1. Introduction
The South Asian high (SAH; also called the Asian monsoon anticyclone, Tibetan anticyclone, or Tibetan high) is a planetary-scale anticyclonic circulation system in the upper troposphere, with its major body located over the Tibetan Plateau (TP) and its neighboring areas during boreal summer (Mason and Anderson 1963; Tao and Zhu 1964; Krishnamurti 1971; Zhu et al. 1980; Zarrin et al. 2010; Liu et al. 2013). The SAH is also characterized by a relatively low potential vorticity (PV) with strong PV gradients toward higher latitudes (Hsu and Plumb 2000; Liu et al. 2007; Garny and Randel 2013). As a prominent component of the Asian summer monsoon system, the SAH can considerably affect the weather, climate, and chemical constituents over the Asian–Pacific region (Zhang et al. 2005; Randel and Park 2006; Zhao et al. 2009; Liu et al. 2013). The SAH has been the focus of much literature because of its multiple spatiotemporal variability and impacts. For example, the SAH shows decadal variation in its intensity accompanied by decadal oscillation of the Indian Ocean–Pacific oceanic thermal anomalies (Zhang et al. 2000; Jiang et al. 2011; Qu and Huang 2012). The interannual variation of the SAH has been studied intensively in recent years. Its connections with the anomalies of the sea surface temperature (Zhang et al. 2000; Yang et al. 2007; Peng et al. 2010; Huang et al. 2011; Jiang et al. 2011), the TP surface condition (Duan et al. 2008; Wang et al. 2014), and the Indian summer monsoon (Wei et al. 2014) were revealed. The impacts of the SAH variability on interannual-to-decadal time scales were also well documented (Luo et al. 1982; Zhang et al. 2005; Jiang et al. 2011; Wei et al. 2014).
Besides the long-term variations, the SAH also exhibits subseasonal variation. However, issues such as the features of the SAH’s subseasonal zonal oscillation and related mechanisms are not yet deeply understood. Mason and Anderson (1963) found that during the summer of 1958, the SAH departed from the TP and extended westward for periods of one week or more. Tao and Zhu (1964) documented that the SAH oscillated zonally around its mean position with a quasi-biweekly period. Such a zonal oscillation of the SAH was classified into three patterns, namely the eastern, the western, and the transitional pattern, which were considered as useful indicators of rainfall patterns over eastern China (Luo et al. 1982). In terms of the pentad reanalysis data, Zhang et al. (2002) further verified that the variations in the location of the climatological center of the SAH at 100 hPa are bimodal and classified them into two modes, the Tibetan mode and the Iranian mode. In the upper-level geopotential height field, the eastward (westward) extension of the SAH has a close conjunction with an anomalous anticyclonic (cyclonic) circulation centered over subtropical eastern Asia, while in the PV field it demonstrates eastward (westward) shedding of the low PV region from the Tibetan anticyclone (Hsu and Plumb 2000; Popovic and Plumb 2001; Liu et al. 2007). The zonal oscillation of the SAH is associated with low-frequency oscillation of other monsoon components, together contributing to the rainfall anomalies over the Asian region (Liu and Lin 1991; Kim et al. 2010; Yuan et al. 2012; Jia and Yang 2013; Chen and Zhai 2014).
The candidate mechanisms for understanding the zonal oscillation of the SAH on subseasonal time scale include (i) the circulation modulation induced by the coupled oscillation system between the circulation and the diabatic heating and (ii) the influences of the midlatitude wave train over the Eurasian continent. The formation of the SAH is tied to the convective heating associated with the Asian summer monsoon, and also affected by topography such as the TP and the surface sensible heating over the continents (Liu et al. 2004; Liu and Wu 2004; Duan and Wu 2005; Liu et al. 2007; Wu et al. 2012; Duan et al. 2013; Liu et al. 2012, 2013; Chen and Bordoni 2014; Zhang et al. 2014). Simple models closely reproduced the climatological features of the anticyclone in the upper level as a wave response to diabatic heating of monsoon system. The anticyclone, as well as other middle-to-low latitude responses, is generally established within several days (Gill 1980; Hoskins and Rodwell 1995; Jin and Hoskins 1995; Lin 2009). Therefore, the SAH is considered as a thermally driven circulation. The variability in the diabatic forcing can be expected to be an important candidate causing subseasonal movement of the SAH.
Using a shallow-water equation model, Hsu and Plumb (2000) discussed the nonlinear dynamics of the upper-tropospheric anticyclone under nonaxisymmetric conditions. They demonstrated that the anticyclone becomes unstable and periodically sheds PV eddies into the background flow when the imposed asymmetry that represents thermal forcing is sufficiently large. By performing a sequence of idealized experiments with a global primitive equation model, Liu et al. (2007) found that if the diabatic heating over the plateau is sufficiently strong, the flow above the plateau is unstable and produces an oscillation on subseasonal time scales. They suggested that the westward “eddy shedding” of the minimum PV area mentioned in Hsu and Plumb (2000) is due to the tendency to reduce the PV above the heating over the plateau and to advection by the consequent anticyclone of high PV around from the east and low PV to the west. The above studies shed light on the features and causation of the SAH’s westward extension on subseasonal time scales. In reality, the eastward extension of the SAH was also observed frequently, accompanied by the westward migration of the western Pacific subtropical high (WPSH) at 500 hPa (Tao and Zhu 1964; Tao and Wei 2006; Liu et al. 2006; Ren et al. 2007; Yuan et al. 2012; Jia and Yang 2013). Accompanied by the opposite movement of the SAH and the WPSH, the precipitation over East China exhibits an unequal distribution both in the spatial and temporal aspects (Liu et al. 2006; Tao and Wei 2006; Chen and Zhai 2014). The nonuniform rainfall pattern could induce anomalous distribution in diabatic heating that may thus have a feedback effect on the movement of the SAH. However, the anomalous pattern of diabatic heating that is connected with the eastward extension of the SAH is still unclear. Furthermore, the role of diabatic heating feedback in the SAH’s eastward extension needs to be explored.
The other possible mechanism for the SAH’s eastward extension involves the subseasonal propagation of Rossby wave train over the Eurasian area (Tao and Zhu 1964; Lu et al. 2002; Fujinami and Yasunari 2004; Tao and Wei 2006; Ding and Wang 2007; Bueh et al. 2008; Yasui and Watanabe 2010; Watanabe and Yamazaki 2014). Case studies showed that when the SAH extends eastward toward eastern Asia, a wave train is seen with two anomalous anticyclonic circulations straddling the western and eastern flanks of the TP and a cyclonic one controlling the area above the TP region (Raman and Rao 1981; Tao and Wei 2006; Liu et al. 2006; Ding and Wang 2007; Watanabe and Yamazaki 2012). The anomalous high on the upstream side of the TP precedes the appearance of the ridges over eastern Asia by several days (Raman and Rao 1981; Krishnan et al. 2009). Previous studies suggested that the wave train originates from the northeastern Atlantic and crosses Europe to central Asia (Ding and Wang 2007; Watanabe and Yamazaki 2014). When the wave train travels quickly eastward along the jet stream, the Rossby wave energy dispersion establishes the anomalous ridge over northeastern Asia (Enomoto et al. 2003; Ding and Wang 2007). As a result, there is always a prolonged heatwave in the middle and lower reaches of the Yangtze River and above-normal rainfall in northern China (Tao and Wei 2006; Liu et al. 2006; Ding and Wang 2007).
