M. Kubota, M. F. Cronin, S. Iwasaki, M. Konda, andH. Ichikawa, 2010: An assessment of surface heat fluxes from
J-OFURO2 at the KEO and JKEO sites. J. Geophys. Res.,
115, C03018, doi:10.1029/2009jc005545.
Tourre, Y. M., Y. Kushnir, and W. B. White, 1999: Evolutionof interdecadal variability in sea level pressure, sea surface

VOLUME 24

temperature, and upper ocean temperature over the Pacific
Ocean. J. Phys. Oceanogr., 29, 1528–1541.
Vecchi, G. A., S.-P. Xie, and A. S. Fischer, 2004: Ocean–atmospherecovariability in the western Arabian Sea. J. Climate, 17, 1213–
1224.
Wallace, J. M., T. P. Mitchell, and C. Deser, 1989: The influence ofsea surface temperature on surface wind in the eastern equatorial Pacific: Seasonal and interannual variability. J. Climate,2, 1492–1499.
Wentz, F. J., and T. Meissner, 2000: Algorithm Theoretical Basis
Document (ATBD) version 2: AMSR ocean algorithm. Remote Sensing Systems Tech. Rep. 121599A-1, 59 pp.
Worley, S. J., S. D. Woodruff, R. W. Reynolds, S. J. Lubker, and
N. Lott, 2005: ICOADS release 2.1 data and products. Int.
J. Climatol., 25, 823–842.
Xie, S.-P., 2004: Satellite observations of cool ocean–atmosphereinteraction. Bull. Amer. Meteor. Soc., 85, 195–208.
Xu, H., H. Tokinaga, and S.-P. Xie, 2010: Atmospheric effects of
the Kuroshio large meander during 2004–05. J. Climate, 23,
4704–4715.