1. Introduction
Recent progress of satellite observations has revealed a significant influence of mesoscale sea surface temperature (SST) features on a marine atmospheric boundary layer (Chelton et al. 2004; Xie 2004; Sampe and Xie 2007; Small et al. 2008; Chelton and Xie 2010; Kelly et al. 2010). These effects are especially prominent over SST fronts associated with western boundary currents, such as, the Gulf Stream (Chelton et al. 2004; Park et al. 2006; Minobe et al. 2008, 2010), the Kuroshio (Xie et al. 2002; Nonaka and Xie 2003), the Kuroshio Extension (Tokinaga et al. 2009; Tanimoto et al. 2011), the Agulhas Return Current (O’Neill et al. 2003, 2005; Liu et al. 2007), and the Brazil and Malvinas Currents (Tokinaga et al. 2005). Furthermore, the influences of SSTs also extend to the troposphere as well as the marine atmospheric boundary layer (e.g., Liu et al. 2007; Minobe et al. 2008; Kobashi et al. 2008). An interesting feature of the atmospheric responses is found in precipitation. Minobe et al. (2008, 2010) reported from satellite observations that the precipitation is enhanced just over the Gulf Stream axis. Similar high rainfall over western boundary currents has been observed over the Kuroshio in the East China Sea in winter and spring (Xie et al. 2002; Small et al. 2008; Xu et al. 2011) and over the Kuroshio Extension in winter (Tokinaga et al. 2009). These atmospheric responses to SST fronts are not the same and depend on season. Tokinaga et al. (2009) found that the enhanced cloud liquid water is located over the warmer flank of the Kuroshio Extension in winter, but becomes blurred in summer. Minobe et al. (2010) suggested that two atmospheric response modes occur to the Gulf Stream. One is the shallow-heating mode prominent in winter, which accompanies strong sensible and latent heating in the lower troposphere. The other is the deep-heating mode mainly occurring in summer, which is characterized by latent heating in the mid- and upper troposphere due to deep convection.
The East China Sea, a marginal sea east of China and west of the Ryukyu Islands, is one of the regions where an air–sea interaction is active (e.g., Small et al. 2008). In this sea, the Kuroshio transports large amount of heat poleward, resulting in warm SSTs over the current axis with a pronounced contrast to colder SST of colder shelf waters. Xie et al. (2002) and Small et al. (2008) showed from satellite observations that the Kuroshio in the East China Sea accompanies the surface wind convergence, high turbulent heat flux, and enhanced cloud liquid water content over the current in cold season, though evidences of deep penetration of atmospheric responses above the marine atmospheric boundary layer were not reported. On the other hand, Xu et al. (2011) demonstrated deep atmospheric responses in spring (March–May), characterized by enhanced convective precipitation, frequent occurrence of cumulus convection, and surface wind convergence based on observational data analysis and regional atmospheric model experiments. They concluded that the spring atmospheric responses are mainly due to the deep-heating mode proposed by Minobe et al. (2010). However, the previous studies for ocean–atmosphere influence over the East China Sea did not investigate all seasons (Xie et al. 2002; Small et al. 2008; Xu et al. 2011), and thus important ocean–atmosphere feedback may not be limited to winter and spring seasons. If one examines precipitation over the Kuroshio in the East China Sea year-round, the rain rate exhibits the outstanding maximum in June, which is more than double of rain rate in any months in the winter and spring seasons (Fig. 1). This high rainfall indicates that ocean–atmosphere influence over the Kuroshio in June is an important research subject, because precipitation is a good measure of atmospheric responses to western boundary currents (Minobe et al. 2008, 2010; Small et al. 2008; Tokinaga et al. 2009; Xu et al. 2011).
Monthly mean precipitation rate in the TRMM 3B43 product averaged over the Kuroshio in the East China Sea (denoted by the gray rectangle box shown in Fig. 4a).
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
The high rainfall over the Kuroshio in June is related to the major rain season, called the baiu in Japan and the mei-yu in China (referred to as the baiu in this study). The baiu rainband is characterized by strong gradients of equivalent potential temperature and specific humidity (Ninomiya 1984), and has been examined by a number of studies (e.g., Akiyama 1973; Ninomiya 1984; Ninomiya and Akiyama 1992; Moteki et al. 2004; Wang 2006; Ninomiya and Shibagaki 2007; Sampe and Xie 2010). The northward moisture transport by low-level wind and associated convective instability play important roles in forming the baiu rainband (e.g., Ninomiya 1984; Ninomiya and Shibagaki 2007). Furthermore, the location of the baiu rainband may be determined by the warm air advection at the midtroposphere from the Tibetan Plateau (Sampe and Xie 2010; Kosaka et al. 2011). Sampe and Xie (2010) also pointed out that energetic weather disturbances over the westerly jet result in favorable conditions for strong convection of the rainband. Nevertheless, the influence of the warm SST of the Kuroshio in the East China Sea on the baiu rainband has not been investigated yet.
The purpose of this study is to describe the influence of the Kuroshio in the East China Sea on the baiu rainband and explore the role of the warm SSTs of the Kuroshio. Using observational datasets, we investigate the precipitation and wind fields over the Kuroshio mainly in June, when the baiu rainband covers the East China Sea and the monthly precipitation is the highest (Fig. 1). Also, we clarify the vertical structure and the strength of the atmospheric response to the Kuroshio. In addition, the importance of the warm SSTs of the Kuroshio for atmosphere responses in June is examined using a regional atmospheric model. The rest of the present paper is organized as follows: observational datasets and the model simulations are explained in section 2. In section 3, we describe the observational mean state of the atmosphere over the Kuroshio in the East China Sea, and explore in section 4 the simulation results. A summary and discussion are presented in section 5.