The aforementioned studies showed that the propagation of the wave train over the Eurasian area is accompanied by a monsoon rainfall anomaly over eastern Asia. It is well known that condensational heating induced by monsoon rainfall is the major component of diabatic heating during summer period (Yanai and Tomita 1998; Jin et al. 2013). We speculate that the midlatitude wave train and the monsoon rainfall–induced diabatic heating feedback over eastern Asia may cooperate in causing the eastward extension of the SAH. Previous studies mentioned some features of propagation of the wave train over the Eurasian area and its connection with the establishment of anomalous anticyclone over northeastern Asia (Tao and Wei 2006; Ding and Wang 2007). Therefore, in this study we focus more on the mechanism related to the diabatic heating feedback. Meanwhile, we suggest that the two mechanisms of wave train and diabatic heating feedback work together to give rise to the zonal oscillation of the SAH on subseasonal time scales. In summary, we intend to address three issues: 1) the features of the zonal movement, especially the eastward extension of the SAH at 200 hPa on subseasonal time scale, and the propagating wave train linked with the eastward extension of the SAH; 2) the anomalous patterns in rainfall and diabatic heating associated with the eastward extension of the SAH; and 3) the role of diabatic heating feedback in contributing to the eastward extension of the SAH.
For better understanding of the above three issues, we use both the geopotential height and PV fields to show the feature of the SAH’s eastward extension, and further perform a PV diagnostic analysis to investigate the connection of the SAH’s movement to diabatic heating (Hoskins et al. 1985). The tool of PV diagnosis has been applied to understanding the dynamics of subseasonal modes, like the Madden–Julian oscillation (MJO) and the boreal summer intraseasonal oscillation over Asian monsoon region (Zhang and Ling 2012; Seo and Song 2012). As mentioned in Zhang and Ling (2012), it is very useful to explore the effect of diabatic heating on the tropical and extratropical circulations, including their generation and evolution (Hoerling 1992; Wu and Wang 2000; Tory et al. 2012). The paper is organized as follows. Data sources and analysis techniques are described in section 2. Section 3 presents the features of the SAH’s eastward extension in terms of geopotential height and PV fields. Wave train propagation and large-scale circulation linked with the SAH’s eastward extension are also examined in section 3. Section 4 analyzes the subseasonal anomalies in rainfall and diabatic heating associated with the SAH’s eastward extension. The role of diabatic heating feedback in the SAH’s eastward extension is investigated in section 5. Finally, conclusions and discussion are provided in section 6.
2. Data and analysis methods
a. Data
The datasets used in this study include 1) daily atmospheric data including isobaric PV value from the global European Centre for Medium-Range Weather Forecasts (ECMWF) interim reanalysis (ERA-Interim; Dee et al. 2011) for 1979–2012, 2) daily gridded precipitation data from the project Asian Precipitation—Highly Resolved Observational Data Integration Toward Evaluation of Water Resources (APHRODITE; http://www.chikyu.ac.jp/precip/) conducted by the Research Institute for Humanity and Nature (RIHN) and the Meteorological Research Institute (MRI) of the Japan Meteorological Agency (JMA) since 2006 (Xie et al. 2007; Yatagai et al. 2009, 2012), we use the APHRO_V1101 dataset that covers the region of 15.0°S–55.0°N, 60.0°–150.0°E with resolution of 0.5° × 0.5° for 1979–2007, and 3) daily mean outgoing longwave radiation (OLR) data on a 2.5° × 2.5° grid from NOAA (Liebmann and Smith 1996) for 1979–2012.
b. Calculations of diabatic heating and PV budget
c. Definition of the SAH’s eastward extension index
To extract the component of subseasonal variations, we use the method adopted by Krishnamurthy and Shukla (2000) and Ren et al. (2013). The procedure is as follows. Processes are first removing very high-frequency fluctuations in the daily data by applying a 5-day running mean, second deducing the daily climatology, and finally removing the interannual signal by subtracting seasonal anomaly.
When the SAH stretches anomalously eastward, a large area with positive geopotential height anomaly prevails over eastern Asia in the upper level. Therefore, an SAH’s eastward extension index (EI) is defined by normalized subseasonal geopotential height anomaly at 200 hPa over subtropical eastern Asia (22.5°–32.5°N, 100°–120°E; box in Figs. 1a,b). The SAH’s eastward and westward events are picked according to the EI index. An eastward (westward) event is identified when the EI is more (less) than 1.0 (−1.0) in at least three consecutive days, and within those three (or more) days at least one day’s EI is more (less) than 1.5 (−1.5). The above criteria ensure the intensity and duration of the events. For an individual event, the day with the biggest (or smallest) EI is called the peak day. Similar to Watanabe and Yamazaki (2012), the interval between each peak day was set to more than 10 days, which renders the subseasonal period for individual events. If the interval of two peak days is smaller than 10 days, their corresponding events are considered as one event. More information about the definition of the EI will be shown in Figs. 1 and 2.
Climatological summer fields at 200 hPa of (a) geopotential height (Z200; gpm, purple contour), (b) potential vorticity (PV200; PVU, purple contour), and (c) horizontal (vector; m s−1) and vertical (shaded; hPa s−1, purple contour) wind fields. The subseasonal standard deviation of Z200 and PV200 (shaded and thin green contour) overlaid on their climatological fields in (a) and (b), respectively. The intervals for green contours are 10 gpm for Z200 and 0.4 PVU for PV200, and for purple contours in (c) are 0.02 m s−1 with zero contours omitted. The red box denotes the area for definition of the SAH eastward EI. The thick gray line denotes an orographic isoline of 4 km corresponding to the main body of the Tibetan Plateau.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
(a) Longitudinal distributions (averaged over 20°–32.5°N) of climatological summer Z200 (blue line, gpm), subseasonal standard deviation of Z200 (purple line, gpm), and correlation coefficients between Z200 anomalies and EI (red line). (b) Power spectra (solid line) of EI (blue) and subseasonal Z200 anomaly averaged over 20°–32.5°N, 80°–130°E (red) . The dotted lines are the spectrum of 95% confidence level.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
The method of time-lagged composition based on the eastward and westward events and the time-lagged regressions of subseasonal variables on EI are used in this study. In the following, day 0 denotes simultaneous composite and regression. A negative lag day (day −n) indicates variable leads EI (or events) and is obtained by shifting backward the number of leading days, and vice versa for a positive lag day (day +n).
3. Features of eastward extension of the SAH
a. Basic features
Before analyzing the features of eastward extension of the SAH on subseasonal time scales, we first depict the climatology of the SAH and other related fields. Figure 1 shows the climatological summer fields at 200 hPa of geopotential height and PV (Z200 and PV200), and horizontal and vertical wind fields. The subseasonal standard deviations of Z200 and PV200 are overlapped on their respectively climatological fields. The SAH in Fig. 1a is characterized by a large anticyclone zonally elongated from Africa to western Pacific with a ridge axis along 28°N. The 12 520-gpm isoline (hereafter 12520 line) at 200 hPa, which is used to represent the main body of the SAH in the upper troposphere, extends zonally from the Arabian Peninsula to southern China. The Asian subtropical westerly jet is located on the poleward side of the ridge axis. Easterlies prevail south of the ridge axis. Strong upward motion is seen on the east and southeast flanks of the SAH. In the PV200 field, the SAH is a relatively low PV region, with values ranging from 0.2 to 2 PVU.
The subseasonal standard deviation of Z200 and PV200 presents relatively larger values to the east of the TP, at the east edge of the SAH’s main body. Consequently, we define the SAH’s eastward extension index as described in section 2c. Figure 2a shows the longitudinal distribution of subseasonal standard deviation of Z200 averaged over 20°–32.5°N. It indicates also that the subseasonal signal over the east edge of the SAH (100°–120°E) is comparatively stronger than that over the SAH’s central area (80°–90°E). Power spectra for the EI and subseasonal Z200 anomaly averaged over another larger area (20°–32.5°N, 80°–130°E) are calculated and plotted in Fig. 2b. A periodicity of 10–30 days associated with quasi-biweekly oscillation can be noted in the power spectra of EI, which was considered as a dominant time scale of the subseasonal zonal oscillation of the SAH (Tao and Zhu 1964; Luo et al. 1982; Liu and Lin 1991; Liu et al. 2007; Jia and Yang 2013). The SAH’s central region has a longer period of intraseasonal oscillation. The correlation coefficients between the EI and the Z200 anomalies suggest that the EI reflect the variation over the east edge of the SAH (Fig. 2a). Therefore, the EI can be used to investigate the subseasonal zonal oscillation of the east edge of the SAH.
We also use PV200 to recalculate the EI (denoted as EI-PV200). The correlation coefficient is 0.75 between EI and EI-PV200, indicating that they are highly consistent in the variation. The levels of 200, 100, or 150 hPa are all popularly used to study the SAH. The correlation coefficient is 0.88 between EI from Z200 and EI from geopotential height at 100 hPa. Therefore, all the levels show similar variability in the index. The EI calculated by Z200 is used in this study. As an example, Fig. 3 displays time series of the daily Z200 and EI, as well as an identified eastward event and a westward one for the summer of 2003. During the summers of 1979–2012, we identified 40 eastward events for a total of 278 days and 43 westward events of 266 days. The 40 eastward events represent the SAH’s eastward extension cases, while the 43 westward events represent westward retraction cases.
Time series of daily Z200 for the region over 22.5°–32.5°N, 100°–120°E (box in Figs. 1a,b) (purple dotted line, gpm) and the SAH eastward EI (red line with crosses) for the summer of 2003. The thick red and blue lines indicate the identified SAH eastward and westward events, respectively. The scales on left side are for daily Z200 and EI, from left to right.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
Figure 4 shows the lagged composites of the 12520 line at 200 hPa, and 5880- and 5860-gpm isolines at 500 hPa for 40 eastward events from day −15 to day +15 with interval of 3 days. The 5880- and 5860-gpm isolines (jointly referred to as the 5880–5860 line) at 500 hPa anchor the western part of the WPSH. The process of the SAH’s eastward extension and then westward retreat is visually depicted in Fig. 4. At the early stage (before day −6), the SAH stretches eastward slowly. On day −15, the east point of the 12520 line is at 100°E. On day −9, it is near 110°E. At the later stage (during days −6 to 0), the SAH moves eastward with big strides. The east point of the SAH is over the oceanic region on days −3 and 0. The WPSH moves in the opposite direction, when the SAH stretches eastward. The west points of the 5880–5860 line remain steady during days −15 to −9. The western part of the WPSH, especially the 5860-gpm line, moves westward conspicuously after day −6 and reach the westernmost location on day 0. After day 0, the SAH moves back toward the TP region, while the WPSH retreats toward the oceanic region.
Composites of (left) 12 520-gpm contour lines at 200 hPa and (right) 5860- and 5880-gpm contour lines at 500 hPa for 40 eastward events at various lags, at (a),(c) leading times from day −15 to day 0 with intervals of 3 days and (b),(d) lag times from day 0 to day +15.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
The simultaneous fields of the subseasonal elements associated with the 40 eastward events are shown in Fig. 5. We also plot the simultaneous composites for the 43 westward events in Fig. 5. The main body of the SAH for the eastward events occupies a larger area than that for the westward events (Figs. 5a,d). The 5860-gpm line of the WPSH advances more toward the inland for the eastward events than for the westward events (Figs. 5b,e). Thus, an anomalous anticyclonic circulation is centered over eastern China and ranges from the central TP to the coastal waters of eastern Asia for the eastward events, while a cyclonic one exists for the westward events in the middle-to-high troposphere. The composited geopotential height fields at three levels (200, 500, and 850 hPa) display a noticeable baroclinic structure from the central TP to the coastal waters for both westward and eastward events. Meanwhile, the centers in the middle-to-high latitudes over the Eurasian area exhibit equivalent barotropic structures for both events. The 2-PV isoline to the east and northeast of the TP has a slightly anticyclonic warp for the eastward events (Fig. 5a) and a greater cyclonic warp for the westward ones (Fig. 5d).
Simultaneous compositions for (left) 40 eastward and (right) 43 westward events. (a),(d) Subseasonal Z200 (black contour with interval of 10 gpm) and OLR (shaded; W m−2) anomalies superimposed on 12 520-gpm (red) and 2-PVU (purple) contour lines. (b),(e) Subseasonal Z500 (black contour with interval of 6 gpm) and rainfall (shaded; mm day−1) anomalies superimposed on 5880–5860-gpm contour lines (thick red). (c),(f) Subseasonal Z850 (contour with interval of 4 gpm) and 850-hPa wind field anomalies superimposed on 1505-gpm contour lines (thick orange).
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
For the composites of the eastward events, negative OLR anomalies are found to the north of 30°–32°N and positive anomalies to the south of 30°–32°N, between 80° and 140°E (Fig. 5a), accompanied by a similarly dipolar structure of enhanced rainfall north of the Yangtze River and reduced rainfall south of it over eastern China (Fig. 5b). The OLR anomalies also occur over the Indian Ocean region, with a reduced (enhanced) value to the east (west) of about 70°E. In the low-level wind fields, the detectable features over the Asian monsoon region are an anomalous southerly (northerly) flow to the south (north) of the Yangtze River and an anomalous cyclonic circulation zonally elongated from the Arabian Sea to the South China Sea. The simultaneous fields of OLR, rainfall, and low-level wind anomalies for the westward events are almost mirror images of those for the eastward events.
b. Wave train propagation and large-scale circulation
Figure 6 displays the regressed fields of Z200, PV200, and 200-hPa wind anomalies on EI. The regressed Z200 and wind fields illustrate that the anomalous anticyclonic circulation centered over eastern China on day 0 is linked with a wave train across the Eurasian continent. On days −12 and −9, the wave train is seen over the northern Atlantic Ocean to the eastern flank of the TP via western Europe, northeastern Europe–western Russia, central Asia, and the western flank of the TP. The individual centers of the wave train west of TP form an arch shape following a “great circle” route. During this period, the anomalous anticyclonic and cyclonic circulations straddling respectively over the western and eastern flanks of TP are still weak. From days −9 to −6, the wave train centers in the midlatitudes are intensified and propagate eastward slightly. The weak high over the western flank of the TP moves eastward rapidly across the northern TP along the subtropical westerly jet, which acts as the Rossby waveguide (Hoskins and Ambrizzi 1993; Zhang et al. 2006; Ding and Wang 2007; Yasui and Watanabe 2010). The propagation of the weak high from the western flank of the TP to the east of the TP contributes to the eastward extension of the SAH during days −12 to −6 shown in Fig. 4a. When crossing the TP, the amplitude of the high center grows slowly, thus leading to the tardy extension of the SAH at its early stage.
Regression of (left) Z200 (contour) and (right) PV200 (contour), and 200-hPa wind field anomalies (vector; m s−1) on EI, from day −12 to day +3 with interval of 3 days. The contour intervals are 5 gpm and 0.1 PVU for Z200 and PV200, respectively. The dashed lines denote negative values. Shaded areas denote regions of statistically significant at 95% confidence level from a Student’s t test for Z200 and PV200. Vectors are plotted only where statistically significant at 95% confidence level.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
From days −6 to −3, the wave train centers in the middle-to-high latitudes propagate eastward continuously. Meanwhile, the intensity of the anomalous high center over eastern China increases dramatically. On day −3, a high–low–high pattern is established with the centers over central Asia of 50°E, northwest of the TP and eastern China, respectively. On day 0, the anomalous high centered over eastern China gets its strongest, while the anomalous low that was located northwest of TP propagates to the Lake Baikal region. The anomalous high centered over central Asia of 50°E migrates to 60°E with slightly weakened intensity. The regression charts of PV fields (Figs. 6g–l) also display a wave train originating from upstream and then downward to East Asia. The centers of negative PV anomalies in the PV wave train are corresponding to the centers of anomalous anticyclone in the Z200 wave train, and vice versa. The negative PV center over eastern China also gets it peak on day 0. The rapid intensification of the Z200 and PV200 perturbations to the east of TP during days −6 to 0 is robustly seen in Fig. 6, corresponding to the remarkable eastward extension of the SAH at its later stage (Fig. 4). On day +3 and afterward, the wave train over Eurasia collapses and a downstream one appears in both the regressed Z200 and PV200 fields.