2. Data
a. Observational datasets
To obtain a fine structure of precipitation distribution, a Tropical Rainfall Measuring Mission (TRMM) 3B43 product derived from the TRMM satellite and other observations is used. This dataset is available on a 0.25° × 0.25° grid extending from 50°S to 50°N at monthly intervals. For near-surface wind estimates, we employ the data measured by the Quick Scatterometer (QuikSCAT) on the QuikBird satellite. This product is gridded to on a 0.25° × 0.25° spatial resolution at monthly intervals by Remote Sensing Systems (RSS). We analyze a scalar mean of wind speed and a vector mean of zonal and meridional wind at 10-m height. A monthly surface evaporation dataset at a 1° × 1° spatial resolution is from the Japanese Ocean Flux dataset version 2.1 (J-OFURO; Kubota and Tomita 2007). The dataset is produced by remote sensing observations and the algorithm of the Coupled Ocean–Atmosphere Response Experiment, version 3.0 (COARE 3.0; Fairall et al. 2003). To investigate the vertical structure of the atmospheric response to the Kuroshio, we use 6-h forecast fields of the National Centers for Environmental Prediction (NCEP) Climate Forecast System Reanalysis (CFSR) product with a spatial resolution of 0.5° × 0.5° with 37 vertical levels (Saha et al. 2010). The original atmospheric model of this product has a resolution of T382 (about 38 km) with 64 vertical levels, and is coupled to a global ocean general circulation model [the Modular Ocean Model, version 4 (MOM)] and a sea ice model. This ocean model assimilates the version 2 of the optimum interpolation SST product from November 1981 (OISSTv2; Reynolds et al. 2007) provided by the National Oceanic and Atmospheric Administration (NOAA). The OISSTv2 dataset assimilated in the NCEP-CFSR product uses in situ and the Advanced Very High Resolution Radiometer (AVHRR) SST data, but since June 2002 also uses the Advanced Microwave Scanning Radiometer (AMSR) data. Thus, this SST product since 2002 improves the spatial resolution of SST fronts and tends to be more accurate compared to that by the AVHRR-only product (Reynolds et al. 2007).
In addition to the atmospheric datasets, we employ two datasets of ocean variables. One is SST data on a 0.25° × 0.25° grid of the OISSTv2 (Reynolds et al. 2007). The other is surface geostrophic velocity data on a ⅓° × ⅓° Mercator grid provided by Archiving, Validation and Interpretation of Satellites Oceanographic data (AVISO). The surface geostrophic velocity fields are given by the sum of sea level anomalies estimated from satellite altimetry combined observations from the Ocean Topography Experiment (TOPEX)/Poseidon, the European Remote Sensing Satellites-1/2 (ERS-1/2), Jason-1, and the Environmental Satellite (Envisat; Ducet et al. 2000) and the mean dynamic topography by Rio and Hernandez (2004). The analysis period in the present study is from January 2003 to December 2008 except for the J-OFURO dataset, which is provided from January 2003 to December 2006. The beginning of our study period is based on the year when the NCEP-CFSR product assimilates the OISSTv2 dataset that contains the AMSR data from the beginning of the year.
b. A regional atmospheric model
The regional atmospheric model used in this study is the Japan Meteorological Agency/Meteorological Research Institute (JMA/MRI) nonhydrostatic model (NHM), which employs the nonhydrostatic formulation in a terrain-following coordinate (Saito et al. 2006, 2007). The model employs the Lambert conformal projection with the center at 31°N, 120°E, and has a 20-km horizontal resolution and 180 meridional × 200 zonal grid points. The model domain roughly covers the area from 15° to 45°N and from 110° to 140°E. The vertical levels of the NHM are 40 with realistic bottom topography. For convective precipitation, the Kain–Fritsch convective parameterization scheme (Kain and Fritsch 1993; Kain 2004) is used.
Six 34-day (from 28 May to 30 June) simulations are conducted for the period 2003 to 2008 (i.e., one run for each year), where initial conditions for each simulation are taken from the Japanese 25-yr Reanalysis/Japan Meteorological Agency Climate Data Assimilation System (JRA-25/JCDAS) data with a 1.25° × 1.25° spatial resolution (Onogi et al. 2007). The lateral boundary condition is also given by 6-hourly data of the JRA-25/JCDAS reanalysis product. The daily OISSTv2 product is employed for the surface boundary condition over the sea. Hereafter, we refer to this run as the CTL run. In addition to the CTL run, to clarify the influence of the warm SSTs of the Kuroshio on precipitation, we also conduct additional experiments forced by spatially smoothing SSTs, where SSTs are smoothed by applying a spatial filter with half-power cutoffs of 5° in latitude and longitude. This run is referred to as the SMTH run.
3. Observed mean state
First, we briefly summarize oceanic conditions in the East China Sea. The Kuroshio enters the East China Sea from the passage east of Taiwan, and leaves through the Tokara Strait into the western North Pacific (Fig. 2a). In the East China Sea, the Kuroshio flows northeastward along the continental slope, and the path of the Kuroshio is nearly steady (e.g., Nakamura et al. 2003). The Kuroshio accompanies a warm SST tongue structure of 26°–28°C on the southeastern side of the Kuroshio axis, which is 1°–2°C warmer than SST at the same latitude in the western East China Sea (Fig. 2b). Because of this, the warm SST tongue is flanked on the northwest by strong SST gradients (color in Fig. 2b).
(a) Absolute velocity of surface geostrophic current (cm s−1) with thick contours of SST at 26° and 27°C in June. (b) SST (contour; °C) and its gradients (color; ×10−6 °C m−1) in June. The contour interval is 1°C with thick contours at 26° and 27°C.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
Before describing precipitation over the East China Sea, we overview large-scale precipitation distributions over East Asia and the western North Pacific and its seasonal variability (Fig. 3). In winter, the precipitation rate is high over the Kuroshio Extension region in the northwestern North Pacific (Figs. 3a,b). The high precipitation is mainly induced by extratropical winter storms. On the other hand, a high rainfall region extends from south China to Japan in the warm season (May–June). During April and May, a high rainband is located over the south of China (Figs. 3d,e). As the season progresses this high rainfall area extends eastward, migrates northward, and is located over Korea and south Japan in July (Fig. 3g). This rainband is the baiu rainband as mentioned in the introduction.
Monthly mean precipitation rate in the TRMM 3B43 product (color) and SSTs (contour) from 2003 to 2008. The contour interval is 1°C.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
We now turn our attention to the precipitation over the East China Sea. Consistent with the rainfall over the Kuroshio (Fig. 1), the rainfall over the East China Sea is the highest in June and the lowest in July (Figs. 3f,g). This rapid seasonal change is due to the northward migration of the baiu rainband. In June, the high rainfall area extends from south of China to the southwestern North Pacific as well as over the East China Sea (Fig. 3f). A closer look shows that the precipitation rate in June reaches 12–15 mm day−1 over the warm SST tongue of the Kuroshio (Fig. 4a), and is twice as large as that in the surrounding area (about 5–7 mm day−1). Table 1 summarizes the precipitation rate over the Kuroshio and its surrounding area. The differences of the precipitation rate over the Kuroshio minus the northwest of the Kuroshio and over the Kuroshio minus over the southeast of the Kuroshio are significantly positive at 95% confidence level, based on a Student’s t test assuming each year is independent. The enhanced rainfall over the Kuroshio is often seen in other seasons, especially in the cold season, but their amplitude is much weaker than in June (see Fig. 3). The enhanced precipitation over the East China Sea is well collocated with the warm SST tongue of the Kuroshio (contours in Fig. 4a), embedded in the large-scale baiu rainband (Fig. 3f). The spatial correspondence and the narrow structure embedded in the large-scale rainband indicate a significant effect of the Kuroshio on the locally enhanced precipitation.
Precipitation rate in June in (a) the TRMM 3B43 product, (b) the NCEP-CFSR product, and (c) the convective precipitation rate of the NCEP-CFSR product (color; mm day−1). The contours are for SST, and the interval is 1°C with thick contours at 26° and 27°C. A gray rectangle box in (a) denotes the analysis region in Fig. 1. The regions of this gray rectangle box and black rectangle boxes are used for moisture budget analysis given in Table 1.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
Precipitation (P) of the TRMM 3B43 product, evaporation (E) of the J-OFURO data, and vertical integrated water vapor flux divergence (WVF div) in the CTL run averaged over the Kuroshio, and the northwest (NW) and southeast (SE) of the Kuroshio. These regions are denoted by rectangle boxes shown in Fig. 4a. Unit is mm day−1.
The NCEP-CFSR product captures the enhanced precipitation over the Kuroshio in June found in the TRMM 3B43 observations (Fig. 4b). The spatial pattern of precipitation rate over the East China Sea in the NCEP-CFSR data resembles that of the TRMM 3B43 product (Fig. 4a), while the high rainband over the Kuroshio slightly shifts eastward compared to the TRMM 3B43 dataset. The peak of the precipitation rate over the Kuroshio of the NCEP-CFSR product is also comparable that of the TRMM 3B43 dataset. Hence, the precipitation field of the NCEP-CFSR dataset reproduces well that by the TRMM 3B43 observations. The precipitation in the NCEP-CFSR product consists of convective precipitation and large-scale precipitation. Since the importance of convective instability for the baiu rainband is well established (e.g., Ninomiya 1984; Ninomiya and Shibagaki 2007), it is informative to examine the relative contribution of the convective precipitation to the enhanced precipitation over the Kuroshio. Figure 4c indicates that half of the enhanced precipitation over the Kuroshio is induced by convective precipitation. Since the warm SST tongue is warmer than 26°C (Fig. 2b), this region is suitable for the deep convection (Graham and Barnett 1987; Waliser et al. 1993). The maximum of the large-scale precipitation of the NCEP-CFSR product is not collocated with the warm SST tongue of the Kuroshio (not shown).
Surface evaporation also exhibits an enhancement over the Kuroshio (Fig. 5c). The high evaporation region in the East China Sea also spatially corresponds well to the warm SST tongue. This high evaporation rate over the Kuroshio is larger in magnitude than the evaporation rate at its surrounding area in the East China Sea and the western North Pacific (Table 1). This strong surface evaporation can be caused by the warm SSTs over the Kuroshio via to two mechanisms: larger saturation vapor pressures for higher temperatures and, as will be shown in the next paragraph, stronger wind speeds over warmer SSTs. From the viewpoint of the moisture budget, the local evaporation over the Kuroshio is several times smaller in amplitude than the precipitation rate there (Fig. 4a and Table 1). This implies that a main part of the moisture for the locally enhanced precipitation over the Kuroshio is supplied by horizontal advection, similar to its role in moisture supply for the large-scale baiu rainband (e.g., Ninomiya 1984; Ninomiya and Shibagaki 2007). However, the contrast of the surface evaporation between over the warm SST tongue of the Kuroshio and its surrounding area is about 3 mm day−1 and corresponds to about half of the aforementioned contrast of the precipitation. This result suggests that the moisture supply by the local surface evaporation plays an important role in enhancing the precipitation over the Kuroshio. This point will be further explored using a numerical model in the next section.
(a) Scalar wind speed (color; m s−1) and wind vector (vector; m s−1), (b) wind convergence (color; ×10−6 s−1) at 10 m in the QuikSCAT data, and (c) evaporation (mm day−1) in the J-OFURO data in June. The thick contours are for SST at 26° and 27°C, as in Fig. 2a.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
We next examine a surface wind response to the warm SST of the Kuroshio in June. The surface wind is an important variable for understanding the influence of warm SSTs on precipitation, because connections between high rainfall and surface wind convergence have been reported in the East China Sea in winter and spring (Xu et al. 2011) and in the Gulf Stream region (Minobe et al. 2008; Kuwano-Yoshida et al. 2010). Southerly surface wind associated with the subtropical high prevails over the southwestern North Pacific and over the southern East China Sea (Fig. 5a). The scalar mean wind speed increases over the warm SST tongue of the Kuroshio (Fig. 5a), consistent with the response over the Kuroshio in winter and spring (Xie et al. 2002; Small et al. 2008; Xu et al. 2011). This correspondence between the warm SST and enhanced wind speed can be explained by the vertical mixing mechanism (e.g., Wallace et al. 1989; Xie et al. 2002; Nonaka and Xie 2003; Chelton et al. 2004) or the boundary layer depth coupling mechanism (Samelson et al. 2006).
Wind convergence observed by QuikSCAT is generally noisy, but an inspection of Fig. 5b reveals that the wind convergence tends to be located over the warm SST tongue of the Kuroshio. This noisy pattern may be due to insufficient sampling of the QuikBird satellite (twice per day). Thus, we also examine the wind convergence field of the NCEP-CFSR product, which assimilates the surface wind of the QuikSCAT data during the study period (Saha et al. 2010). A prominent surface wind convergence band in the NCEP-CFSR data occurs just over the warm SST tongue of the Kuroshio (Fig. 6). The spatial distribution of this wind convergence corresponds well to that of the enhanced precipitation over the Kuroshio (Fig. 4a).