To illustrate energy propagation along the wave train, the horizontal component of wave-activity flux of quasi-stationary Rossby waves are calculated based on the definition by Takaya and Nakamura (2001). According to Takaya and Nakamura (2001) and related studies (Nakamura and Fukamachi 2004; Watanabe and Yamazaki 2014), the wave-activity flux is parallel to the group velocity of Rossby waves in the plane wave limit. It can indicate an amplification of wave energy. Thus, the flux is a useful tool for illustrating the propagation of wave train in a zonally varying time-mean flow, especially during summertime. Figure 7 shows compositions of subseasonal anomalies of streamfunction and wave-activity flux at 200 hPa for 40 eastward events. During days −12 to −9, the development of the anomalous anticyclone over northeastern Europe coincides with strong wave-activity flux across the upstream region and with convergence to the center of the anticyclone, suggesting that the inflow of wave energy originates from the western North Atlantic Ocean. The anomalous anticyclone west of the TP is associated with a weak but detectable inflow of wave energy directing from central Asia along the summertime subtropical jet (Zhang et al. 2006). During days −6 and −3, the flux moves eastward and southeastward in association with the propagation of the wave train and finally reaches the northeast of the TP region, being favorable to the intensification of the anomalous high Z200 (negative PV200) locally. After day −3, the flux signal is much weak over the negative PV region to the east of the TP (figure not shown).
Compositions of subseasonal anomalies of streamfunction (contour with interval of 12 × 105 m2 s−1) and wave-activity flux (vector; m2 s−2) at 200 hPa for 40 eastward events, at days (a) −12, (b) −9, (c) −6, and (d) −3. Wave-activity fluxes were calculated according to Takaya and Nakamura (2001).
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
Figure 8 presents the regression of anomalies in geopotential height at 500 and 850 hPa (Z500 and Z850), vertical velocity at 400 hPa, and wind field at 850 hPa on EI. The arch-shaped wave train across the Eurasian continent at 200 hPa is also clearly seen in the midtroposphere, while the weak high disturbance over the western flank of TP at 200 hPa becomes rather indistinct at this level because of the topography of the TP. Accompanied by the wave train of geopotential height at 500 hPa, a wave train–like structure of 400-hPa vertical velocity spans from the eastern European area, west of Lake Baikal, to subtropical eastern Asia, with anomalous downward (upward) motion in the east and southeast of the anomalous high (low). Over subtropical eastern Asia, a dipolar structure of anomalous upward (downward) motion to the north (south) of 30° around the TP region incipiently forms on day −6. During days −6 to −3, the dipolar structure becomes robust and propagates eastward. On day −3, the northern lobe with anomalous upward motion stretches zonally from the central and northern part of the TP to northern China, while the southern lobe with downward motion is zonally elongated from the south flank of the TP to southern China. On day 0, the main body of the dipolar structure is located over 90°–130°E. On day +3, the wave train–like structure of vertical velocity is indistinct. The dipolar structure over subtropical eastern Asia disappears.
As in Fig. 6, but for (left) 500-hPa geopotential height (contour) and 400-hPa vertical velocity (shaded; 0.01 hPa s−1) anomalies and (right) 850-hPa geopotential height (contour) and 850-hPa wind field (vector; m s−1) anomalies, from day −9 to day +3 with intervals of 3 days. The contour intervals are 3 gpm for geopotential height with negative contours dashed. The shading areas (vectors) are significant at 95% confidence level for 400-hPa vertical velocity (850-hPa wind field).
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
The wave train pattern of the circulation at 850 hPa associated with the eastward extension of the SAH is weak but still detectable. Eastward propagation of the wave train centers over the Eurasian landmass is also discernible. Similar to the composition results in Fig. 5, the geopotential height anomalies along the wave train exhibit an equivalent barotropic vertical structure in the middle-to-high latitudes. The equivalent barotropic structure in the wave train can be linked with Rossby wave energy dispersion from upstream (Ding and Wang 2007; Jiang and Lau 2008). In contrast, the anomalies have a distinct baroclinic character over subtropical eastern Asia, with a low center in the lower troposphere being overlain by a high center in the middle and upper levels. The baroclinic structure over subtropical eastern Asia combined with anomalous upward or downward motions in this area suggests anomalies in rainfall and diabatic heating, which we will investigate in the next section.
4. Rainfall and diabatic heating anomalies associated with SAH eastward extension
In this section, we analyze the subseasonal anomalies in rainfall and diabatic heating associated with the eastward extension of the SAH. Figure 9 plots the distribution of the regressed fields of anomalies in rainfall and 850-hPa wind fields on EI at time lags ranging from days −12 to +9 with interval of 3 days. The corresponding lagged composites of the 12520 line and 5880–5860 line for 40 eastward events are also superimposed on the regressed rainfall and wind fields. On days −12 and −9, reduced rainfall scatters over the TP and to the southeast of the TP. The dipolar structure of enhanced or reduced rainfall over eastern Asia shown in Fig. 5 is incipiently formed on day −6, which is consistent with the zonally north–south dipolar structure of vertical motion over the same area in Fig. 8b. The rainfall signals are mostly located around and to the south and north of the TP region. From days −6 to 0, the eastward extension of the SAH and the westward advance of the WPSH are both distinct, as mentioned in section 3. Meanwhile, the rainfall dipolar structure becomes prominent and also propagates eastward, similar to the results of regressed vertical motion in Fig. 8. Besides, the rainfall anomaly in the southern lobe of the dipole (22.5°–30°N, 80°–120°E) is better organized and moves slightly ahead of the one in the northern lobe (32.5°–45°N, 80°–120°E) during this week. The anomalous southerly flow at 850 hPa prevails over northern China on day −6, because of the anomalous low located to the northwest of the TP. Meanwhile, the monsoon flow over southern China is switched off anomalously, being unfavorable to the local rainfall condition. On day −3, the anomalous southerly flow prevails from southern China to northern China. It originates from the SCS and the western Pacific Ocean, being favorable to strong moisture transportation to northern China (Zhou and Yu 2005).
Regression of rainfall anomalies (shaded areas, which are significant at 90% confidence level; mm day−1) on EI, from day −12 to day +9 with interval of 3 days, superimposed on regressed 850-hPa wind field anomalies (vector, which are significant at 95% confidence level; m s−1). The thick red and orange contours indicate the locations of corresponding lagged composites of the 12520 line at 200 hPa and 5860–5880 line at 500 hPa for 40 eastward events, respectively.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
On day 0, an anomalous northerly flow associated with the anomalous low over the Lake Baikal region is dominant over Mongolia and northern China and conflicts with the southerly flow over the Yangtze–Huaihe River region over eastern China. This anomalous convergence and ascending motion over eastern China provide necessary dynamical and thermal conditions for the persistent rainfall north of Yangtze River region (Tao and Chen 1987). Thus, the rainfall anomaly in the northern lobe (32.5°–45°N, 80°–120°E) of the dipolar structure reaches maximum intensity on day 0. Meanwhile, the eastern part of the SAH and the western part of the WPSH almost overlap. After day 0, the dipolar structure of rainfall anomalies over eastern Asia disappears soon and is replaced by an opposite pattern of reduced (enhanced) rainfall to the north (south) of 30°–32°N over eastern China. Besides the above dipolar structure over eastern Asia, increased rainfall over Indo-China and the Philippines between 5° and 20°N is scattered but persistent from days −12 to −3, accompanied by strengthened westerly locally and anomalous ascending motion shown in Fig. 8. After day 0, a decreased rainfall is seen in this band.
Figure 10 displays the longitude–time section of regressed anomalies of rainfall, OLR, and vertical velocity (ω) for the three bands from south to north: 5°–20°N, 22.5°–30°N, and 32.5°–45°N. According to the above analyses, the enhanced (reduced) rainfall is linked with the anomalous ascending (descending) motion and below-normal (above normal) OLR in the three bands. The intensified convective activity (denoted by below-normal OLR) and anomalous ascending motion occur foremost in the band of 5°–20°N around day −12, and are zonally located over the Bay of Bengal and the SCS and subtropical western Pacific regions (Fig. 10c). The signals propagate eastward and westward, respectively, and reach their individual maximums several days before day 0. They become weak after day −3 and disappear on day 0. For the band 22.5°–30°N (Fig. 10b), the reduced convection and anomalous descending motion as well as below-normal rainfall first occur at the southern foot of the TP between days −9 and −6, and then the signals are strengthened locally and expand eastward to southern China. They reach their peak about day −3 and sustain to day 0 over the region between 90° and 110°E for the band 32.5°–45°N (Fig. 10a). The below-normal OLR and anomalous ascending motion grow first over northwest of the TP and its neighbors between days −9 and −6. The anomalies become stronger when propagating eastward. They reach the peaks between days −3 and 0 to the east of the TP. After day 0, the anomalous signals in this band creep eastward and spread slowly to the northwest Pacific.
Longitude–time section of regressed subseasonal anomalies of OLR (thin blue contour; W m−2), rainfall (thick maroon contour; mm day−1), and 400-hPa vertical velocity ω (shaded; 0.01 hPa s−1) averaged over (a) 32.5°–45°N, (b) 22.5°–30°N, and (c) 5°–20°N. Negative (positive) lags on the vertical axis (days) mean OLR (rainfall and ω) anomalies leading (lagging) EI index. Lag = 0 corresponds to simultaneous regression. The thick maroon contours for rainfall are 0.2 (solid) and −0.2 mm day−1 (dashed), respectively.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
It is known that diabatic heating over the Asian monsoon region is mostly from latent heat released from convection and precipitation (Yanai et al. 1973; Yanai and Tomita 1998; Jin et al. 2013). Therefore, the anomalous convection/rainfall pattern over East Asia is inevitably linked with an anomalous pattern of diabatic heating. Figures 11a–c display the evolution of regressed anomalies of the vertically integrated heating source 
(left) As in Fig. 10, but for vertically integrated heating source 
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
In addition, Fig. 11 also shows that the anomalous signal with positive 
5. Diabatic heating feedback and its role in the eastward extension of the SAH
In this section, we first diagnose the evolutions of three terms contributing to the local PV change at 200 hPa, with an emphasis on the processes related with diabatic heating and rainfall. The three terms include the PV generation term due to diabatic heating and horizontal and vertical advection terms, presented in section 2. Then, we investigate the connection of the anomalous circulation at 850 hPa with the diabatic heating over the SCS and subtropical western Pacific regions. We intend to reveal the role of diabatic heating feedback in favor of the SAH’s eastward extension during days −6 to 0, especially days −3 to 0.
a. Diabatic heating and subseasonal PV changes at 200 hPa
The above analyses showed that the signals of anomalous diabatic heating and vertical motion over eastern Asia are robust during days −6 to 0 (Figs. 10 and 11). The sudden enhancement of the negative PV200 (positive Z200) anomaly to the east of TP (Fig. 6) and the dramatically eastward extension of the SAH (Fig. 4) also occurs at the same period. Thus, we focus on the period from days −6 to 0.
Figure 12 shows the regression fields of subseasonal anomalies of PVG [(∂PV′/∂t)PVG], the first term on the right-hand side of (7) [
Regressions of subseasonal anomalies of (left) PVG at 200 hPa [PV generation term due to vertical gradients of diabatic heating, (∂PV′/∂t)PVG; 10−2 PVU day−1], (center) the first term on the right-hand side of (7) at 200 hPa [
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
Figure 13 plots the regression fields of subseasonal anomalies of part-w [the vertical advection term, (∂PV′/∂t)part-w], superimposed on regressed subseasonal anomaly of vertical velocity (ω′) at 200 hPa. The calculation shows that the results of (∂PV′/∂t)part-w are contributed mostly by the second term on the right-hand side of (6) (figure not shown). The spatial pattern and propagation of (∂PV′/∂t)part-w are determined by ω′ at 200 hPa. The reason is that the term of 
As in Fig. 12, but for (left) part-w [vertical advection term, (∂PV′/∂t)part-w; 10−2 PVU day−1], and vertical velocity (ω′) at 200 hPa with blue solid (dashed) line of 0.0025 (−0.0025) hPa s−1, and (right) the sum of (∂PV′/∂t)PVG and (∂PV′/∂t)part-w (shaded), superimposed on regressed PV200 anomalies (thick maroon; PVU).
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
Figures 13d–f plot the sum of (∂PV′/∂t)PVG and (∂PV′/∂t)part-w, superimposed on the regressed PV200. It is seen that those two terms are negative in the region to the north of 30°N between 90° and 130°E during days −6 to 0. Thus, they are favorable to the increase of negative PV anomaly locally. To illustrate more clearly the above connection, we show in Fig. 14 longitude–time sections of regressed (∂PV′/∂t)PVG at 200 hPa, vertical gradient of diabatic heating (−∂Q1′/∂p) between 100 and 300 hPa, and (∂PV′/∂t)part-w at 200 hPa, all averaged over 32.5°–42.5°N. The corresponding regressed PV200 and Z200 are also plotted in Fig. 14. The weak center of negative PV200 (positive Z200) propagates eastward across the north of the TP with almost invariable value during days −12 to −6, as shown in Fig. 6. The nearly unchanged negative PV anomalies and small anomalies of (∂PV′/∂t)PVG and (∂PV′/∂t)part-w indicate that the horizontal advection is the dominant process associated with the eastward extension of the SAH during days −12 to −6. Other processes in this area are negligible.
Longitude–time sections (averaged over 32.5°–42.5°N) of regressed subseasonal anomalies of (a) PV200 (shaded; PVU) and Z200 (blue contour; gpm), (b) PVG at 200 hPa (shaded; 10−2 PVU day−1) and vertical gradient of diabatic heating (−∂Q1′/∂p) between 100 and 300 hPa (thick maroon; K day−1 hPa−1), and (c) part-w (vertical advection term, shaded; 10−2 PVU day−1). The black contours in (b) and (c) are regressed subseasonal anomalies of PV200 as in (a) with interval of 0.1 PVU, and dashed lines denoting negative values. The interval for thick maroon contour in (b) is 0.1 K day−1 hPa−1 with zero contour omitted.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
After day −6, the negative PV200 center is located to the east of TP with dramatically increased amplitude. The horizontal advection process is primary in the midlatitudes as a result of the westerly jet. However, the increased amplitude of negative PV anomaly suggests that other processes, like diabatic heating and rainfall, also have an effect on the PV variation to a certain extent. Figure 14 shows that the evolutions of negative (∂PV′/∂t)PVG and (∂PV′/∂t)part-w over the northern TP and northern China are similar to the negative PV200 center during days −6 to 0. The negative (∂PV′/∂t)PVG and (∂PV′/∂t)part-w during days −3 to 0 over the region of 90°–120°E largely overlap with the negative PV200 center. As mentioned earlier, the spatiotemporal features of (∂PV′/∂t)PVG and (∂PV′/∂t)part-w are nearly identical to those of anomalous diabatic heating and its vertical gradient. Moreover, the latter are closely connected with the anomalous rainfall and vertical motion. Therefore, the above results suggest that the positive diabatic heating (enhanced rainfall and ascending motion) anomalies over the northern TP and northern China regions are favorable to the increasing of the negative PV locally during days −3 to 0 through the PV generation term and vertical advection term. That positive diabatic heating signal originates from the eastward propagation of the signals of enhanced rainfall and ascending motion over the northwestern TP before day −3. Thus, the intensified negative PV200 (positive Z200) center to the east of TP appears a connection with the eastward propagation of enhanced rainfall and anomalous ascending motion stemming from the northwestern TP.