Wind convergence (color; ×10−6 s−1) and wind vector (vector; m s−1) at 10 m in the NCEP-CFSR data in June. The thick contours are for SST at 26° and 27°C, as in Fig. 2a.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
Interestingly, this location of the surface wind convergence just over the warm SST tongue of the Kuroshio is different from that in winter and spring, when the surface wind convergence is located over the northwestern flank of the Kuroshio axis (Xie et al. 2002; Small et al. 2008; Xu et al. 2011). In the vertical mixing mechanism, a connection between warm SSTs and high surface wind speed yields wind convergence when the wind blows from warm to cool water across a SST front (e.g., Chelton et al. 2004; Chelton and Xie 2010). Thus, the vertical mixing mechanism cannot explain the wind convergence just over the warm SST tongue in June. This is partly because the vector mean winds are weak in the eastern East China Sea (see vectors in Fig. 5a), although the scalar mean winds are enhanced over the warm SST tongue. This difference between the vector and scalar mean winds is probably because of submonthly change in wind direction over the Kuroshio associated with the change of the subtropical high. Hence, another mechanism should be responsible for the formation of the surface wind convergence in June. An alternative mechanism is the pressure adjustment mechanism (Lindzen and Nigam 1987), in which SSTs modify atmospheric temperature, so that the resultant pressure anomalies induce wind convergence (divergence) over warm (cold) SSTs. To test this hypothesis, we employ the method by Minobe et al. (2008), who suggested that an importance of the mechanism is measured by a relation of surface wind convergence to the Laplacian of sea level pressures (SLPs) and SSTs, confirmed by the diagnostic study using a regional atmospheric model by Takatama et al. (2012). We use the 1000-hPa geopotential height (Z1000) instead of SLP, because the SLP Laplacian is much noisier than the Z1000 Laplacian, suggesting a problem in estimation of SLPs in the NCEP-CFSR data. The positive values of the sign-reversed Laplacian of SSTs around the Kuroshio (Fig. 7b) are confined along its axis, and exhibit a good correspondence with the surface wind convergence (Fig. 6) and the positive values of the Laplacian of Z1000 (Fig. 7a). This correspondence indicates that the pressure adjustment to the warm SST of the Kuroshio can be attributed to the surface wind convergence over the Kuroshio.
The Laplacian of (a) geopotential height at 1000 hPa in the NCEP-CFSR data, (b) sign-reversed SST of the OISSTv2 data, and sign-reversed thickness between two isobaric levels of 1000 and 700 hPa calculated using (c) temperature and (d) virtual temperature in the NCEP-CFSR data in June. The thick contours are for SST at 26° and 27°C, as in Fig. 2a. The unit is ×1010 m−1 in (a),(c),(d), and 5.0 × 109 K m−2 in (b).
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
Finally, we consider the vertical structure of the atmospheric response and its seasonal variability over the Kuroshio. The enhanced rainband over the Kuroshio in June is collocated with the large ascent motion extending to the upper troposphere (Fig. 8a). The upward velocity gradually increases from the surface to the midtroposphere, and its peak is at about 400 hPa with an amplitude of 0.1 Pa s−1. The horizontal pattern of the upward velocity at 400 hPa (Fig. 9b) is similar with the precipitation pattern (Fig. 4a). This ascent motion in June is much stronger than that in any other month (Fig. 8a). The second largest upward velocity at the troposphere over the Kuroshio is observed in September, but its value is one-fifth of the value in June. A similar upward motion from the surface to the upper troposphere over the Kuroshio in spring (March–May) is reported by Xu et al. (2011) using the JRA-25/JCDAS reanalysis product, but the spring ascent motion is much smaller than that in June (Fig. 8a).
Vertical sections of monthly mean (a) upward pressure velocity and (b) net diabatic heating over the Kuroshio. (c) Vertical section of net diabatic heating rate (black thick line), large-scale condensate heating rate (line with open circle), deep convective heating rate (line with closed circle), shallow convective heating rate (gray thick line), sensible heating rate (thin solid line), and the sum of solar and longwave radiative heating rates (line with triangle) over the Kuroshio in June. In (a)–(c), the averaged region over the Kuroshio is the gray rectangle box in Fig. 4a, and the NCEP-CFSR data are used. The contour interval is 2 × 10−2 Pa s−1 in (a) and 1 K day−1 in (b).
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
(a) Precipitation rate in July in the TRMM 3B43 product (color; mm day−1). Upward pressure velocity (color; 10−2 Pa s−1) at 400 hPa (b) in the Kuroshio region in June and (c) in the Gulf Stream region in July in the NCEP-CFSR data. The contours are for SST and the interval is 2°C with thick contours at 26° and 27°C.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
The ascent motion over the warm SST tongue in June is accompanied by large diabatic heating (Fig. 8b), where the diabatic heating is defined as the sum of vertical diffusion, deep convective, shallow convective, large-scale condensate, and solar and longwave radiative heating rates in the NCEP-CFSR product. The vertical diffusion heating rate and the sum of the deep convective, shallow convective, and large-scale condensate heating rates is referred to as sensible heating and latent heating, respectively, according to Minobe et al. (2010). Again, the large heating in June occurs in the midtroposphere over the Kuroshio, and the maximum reaches more than 6.5 K day−1 at 400 hPa (Fig. 8b). The corresponding horizontal pattern of the diabatic heating rate at 400 hPa (Fig. 10a) resembles the precipitation pattern (Fig. 4a). In the midtroposphere, the deep convective heating rate is the strongest among the six components followed by the slightly weaker large-scale condensate heating rate (Fig. 8c), consistent with comparable magnitudes between the convective precipitation and large-scale precipitation over the Kuroshio (Figs. 4b–c). These two components are much larger than the shallow convective heating and vertical diffusion rates. This diabatic heating over the Kuroshio in the midtroposphere in June is larger than that in any other season (Fig. 8b). On the other hand, the large heating rate below 800 hPa is observed in winter and early spring and is probably responsible to atmospheric responses in those seasons reported by Xie et al. (2002) and Small et al. (2008). This shallow heating is mainly induced by the sensible heating in contrast to the case in June (not shown). A similar seasonal difference of the diabatic heating between summer and winter in the Gulf Stream region is reported by Minobe et al. (2010). Note that the ascent motion and diabatic heating over the Kuroshio are quite weak in July in spite of the warm SSTs around the Kuroshio (Fig. 3g). This may be due to the relatively stable atmospheric conditions for convection over the East China Sea in July associated with the northward migration of the baiu rainband and the westerly jet (Sampe and Xie 2010).
(a),(b) Net diabatic heating rate; (c),(d) deep convective heating rate; and (e),(f) large-scale condensate heating rate at 400 hPa in the NCEP-CFSR data. (a),(c),(e) The Kuroshio region in June and (b),(d),(f) the Gulf Stream region in July. The contours are for SST and the interval is 2°C with thick contours at 26° and 27°C.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
The deep atmospheric responses to the warm SST tongue of the Kuroshio in June resemble those due to the deep-heating mode, which is the atmospheric response mode prominent over the Gulf Stream in summer proposed by Minobe et al. (2010). The aforementioned vertical structures of the ascent motion and the diabatic heating over the Kuroshio are quite similar with those documented for the deep-heating mode (Minobe et al. 2010). Thus, it is interesting to compare magnitudes of the atmospheric responses between the two regions. The rain rate of the TRMM 3B43 product over the Kuroshio in June is larger than that over the Gulf Stream in any months in the summer season. As an example, Fig. 9a shows the precipitation rate of the TRMM 3B43 product in the western North Atlantic in July, when the precipitation rate is the highest over the Gulf Stream in the summer season. Similar to the case over the Kuroshio, the enhancement of the precipitation occurs over the warm SST tongue of the Gulf Stream, consistent with the finding by Minobe et al. (2010). However, the amplitude of the maximum precipitation over the Gulf Stream in July is about two-thirds of that over the Kuroshio in June (Fig. 4a).