The negative PV200 (positive Z200) center to the east of the TP during days −3 to 0 also expands to southern China (Figs. 13d–f). To analyze the causation, Fig. 15 displays latitude–time sections of (∂PV′/∂t)PVG, (∂PV′/∂t)part-w, and the horizontal advection term [(∂PV′/∂t)part-h], averaged over 95°–125°E. Latitudinal distribution of climatological zonal and meridional winds at 200 hPa in summer is also plotted. Figure 15 shows that the negative PV anomaly over southern China is related with the process of PV horizontal advection. During the summertime, the climatological northerly flow is located to the east and southeast of TP, where the climatological westerly is relatively weaker than the jet stream region (Fig. 15d). The dipolar structure of negative (positive) (∂PV′/∂t)PVG and (∂PV′/∂t)part-w to the north (south) of 30°–34°N around days −3 to 0 (Figs. 15a,b) indicates large values of the meridional gradient of PV anomalies (−∂PV′/∂y). The north–south dipolar structure of PV anomaly with the climatological northerly flow is contributed to the southward advection of negative PV anomalies from the north lobe to southern China during days −3 to 0. After day 0, the dipolar pattern of diabatic heating and rainfall anomalies over eastern Asia collapses rapidly. The negative PV anomaly to the east of the TP decreases. The eastward extension of the SAH terminates.
As in Fig. 14, but for latitude–time sections (averaged over 95°–125°E) of (a) PVG (shaded; 10−2 PVU day−1) and −∂Q1′/∂p (thick maroon; K day−1 hPa−1), (b) part-w (shaded; 10−2 PVU day−1), and (c) part-h [horizontal advection term, (∂PV′/∂t)part-h, shaded; 10−2 PVU day−1]. (d) Latitudinal distribution of climatological summertime zonal (purple line) and meridional (yellow line) winds at 200 hPa averaged over 95°–125°E (m s−1). The zonal wind in (d) is multiplied by 0.25.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
b. Diabatic heating around the SCS and subtropical western Pacific regions
As noted in section 4, the robust diabatic heating anomaly is located around the SCS and subtropical western Pacific regions before day −6. It dramatically decreases on days −3 to 0 (Fig. 11c). To show the connection between the diabatic heating anomalies around these regions and the eastward extension of the SAH, we calculate lead correlation between the diabatic heating anomalies in the lower level (850–600 hPa averaged) and the EI with diabatic heating leading EI. Figure 16 plots the maximum and minimum lead correlation and the corresponding lead time that maximum and minimum correlations occur. It is seen that the regions with large maximum and minimum correlation display different features. The large maximum correlation region has two branches. One is located to the east of the TP corresponding to the anomalous diabatic heating shown in Fig. 11d. Another is around the SCS and western Pacific regions corresponding to the anomalous heating shown in Fig. 11f. The value of maximum correlation coefficient over the SCS and western Pacific regions is small but statistically significant. The large minimum correlation region is over southern China corresponding to anomalous cooling shown in Fig. 11e. On average, the lead time is from −5 to 0 days for the regions of southern China and to the east of the TP, which confirms again the close relationship between diabatic heating anomalies over eastern Asia and negative PV200 (high Z200) anomalies to the east of TP. The lead time from −10 to −6 days around the SCS and western Pacific regions is much earlier than that for the regions of southern China and to the east of the TP, although the correlation in this region is smaller than that in the other two regions.
(a) Maximum and (c) minimum lead correlation between subseasonal anomaly of diabatic heating in the lower level (850–600 hPa averaged) and EI (diabatic heating leads EI), and lead time (days) that (b) maximum and (d) minimum correlations occur. Shading areas denote regions of statistically significant at 95% confidence.
Citation: Journal of Climate 28, 17; 10.1175/JCLI-D-14-00682.1
As shown in section 3, the negative PV200 center to the east of the TP is not well established before day −6. Thus, we speculate that the anomalous diabatic heating around the SCS and western Pacific regions plays a role in the eastward extension of the SAH through influencing the regional low-level circulation before day −6. Figures 8f and 8g show that an anomalous cyclonic circulation at 850 hPa prevails from the SCS to the east of the Philippine Sea from days −9 to −6. It is located exactly in the northwest part of the positive diabatic heating anomalies (Fig. 16a, also seen in Fig. 11c), which is a local response to the positive heating (Gill 1980). The WPSH lingers over the western Pacific Ocean due to the above anomalous cyclone over the coastal waters. Therefore, the transportation channel for water vapor from the coastal waters to the southeastern Asia is cut off anomalously. Such a condition contributes to the negative anomalies of rainfall (diabatic heating) and thus positive PV anomaly over southern China during the SAH’s eastward extension period. The above analysis indicates that the anomalous diabatic heating around the SCS and western Pacific regions before day −6 is favorable to the eastward extension of the SAH.
6. Conclusions and discussion
a. Conclusions
The present study investigates the features of eastward extension of the SAH at 200 hPa on subseasonal time scales, and its connection to the diabatic heating and rainfall over eastern Asia. The midlatitude wave train over the Eurasian region related with the SAH’s eastward extension are examined by regressing PV and geopotential height fields at 200 hPa onto the definition of the SAH’s eastward extension index (EI). The role of diabatic heating feedback in the SAH’s eastward extension is revealed through performing PV diagnostic analysis.
The index EI, which is time series of normalized subseasonal anomaly of Z200 over the region 22.5°–32.5°N, 100°–120°E, has a periodicity of 10–30 days. It can reflect the variation over the east edge of the SAH. Based on the EI, a total of 40 (43) eastward (westward) events are identified for the summers of 1979–2012. The SAH’s eastward extension and then westward retreat on subseasonal time scale is visually presented by lagged composites of the 12520 line at 200 hPa for 40 eastward events. When the SAH is at its easternmost position (denoted as day 0), a strong anomalous high (low) in the Z200 (PV200) fields is sustained over eastern Asia. At the early stage (before day −6), the SAH stretches eastward tardily. At the later stage (from day −6 to day 0), the SAH moves eastward with big strides. At 500 hPa, the 5860-gpm line advances toward the inland accompanied by the eastward extension of the SAH.
The eastward extension of the SAH is characterized by the eastward propagation of a wave train across the Eurasian continent on subseasonal time scale. The wave train exhibits an equivalent barotropic vertical structure in the middle-to-high latitudes, while it has a distinct baroclinic character over subtropical eastern Asia. During days −9 to −6, the weak high over the western flank of TP, which is one of the wave train centers, propagates eastward rapidly along the subtropical westerly jet with a small change in the amplitude, and arrives at eastern China. This process leads to the slow extension of the SAH at its early stage. On day 0, the anomalous high originating from the western flank of TP becomes strongest over eastern China, while the anomalous low stemming from central Asia migrates to the Lake Baikal region. The pattern of wave-activity flux substantiates the connection of the SAH’s eastward extension and the eastward propagation of the wave train.
The eastward extension of the SAH has a strong relationship with the anomalies of diabatic heating and rainfall and vertical motion in three zonally elongated regions over East Asia: the SCS and subtropical western Pacific, the southern foot of TP and southern China, and the northern TP and northern China. The above-normal diabatic heating and rainfall and anomalous ascending motion first appear over the SCS and subtropical western Pacific regions on day −12 as a precursor. The largest correlation occurs at lag time of −10 to −6 days when the lower-level diabatic heating anomaly in that region leads the EI. The anomalous heating induces a lower-level anomalous cyclone over the coastal region, which is unfavorable to the water vapor transportation from the coastal waters to the southeastern Asian region. Such a situation contributes to the below-normal rainfall in southern China during the SAH’s eastward extension.
The anomalous cooling and below-normal rainfall and descending over the southern foot of the TP and southern China and anomalous heating and above-normal rainfall and ascending over the northern TP and northern China are formed on day −6. The above signals in the two regions construct a north–south dipolar structure with the boundary at 30°–32°N over eastern China. The dipolar pattern propagates eastward and matures to the east of TP at days −3 to 0. Around day 0, to the east of TP, an anomalous northerly flow induced by the anomalous low over the Lake Baikal converges with an anomalous southerly flow originated from the SCS and the subtropical western Pacific Ocean, contributing to the above-normal rainfall in northern China.
According to the results of PV diagnosis at 200 hPa, the subseasonal anomaly of local PV change is mainly controlled by three terms: subseasonal anomaly of the PV generation term and vertical and horizontal advection terms. The horizontal advection process is primary in the midlatitudes as a result of the westerly jet. The subseasonal anomaly of the PV generation term is identical to that of diabatic heating in the structure and propagation aspects. The subseasonal anomaly of vertical advection is decided by the anomalous vertical motion, which is analogous to the rainfall and diabatic heating.
We suggest three major diabatic heating (rainfall)-related processes connecting to the eastward extension of the SAH, during days −6 to −3, especially days −3 to 0. 1) The above-normal diabatic heating over northern TP and northern China produces a negative PV anomaly through the diabatic heating feedback process. 2) The vertical advection term induces a negative PV anomaly over northern TP and northern China, through the process of the anomalous ascending velocity carrying low PV value in the lower level upward. Our results show that the above positive diabatic heating and rainfall anomalies originate from the eastward propagation of the signals of enhanced rainfall and ascending motion over the northwestern TP before day −3, which is part of the features linked with the wave train. 3) The horizontal advection process is conducive to the negative PV anomalies over southern China through southward advection of the negative PV anomaly. The negative anomalies in diabatic heating and descent over southern China can generate positive PV anomalies locally, favoring the formation of a meridional gradient of PV anomalies over eastern China.
b. Discussion
Asian monsoon circulations demonstrate multiform variability. A subseasonal signal is an attractive time scales because it is associated with active and break cycles of the Indian monsoon, as well as the advanced northward progression of the rainfall belt over East Asia (Annamalai and Slingo 2001; Mao et al. 2010; Jia and Yang 2013; Moon et al. 2013; Lu et al. 2014). The subseasonal signal shows its preferred direction of propagation. For example, the northward propagation is one of the most striking features of the boreal summer intraseasonal oscillation over the Indian Ocean and western Pacific monsoon regions (Jiang et al. 2004; Drbohlav and Wang 2005). Besides the northward movement, the Asian monsoon circulations, like the SAH and the WPSH, are characterized by a subseasonal zonal shift (Ren et al. 2013; Yang et al. 2014). However, our understanding of the mechanism responsible for the zonal oscillation is still limited. In this paper we investigate the cause for the SAH eastward extension based on the joint roles of the midlatitude wave train across the Eurasian landmass and the diabatic heating feedback over eastern Asia. During the SAH’s eastward extension, the diabatic heating interacts with the atmospheric circulation. We investigate only the role of diabatic heating feedback in the SAH’s eastward extension. How the SAH’s zonal shift induces the diabatic heating and rainfall anomalies is not intensively examined in this paper. Another puzzling question is why the WPSH moves inland, when the SAH stretches eastward. In the present investigation, the wave train and diabatic heating feedback collaborate and contribute to the eastward extension of the SAH. These results imply that there is linkage between East Asian monsoon convection and midlatitude wave train on subseasonal time scale, and this needs further analysis (Moon et al. 2013). It should be pointed out that the present results are based on statistical analysis. Can the wave train propagation or the East Asian monsoonal rainfall processes cause the eastward extension of the SAH solely for a certain extreme event? To answer these questions, more comprehensive investigations need to be carried out in future study.
Acknowledgments
This work was jointly supported by the 973 Program (Grant 2012CB417203) and the National Natural Science Foundation of China under Grants 41330420 and 41275068. We thank three anonymous reviewers for their insightful comments, which led to great improvements of this manuscript. The ERA-Interim daily and monthly data were downloaded from http://apps.ecmwf.int/datasets/.
REFERENCES
Annamalai, H., and J. M. Slingo, 2001: Active/break cycles: Diagnosis of the intraseasonal variability of the Asian summer monsoon. Climate Dyn., 18, 85–102, doi:10.1007/s003820100161.
Bueh, C., N. Shi, L. Ji, J. Wei, and S. Tao, 2008: Features of the EAP events on the medium-range evolution process and the mid- and high-latitude Rossby wave activities during the Meiyu period. Chin. Sci. Bull., 53, 610–623, doi:10.1007/s11434-008-0005-2.
Chen, J., and S. Bordoni, 2014: Orographic effects of the Tibetan Plateau on the East Asian summer monsoon: An energetic perspective. J. Climate, 27, 3052–3072, doi:10.1175/JCLI-D-13-00479.1.
Chen, Y., and P. Zhai, 2014: Precursor circulation features for persistent extreme precipitation in central-eastern China. Wea. Forecasting, 29, 226–240, doi:10.1175/WAF-D-13-00065.1.
Dee, D. P., and Coauthors, 2011: The ERA-Interim Reanalysis: Configuration and performance of the data assimilation system. Quart. J. Roy. Meteor. Soc., 137, 553–597, doi:10.1002/qj.828.
Ding, Q., and B. Wang, 2007: Intraseasonal teleconnection between the summer Eurasian wave train and the Indian monsoon. J. Climate, 20, 3751–3767, doi:10.1175/JCLI4221.1.
Drbohlav, H. L., and B. Wang, 2005: Mechanism of the northward-propagating intraseasonal oscillation: Insights from a zonally symmetric model. J. Climate, 18, 952–972, doi:10.1175/JCLI3306.1.
Duan, A., and G. Wu, 2005: Role of the Tibetan Plateau thermal forcing in the summer climate patterns over subtropical Asia. Climate Dyn., 24, 793–807, doi:10.1007/s00382-004-0488-8.
Duan, A., G. Wu, and X. Liang, 2008: Influence of the Tibetan Plateau on the summer climate patterns over Asia in the IAP/LASG SAMIL model. Adv. Atmos. Sci., 25, 518–528, doi:10.1007/s00376-008-0518-2.
Duan, A., J. Hu, and Z. Xiao, 2013: The Tibetan Plateau summer monsoon in the CMIP5 simulations. J. Climate, 26, 7747–7766, doi:10.1175/JCLI-D-12-00685.1.
Enomoto, T., B. J. Hoskins, and Y. Matsuda, 2003: The formation mechanism of the Bonin high in August. Quart. J. Roy. Meteor. Soc., 129, 157–178, doi:10.1256/qj.01.211.
Fujinami, H., and T. Yasunari, 2004: Submonthly variability of convection and circulation over and around the Tibetan Plateau during the boreal summer. J. Meteor. Soc. Japan, 82, 1545–1564, doi:10.2151/jmsj.82.1545.
Garny, H., and W. J. Randel, 2013: Dynamic variability of the Asian monsoon anticyclone observed in potential vorticity and correlations with tracer distributions. J. Geophys. Res., 118, 13 421–13 433, doi:10.1002/2013JD020908.
Gill, A. E., 1980: Some simple solutions for heat-induced tropical circulation. Quart. J. Roy. Meteor. Soc., 106, 447–462, doi:10.1002/qj.49710644905.
Hoerling, M. P., 1992: Diabatic sources of potential vorticity in the general circulation. J. Atmos. Sci., 49, 2282–2292, doi:10.1175/1520-0469(1992)049<2282:DSOPVI>2.0.CO;2.
Hoskins, B. J., and T. Ambrizzi, 1993: Rossby wave propagation on a realistic longitudinally varying flow. J. Atmos. Sci., 50, 1661–1671, doi:10.1175/1520-0469(1993)050<1661:RWPOAR>2.0.CO;2.
Hoskins, B. J., and M. J. Rodwell, 1995: A model of the Asian summer monsoon. Part I: The global scale. J. Atmos. Sci., 52, 1329–1340, doi:10.1175/1520-0469(1995)052<1329:AMOTAS>2.0.CO;2.
Hoskins, B. J., M. E. McIntyre, and A. W. Robertson, 1985: On the use and significance of isentropic potential vorticity maps. Quart. J. Roy. Meteor. Soc., 111, 877–946, doi:10.1002/qj.49711147002.