Also, the upward wind over the Kuroshio region is much stronger than that over the Gulf Stream (Figs. 9b,c). In addition, the amplitude of the diabatic heating over the Kuroshio in June is about twice as large as that over the Gulf Stream in July (Figs. 10a,b). It is worth noting that the deep convective heating over the Kuroshio in June is as the same magnitude as that over the Gulf Stream in July (Figs. 10c,d), while the large-scale condensate heating over the Kuroshio in June is much larger than that over the Gulf Stream in July (Figs. 10e,f). Thus, the influence of the Kuroshio in June on the atmosphere is comparable in amplitude to that of the Gulf Stream in July, if the effect is measured by amplitude of the deep-convective heating rate. Such strong ascent or diabatic heating in the midtroposphere is not found for the other midlatitude ocean (poleward of 20°N/S) in the NCEP-CFSR product (not shown). Consequently, the East China Sea in June is the region of the strongest deep atmospheric response to western boundary currents along with the Gulf Stream region in summer.
4. Simulated mean state
We have shown in the previous section from observational data that precipitation is enhanced over the Kuroshio in June. The spatial distribution of the high rainfall over the Kuroshio corresponds well to the warm SST tongue of the Kuroshio and the high surface evaporation. Furthermore, the warm SST tongue is accompanied by the surface wind convergence, which is likely to be induced by the adjustment of pressure in the marine atmospheric boundary layer to the warm SSTs. In this section, we employ the nonhydrostatic regional atmospheric model to clarify the influence of the warm SST tongue on the enhanced rainfall and examine the moisture balance over the Kuroshio. We also investigate the spatial changes of properties in the marine atmospheric boundary layer over the Kuroshio.
Figure 11a shows the precipitation rate in June in the CTL run. The observed high rainfall over the Kuroshio is well simulated by the nonhydrostatic regional atmospheric model (Figs. 4a and 11a), although the model fails to simulate the high rainfall over southeastern China and that to the south of Japan. The amplitude of the precipitation over the Kuroshio in the model is also comparable to that by the TRMM 3B43 observations. These results show that the NHM simulation successfully reproduces the high rainband over the Kuroshio in the East China Sea in June, while the results may depend on the resolution and parameterization.
Simulated precipitation rate (color; mm day−1) in June (a) in the CTL run with the observed SSTs (contours), (b) in the SMTH run with the smoothed SSTs, and (c) their differences. The contour interval is 1°C with thick contours at 26° and 27°C in (a) and (b), respectively; and 0.5°C with thick contours at 1°C in (c).
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
To clarify the role of the warm SST tongue of the Kuroshio in the enhanced precipitation over the East China Sea, we compare the precipitation fields between the CTL and SMTH runs. The SST difference between the two runs reaches about 1.5°C around the warm SST tongue of the Kuroshio in the East China Sea (contours in Fig. 11c), and the SST front around the northwest flank of the Kuroshio axis becomes quite weak in the SMTH run (Fig. 11b). When the SST is smoothed, the precipitation over the warm SST tongue becomes weak, and the enhanced precipitation disappears (Fig. 11b). The difference of the precipitation rate between the CTL and SMTH runs closely follows the SST difference of the Kuroshio (Fig. 11c). Except for this change, the large-scale pattern of the baiu rainband in the SMTH run (e.g., the extent to which the precipitation rate is larger than 7.5 mm day−1) resembles that in the CTL run. Hence, the warm SST tongue associated with the Kuroshio acts to enhance locally the rainfall.
Consistently, the surface wind response exhibits significant changes between the CTL and SMTH runs. The surface wind convergence over the Kuroshio in the CTL run (Fig. 12a) is similar and comparable in amplitude that of the NCEP-CFSR product (Fig. 6). This surface wind convergence over the Kuroshio almost disappears in the SMTH run (Fig. 12b). The large difference between the two runs is again confined over the Kuroshio (Fig. 12c). Also, the Laplacian of Z1000 over the Kuroshio in the CTL exhibits a similar structure with that of the NCEP-CFSR data, but this structure disappears in the SMTH run (not shown). Consistently, the Laplacian of SST in the SMTH run is quite weak around the Kuroshio. These results suggest that the sharpness of the SST front of the Kuroshio, as well as the warmth of SSTs, plays an important role in inducing the surface wind convergence over the Kuroshio.
As in Fig. 11, but for the surface wind convergence (color; ×10−6 s−1).
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
From a viewpoint of the moisture balance, a large part of the moisture for the enhanced precipitation over the Kuroshio is supplied by horizontal advection in the CTL run (Fig. 13 and Table 1). Consistent with the observation (Fig. 5c), the evaporation in the CTL run is enhanced around the warm SST tongue (Fig. 13a), but is a few times smaller than the precipitation there (Fig. 11a). The vertically integrated water vapor flux convergence is large over the East China Sea (Fig. 13b), which is comparable in magnitude to the precipitation there (Table 1). This water vapor convergence over the Kuroshio is likely consistent with the surface wind convergence (Fig. 12a). The water vapor is mainly transported from the south by large-scale low-level winds associated with the subtropical high over the North Pacific. This is consistent with a large-scale picture of the moisture budget for the baiu front region (e.g., Ninomiya 1984; Ninomiya and Shibagaki 2007).
(a) Evaporation (mm day−1) and (b) vertical integrated water vapor flux (vector; ×10−1 m2 s−1) and its convergence (color; mm day−1) in the CTL run in June. The thick contours are for SST at 26° and 27°C, as in Fig. 11a.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
Interestingly, however, the evaporation change between the CTL and SMTH runs contributes about half of the precipitation changes over the Kuroshio (Fig. 14). This difference of the surface evaporation around the Kuroshio well corresponds to the SST difference. The difference of the water vapor flux convergence also significantly contributes to the rainfall changes over the Kuroshio (Fig. 14b), although the spatial scale of the water vapor flux convergence is smaller than the rainfall changes (Fig. 11c). Therefore, the moisture supply not only by the water vapor flux convergence but also by the energetic evaporation is attributed to the enhancement of the precipitation amount locally.
The difference of the (a) evaporation and (b) vertical integrated water vapor flux convergence (color; mm day−1) between the CTL minus SMTH runs. The thick contours are for the SST difference at 1°C as in Fig. 11c.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
As shown in Fig. 12, the warm SST of the Kuroshio influences the surface wind convergence through the temperature and humidity changes in the marine atmospheric boundary layer. To further clarify the effect of the warm SST of the Kuroshio on the properties in the marine atmospheric boundary layer, we show the difference of the equivalent potential temperature θe, mixing ratio, and temperature at 950 hPa between the CTL and SMTH runs (Figs. 15a–c). Over the Kuroshio, the equivalent potential temperature in the CTL run is about 1°C warmer than that in the SMTH run. This signature of the warm SST tongue in the equivalent potential temperature extends to 800 hPa (not shown). The mixing ratio over the Kuroshio is about 0.2 g kg−1 larger than that in the SMTH run (Fig. 15c). About three-fourths of the positive equivalent potential temperature anomaly at 950 hPa result from this mixing ratio changes. Although the SST difference is about 1°C, the difference of the temperature at 950 hPa is only about 0.2°C (Fig. 15b). The contribution by the temperature to the equivalent potential temperature is only one-third by that of the humidity. Thus, the more humidity of the lower layer results in warmer marine atmospheric boundary layer in the sense of the equivalent potential temperature. Of course, since saturation vapor pressure increases with temperature, the warmer temperature in the marine atmospheric boundary layer contributes to the higher moisture.
As in Fig. 14, but for (a) equivalent potential temperature (°C), (b) temperature (°C), and (c) mixing ratio (g kg−1) at 950 hPa, and CAPE (m2 s−2). The equivalent potential temperature is calculated following Bolton (1980).
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
This warmer equivalent potential temperature in the marine atmospheric boundary layer suggests that the atmospheric condition over the Kuroshio in the CTL run is more convectively unstable than that in the SMTH run. As shown in Fig. 4c, the half of the high rainband over the Kuroshio is induced by convection process. To examine stability changes over the Kuroshio, Fig. 15d shows the difference of the convective available potential energy (CAPE) between the CTL and SMTH runs. CAPE is estimated using 3-h outputs, where a parcel is adiabatically raised from 950 hPa to the lifting condensation level, and is moist adiabatically raised to the level of neutral buoyancy without entrainment from environment. The change of the CAPE exhibits high positive values over the warm SST tongue of the Kuroshio, especially over the southern part of the Kuroshio (Fig. 15d). Hence, the atmospheric conditions over the Kuroshio in the CTL run are more unstable than those in the SMTH run and favor the convective instability. This difference of the CAPE is largely attributed to the difference of θe at the lower layer between the CTL and SMTH runs (Fig. 15a), which is mainly due to increased moisture as mentioned above. The CAPE changes and the moisture supply by the enhanced local evaporation indicate suitable conditions for energetic convection.
5. Summary and discussion
The influence of the Kuroshio in the East China Sea on the baiu rainband is examined using the satellite observations, the NCEP-CFSR reanalysis dataset, and the nonhydrostatic regional atmospheric model from 2003 to 2008. We focused on the precipitation field in June, when the baiu rainband located over the East China Sea and the monthly precipitation over the Kuroshio is the highest (Fig. 1). The satellite observations and reanalysis data reveal that the precipitation in June is enhanced over the warm SST tongue of the Kuroshio, and this enhancement is stronger than any other season (Figs. 3–4). The amplitude of the precipitation rate over the warm SST tongue of the Kuroshio is twice as large as that in its surrounding area, and the enhanced precipitation is embedded in the large-scale baiu rainband (Table 1). We also found the energetic surface evaporation, the enhanced surface wind speed, and the surface wind convergence over the Kuroshio (Figs. 5–6). This wind convergence likely results from the warm SST influence on the pressure in the marine atmospheric boundary layer through temperature and humidity changes (Fig. 7). In addition, the enhanced precipitation over the Kuroshio accompanies ascent motion and diabatic heating extending from the surface to the upper troposphere (Fig. 8). The horizontal patterns of the ascent motion and diabatic heating in the midtroposphere (Figs. 9b and 10a) resemble the enhanced precipitation pattern (Fig. 4b). The ascent motion and large diabatic heating over the Kuroshio in the troposphere in June are much stronger than any other month. The large diabatic heating in the midtroposphere is induced by the sum of the deep convective heating and the large-scale condensate heating (Fig. 8c).
Using the nonhydrostatic regional atmospheric model and the high-resolution SST product, we investigate the influence of the warm SST tongue of the Kuroshio on the high rainfall. The comparison between the CTL and SMTH runs indicates that the warm SST tongue acts to concentrate the precipitation there (Fig. 11), while the large-scale baiu rainband is not essentially changed. The surface wind convergence over the Kuroshio seen in the CTL run disappears in the SMTH run (Fig. 12). This result is consistent with the fact that the pressure modification by temperature and humidity changes plays an important role in inducing the wind convergence in the NCEP-CFSR product (Figs. 6–7). The moisture budget analysis reveals that the moisture is mainly transported to the East China Sea by large-scale low-level winds (Fig. 13 and Table 1), but the local energetic evaporation induced by the warm SST of the Kuroshio also significantly contributes to the increase of precipitation over the Kuroshio (Fig. 14). In addition, the large surface evaporation around the Kuroshio supplies more moisture to the marine atmospheric boundary layer (Figs. 15a–c), and this anomalous humidity results in the decrease of the vertical stability in the CTL run compared to that in the SMTH run (Fig. 15d). These results suggest that the energetic evaporation caused by the warm SST of the Kuroshio influences the baiu rainband through moisture supply, pressure modification, and stability change.
The atmospheric response to the warm SST tongue of the Kuroshio in June resembles the deep-heating mode, the atmospheric response to the Gulf Stream in summer shown by Minobe et al. (2010). In addition to the similarity, the amplitudes of the precipitation rate, ascent motion, and diabatic heating over the Kuroshio are larger those over the Gulf Stream region in the NCEP-CFSR product (Figs. 9–10). Minobe et al. (2010) pointed out two necessary conditions for the deep-heating mode. One is the warm SSTs above the threshold for the deep convection (26°–27°C), and the other is the moisture transport from low latitudes by a large-scale atmospheric circulation. The East China Sea in June meets both conditions (Figs. 2b and 13), consistent with the results of Minobe et al. (2010). In addition, these strong deep atmospheric responses found over the Kuroshio and the Gulf Stream are not observed in other regions of western boundary currents, such as the Kuroshio Extension, the Agulhas Return Current, and the Brazil–Malvinas Currents, in the TRMM 3B43 and NCEP-CFSR products (not shown). The East China Sea in June is therefore one of the regions of the strongest deep-heating mode over the midlatitude ocean.
The baiu rainband fluctuates in its location and intensity on interannual to decadal time scales (e.g., Wang et al. 2001; Tomita et al. 2004). For example, the precipitation rate in June at Naha, which is the capital city of Okinawa prefecture in Japan, on the Okinawa Island shows large fluctuations and was the highest in 2005 in the instrument record (Fig. 16). This record-breaking high precipitation is associated with floods and landslides in the island. The satellite observation indicates that the strong rain is confined over the Kuroshio and east of it, suggesting that the warm SST tongue of the Kuroshio plays an important role in this disastrous heavy rain. Thus, it is interesting to examine interannual variability of precipitation from 2003 to 2008 using the nonhydrostatic regional atmospheric model outputs, although the analysis used only one run for each year. The numerical results show the largest difference of the precipitation over the Kuroshio between the CTL and SMTH runs in 2005 (Fig. 17a), when the precipitation at Naha was the highest (Fig. 16). The maximum difference of the precipitation rate is about 8 mm day−1 over the warm SST tongue of the Kuroshio. This difference is twice as large as that of the mean value from 2003 to 2008 (Fig. 11c). This result suggests that the influence of the warm SST tongue of the Kuroshio also contributes to interannual variability in precipitation. In contrast to the mean moisture balance (Fig. 14), the rainfall changes over the Kuroshio mainly result from the change of the water vapor flux convergence (Figs. 17b,c). Note that the SST difference between the CTL and SMTH runs around the warm SST tongue in 2005 (contours in Fig. 17) is quantitatively similar with that in the mean difference. Thus, the SST variability itself is not the cause of the enhancement of precipitation. A detailed examination of the origin of the heavy rain and its relation to the Kuroshio remains for a future study.
Precipitation rate in June at Naha (26.2°N, 127.7°E) from rain gauge data of World Monthly Surface Station Climatology (gray) from 1891 to 2006 (downloaded from http://dss.ucar.edu/datasets/ds570.0), and the Automated Meteorological Data Acquisition System (AMeDAS) data from 1976 to 2008 (black) provided by the Japan Meteorological Agency.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
The difference of the (a) precipitation rate, (b) evaporation, and (c) vertical integrated water vapor flux convergence (color; mm day−1) between the CTL minus SMTH runs in 2005. The thick contours are for the corresponding SST difference at ±1°C.
Citation: Journal of Climate 25, 19; 10.1175/JCLI-D-11-00727.1
Our results showed that the humidity in the marine atmospheric boundary layer is an important variable for understanding the influence of the Kuroshio. For further understanding the convective process and resultant adjustment in the marine atmospheric boundary layer from observations, vertically and temporally high-resolution data for the humidity in the marine atmospheric boundary layer are required, but it is difficult to obtain this information from satellite observations. Nevertheless, there are quite few in situ observations for the atmospheric conditions over the open ocean in the East China Sea (e.g., Maeda et al. 2008). It is therefore desirable to use in situ observations for more detailed analyses of the influence of the Kuroshio on the atmosphere.
Acknowledgments
This research was supported by the Innovative Program of Climate Change Projection for the 21st Century (Kakushin 4 program), the Research Program on Climate Change Adaptation, the Grant-in-Aid for Scientific Research on Innovative Areas 22106008, the Grant-in-Aid for Scientific Research (A) 22244057, and the Grant-in-Aid for Young Scientists (B) 18740293, which are all funded by the Ministry of Education, Culture, Sports, Science, and Technology of Japan, and by the Global Environmental Research Fund S-5-3 of the Ministry of the Environment of Japan. The JMA/MRI NHM was used with permission of the Japan Meteorological Agency and technical support was provided by Drs. K. Saito, N. Ishizaki, and S. Hayashi. The model simulation was performed using the Hokkaido University High Performance Computing System.
REFERENCES
Akiyama, T., 1973: The large-scale aspects of the characteristics features of the Baiu front. Pap. Meteor. Geophys., 24, 157–188.
Bolton, D., 1980: The computation of equivalent potential temperature. Mon. Wea. Rev., 108, 1046–1053.
Chelton, D. B., and S.-P. Xie, 2010: Coupled ocean-atmosphere interaction at oceanic mesoscales. Oceanography, 23, 52–69.
Chelton, D. B., M. G. Schlax, M. H. Freilich, and R. F. Milliff, 2004: Satellite measurements reveal persistent small-scale features in ocean winds. Science, 303, 978–983.
Ducet, N., P. Y. Le Traon, and G. Reverdin, 2000: Global high-resolution mapping of ocean circulation from TOPEX/Poseidon and ERS-1 and -2. J. Geophys. Res., 105 (C8), 19 477–19 498.
Fairall, C. W., E. F. Bradley, J. E. Hare, A. A. Grachev, and J. B. Edson, 2003: Bulk parameterization of air-sea fluxes: Updates and verification for the COARE algorithm. J. Climate, 16, 571–591.
Graham, N. E., and T. P. Barnett, 1987: Sea surface temperature, surface wind divergence, and convection over tropical oceans. Science, 238, 657–659.
Kain, J., 2004: The Kain–Fritsch convective parameterization: An update. J. Appl. Meteor., 43, 170–181.
Kain, J., and J. Fritsch, 1993: Convective parameterization for mesoscale models: The Kain–Fritsch scheme. The Representation of Cumulus Convection in Numerical Models, Meteor. Monogr., No. 46, Amer. Meteor. Soc., 165–170.
Kelly, K. A., R. J. Small, R. M. Samelson, B. Qiu, T. M. Joyce, Y.-O. Kwon, and M. F. Cronin, 2010: Western boundary currents and frontal air–sea interaction: Gulf Stream and Kuroshio Extension. J. Climate, 23, 5644–5667.
Kobashi, F., S.-P. Xie, N. Iwasaka, and T. T. Sakamoto, 2008: Deep atmospheric response to the North Pacific oceanic subtropical front in spring. J. Climate, 21, 5960–5975.
Kosaka, Y., S.-P. Xie, and H. Nakamura, 2011: Dynamics of interannual variability in summer precipitation over East Asia. J. Climate, 24, 5435–5453.
Kubota, M., and H. Tomita, 2007: The present state of the J-OFURO air-sea interaction data product. Flux News, No. 4, WCRP, 13–15. [Available online at ftp://ftp1.esrl.noaa.gov/users/cfairall/wcrp_wgsf/newsletter/voi4_fluxnews_10_07.pdf.]
Kuwano-Yoshida, A., S. Minobe, and S.-P. Xie, 2010: Precipitation response to the Gulf Stream in an atmospheric GCM. J. Climate, 23, 3676–3698.
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.
Liu, W. T., X. Xie, and P. P. Niiler, 2007: Ocean–atmosphere interaction over Agulhas extension meanders. J. Climate, 20, 5784–5797.
Maeda, S., K. Tsuboki, Q. Moteki, T. Shinoda, H. Minda, and H. Uyeda, 2008: Detailed structure of wind and moisture fields around the Baiu frontal zone over the East China Sea. SOLA, 4, 141–144.
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.
Minobe, S., M. Miyashita, A. Kuwano-Yoshida, H. Tokinaga, and S.-P. Xie, 2010: Atmospheric response to the Gulf Stream: Seasonal variations. J. Climate, 23, 3699–3719.
Moteki, Q., H. Ueda, T. Maesaka, T. Shinoda, M. Yoshizaki, and T. Kato, 2004: Structure and development of two merged rainbands observed over the East China Sea during X-BAIU-99. Part II: Meso-α-scale structure and build-up processes of convergence in the Baiu frontal region. J. Meteor. Soc. Japan, 82, 45–65.
Nakamura, H., H. Ichikawa, A. Nishina, and H.-J. Lie, 2003: Kuroshio path meander between the continental slope and the Tokara Strait in the East China Sea. J. Geophys. Res., 108, 3360, doi:10.1029/2002JC001450.
Ninomiya, K., 1984: Characteristics of Baiu front as predominant subtropical front in the summer Northern Hemisphere. J. Meteor. Soc. Japan, 62, 880–894.
Ninomiya, K., and T. Akiyama, 1992: Multi-scale features of Baiu, the Summer Monsoon over Japan and the East Asia. J. Meteor. Soc. Japan, 70, 467–495.
Ninomiya, K., and Y. Shibagaki, 2007: Multi-scale features of the Meiyu-Baiu front and associated precipitation systems. J. Meteor. Soc. Japan, 85, 103–122.
Nonaka, M., and S.-P. Xie, 2003: Covariations of sea surface temperature and wind over the Kuroshio and its extension: Evidence for ocean–atmosphere feedback. J. Climate, 16, 1404–1413.
O’Neill, L. W., D. B. Chelton, and S. K. Esbensen, 2003: Observations of SST-induced perturbations of the wind stress field over the Southern Ocean on seasonal timescales. J. Climate, 16, 2340–2354.
O’Neill, L. W., D. B. Chelton, S. K. Esbensen, 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.
Onogi, K., and Coauthors, 2007: The JRA-25 reanalysis. J. Meteor. Soc. Japan, 85, 369–432.
Park, K.-A., P. C. 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.
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.
Rio, M.-H., and F. Hernandez, 2004: A mean dynamic topography computed over the world ocean from altimetry, in situ measurements, and a geoid model. J. Geophys. Res., 109, C12032, doi:10.1029/2003JC002226.
Saha, S., and Coauthors, 2010: The NCEP Climate Forecast System Reanalysis. Bull. Amer. Meteor. Soc., 91, 1015–1057.
Saito, K., and Coauthors, 2006: The operational JMA nonhydrostatic mesoscale model. Mon. Wea. Rev., 134, 1266–1298.
Saito, K., J. Ishida, K. Aranami, T. Hara, T. Segawa, M. Narita, and Y. Honda, 2007: Nonhydrostatic atmosperic models and operational development at JMA. J. Meteor. Soc. Japan, 85, 271–304.
Samelson, R. M., E. D. Skyllingstad, D. B. Chelton, S. K. Esbensen, L. W. O’Neill, and N. Thum, 2006: On the coupling of wind stress and sea surface temperature. J. Climate, 19, 1557–1566.
Sampe, T., and S.-P. Xie, 2007: Mapping high sea winds from space: A global climatology. Bull. Amer. Meteor. Soc., 88, 1965–1978.
Sampe, T., and S.-P. Xie, 2010: Large-scale dynamics of the meiyu–baiu rainband: Environmental forcing by the westerly jet. J. Climate, 23, 113–134.
Shimada, T., and S. Minobe, 2011: Global analysis of the pressure adjustment mechanism over sea surface temperature fronts using AIRS/Aqua data. Geophys. Res. Lett., 38, L06704, doi:10.1029/2010GL046625.
Small, R. J., and Coauthors, 2008: Air–sea interaction over ocean fronts and eddies. Dyn. Atmos. Oceans, 45, 274–319.
Takatama, K., S. Minobe, M. Inatsu, and R. J. Small, 2012: Diagnostics for near-surface wind convergence/divergence response to the Gulf Stream in a regional atmospheric model. Atmos. Sci. Lett., 13, 16–21, doi:10.1002/asl.355.
Tanimoto, Y., T. Kanenari, H. Tokinaga, and S.-P. Xie, 2011: Sea level pressure minimum along the Kuroshio and its extension. J. Climate, 24, 4419–4434.
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.
Tokinaga, H., Y. Tanimoto, S.-P. Xie, T. Sampe, H. Tomita, and H. Ichikawa, 2009: Ocean frontal effects on the vertical development of clouds over the western North Pacific: In situ and satellite observations. J. Climate, 22, 4241–4260.
Tomita, T., T. Yoshikane, and T. Yasunari, 2004: Biennial and lower-frequency variability observed in the early summer climate in the western North Pacific. J. Climate, 17, 4254–4266.
Waliser, D. E., N. E. Graham, and C. Gautier, 1993: Comparison of the highly reflective cloud and outgoing longwave radiation datasets for use in estimating tropical deep convection. J. Climate, 6, 331–353.
Wallace, J., T. Mitchell, and C. Deser, 1989: The influence of sea-surface temperature on surface wind in the eastern equatorial Pacific: Seasonal and interannual variability. J. Climate, 2, 1492–1499.
Wang, B., Ed., 2006: The Asian Monsoon. Springer-Praxis, 844 pp.
Wang, B., R. Wu, and K.-M. Lau, 2001: Interannual variability of the Asian summer monsoon: Contrasts between the Indian and the western North Pacific–East Asian monsoons. J. Climate, 14, 4073–4090.
Xie, S.-P., 2004: Satellite observations of cool ocean–atmosphere interaction. Bull. Amer. Meteor. Soc., 85, 195–208.
Xie, S.-P., J. Hafner, Y. Tanimoto, W. T. Liu, H. Tokinaga, and H. Xu, 2002: Bathymetric effect on the winter sea surface temperature and climate of the Yellow and East China Seas. Geophys. Res. Lett., 29, 2228, doi:10.1029/2002GL015884.
Xu, H., M. Xu, S.-P. Xie, and Y. Wang, 2011: Deep atmospheric response to the spring Kuroshio Current over the East China Sea. J. Climate, 24, 4959–4972.