Hsu, C., and R. Plumb, 2000: Nonaxisymmetric thermally driven circulations and upper-tropospheric monsoon dynamics. J. Atmos. Sci., 57, 1255–1276, doi:10.1175/1520-0469(2000)057<1255:NTDCAU>2.0.CO;2.
Huang, G., X. Qu, and K. Hu, 2011: The impact of the tropical Indian Ocean on the South Asian high in boreal summer. Adv. Atmos. Sci., 28, 421–432, doi:10.1007/s00376-010-9224-y.
Jia, X., and S. Yang, 2013: Impact of the quasi-biweekly oscillation over the western North Pacific on East Asian subtropical monsoon during early summer. J. Geophys. Res. Atmos., 118, 4421–4434, doi:10.1002/jgrd.50422.
Jiang, X., and N.-C. Lau, 2008: Intraseasonal teleconnection between North American and western North Pacific monsoons with 20-day time scale. J. Climate, 21, 2664–2679, doi:10.1175/2007JCLI2024.1.
Jiang, X., T. Li, and B. Wang, 2004: Structures and mechanisms of the northward propagating boreal summer intraseasonal oscillation. J. Climate, 17, 1022–1039, doi:10.1175/1520-0442(2004)017<1022:SAMOTN>2.0.CO;2.
Jiang, X. W., Y. Li, S. Yang, and R. Wu, 2011: Interannual and interdecadal variations of the South Asian and western Pacific subtropical highs and their relationships with Asian-Pacific summer climate. Meteor. Atmos. Phys., 113, 171–180, doi:10.1007/s00703-011-0146-8.
Jin, F., and B. J. Hoskins, 1995: The direct response to tropical heating in a baroclinic atmosphere. J. Atmos. Sci., 52, 307–319, doi:10.1175/1520-0469(1995)052<0307:TDRTTH>2.0.CO;2.
Jin, Q., X.-Q. Yang, X.-G. Sun, and J.-B. Fang, 2013: East Asian summer monsoon circulation structure controlled by feedback of condensational heating. Climate Dyn., 41, 1885–1897, doi:10.1007/s00382-012-1620-9.
Kim, K.-Y., J.-W. Roh, D.-K. Lee, and J.-G. Jhun, 2010: Physical mechanisms of the seasonal, subseasonal, and high-frequency variability in the seasonal cycle of summer precipitation in Korea. J. Geophys. Res., 115, D14110, doi:10.1029/2009JD013561.
Krishnamurthy, V., and J. Shukla, 2000: Intraseasonal and interannual variability of rainfall over India. J. Climate, 13, 4366–4377, doi:10.1175/1520-0442(2000)013<0001:IAIVOR>2.0.CO;2.
Krishnamurti, T. N., 1971: Tropical east–west circulations during the northern summer. J. Atmos. Sci., 28, 1342–1347, doi:10.1175/1520-0469(1971)028<1342:TEWCDT>2.0.CO;2.
Krishnan, R., V. Kumar, M. Sugi, and J. Yoshimura, 2009: Internal feedbacks from monsoon–midlatitude interactions during droughts in the Indian summer monsoon. J. Atmos. Sci., 66, 553–578, doi:10.1175/2008JAS2723.1.
Liebmann, B., and C. A. Smith, 1996: Description of a complete (interpolated) outgoing longwave radiation dataset. Bull. Amer. Meteor. Soc., 77, 1275–1277.
Lin, H., 2009: Global extratropical response to diabatic heating variability of the Asian summer monsoon. J. Atmos. Sci., 66, 2697–2713, doi:10.1175/2009JAS3008.1.
Ling, J., and C. Zhang, 2013: Diabatic heating profiles in recent global reanalyses. J. Climate, 26, 3307–3325, doi:10.1175/JCLI-D-12-00384.1.
Liu, B., J. He, and L. Wang, 2012: On a possible mechanism for southern Asian convection influencing the South Asian high establishment during winter to summer transition. J. Trop. Meteor., 18 (4), 473–484.
Liu, B., G. Wu, J. Mao, and J. He, 2013: Genesis of the South Asian high and its impact on the Asian summer monsoon onset. J. Climate, 26, 2976–2991, doi:10.1175/JCLI-D-12-00286.1.
Liu, F., and H. Lin, 1991: Relationship between the atmospheric low-frequency oscillation and the variability of the South Asian high (in Chinese). Plateau Meteor., 10 (1), 61–69.
Liu, H., S. Zhao, C. Zhao, and Z. Lu, 2006: Weather abnormal and evolutions of western Pacific subtropical high and South Asian high in summer 2003 (in Chinese). Plateau Meteor., 25 (2), 169–178.
Liu, Y., and G. Wu, 2004: Progress in the study on the formation of the summertime subtropical anticyclone. Adv. Atmos. Sci., 21, 322–342, doi:10.1007/BF02915562.
Liu, Y., G. Wu, and R. Ren, 2004: Relationship between the subtropical anticyclone and diabatic heating. J. Climate, 17, 682–698, doi:10.1175/1520-0442(2004)017<0682:RBTSAA>2.0.CO;2.
Liu, Y., B. Hoskins, and M. Blackburn, 2007: Impact of Tibetan orography and heating on the summer flow over Asia. J. Meteor. Soc. Japan, 85B, 1–19, doi:10.2151/jmsj.85B.1.
Lu, R., J. Oh, and B. Kim, 2002: A teleconnection pattern in upper-level meridional wind over the North African and Eurasian continent in summer. Tellus, 54A, 44–55, doi:10.1034/j.1600-0870.2002.00248.x.
Lu, R., H. L. Dong, Q. Su, and H. Ding, 2014: The 30–60-day intraseasonal oscillations over the subtropical western North Pacific during the summer of 1998. Adv. Atmos. Sci., 31, 1–7, doi:10.1007/s00376-013-3019-x.
Luo, S., Z. Qian, and Q. Wang, 1982: The climatic and synoptical study about the relationship between the Qinghai-Xizang high pressure on the 100 mb surface and the flood and drought in East China in summer (in Chinese). Plateau Meteor., 1, 1–10.
Mao, J., Z. Sun, and G. Wu, 2010: 20–50-day oscillation of summer Yangtze rainfall in response to intraseasonal variations in the subtropical high over the western North Pacific and South China Sea. Climate Dyn., 34, 747–761, doi:10.1007/s00382-009-0628-2.
Mason, R. B., and C. E. Anderson, 1963: The development and decay of the 100-mb summertime anticyclone over southern Asia. Mon. Wea. Rev., 91, 3–12, doi:10.1175/1520-0493(1963)091<0003:TDADOT>2.3.CO;2.
Moon, J.-Y., B. Wang, K.-J. Ha, and J.-Y. Lee, 2013: Teleconnections associated with Northern Hemisphere summer monsoon intraseasonal oscillation. Climate Dyn., 40, 2761–2774, doi:10.1007/s00382-012-1394-0.
Nakamura, H., and T. Fukamachi, 2004: Evolution and dynamics of summertime blocking over the Far East and the associated surface Okhotsk high. Quart. J. Roy. Meteor. Soc., 130, 1213–1233, doi:10.1256/qj.03.101.
Peng, L., Z. Sun, H. Chen, and D. Ni, 2010: The persistent anomaly of the South Asia high and its association with ENSO events (in Chinese). Acta Meteor. Sin., 68 (6), 855–864.
Popovic, J. M., and R. A. Plumb, 2001: Eddy shedding from the upper-tropospheric Asian monsoon anticyclone. J. Atmos. Sci., 58, 93–104, doi:10.1175/1520-0469(2001)058<0093:ESFTUT>2.0.CO;2.
Qu, X., and G. Huang, 2012: An enhanced influenced of tropical Indian Ocean on the South Asian high after the late 1970s. J. Climate, 25, 6930–6941, doi:10.1175/JCLI-D-11-00696.1.
Raman, C. R. V., and Y. P. Rao, 1981: Blocking highs over Asia and monsoon droughts over India. Nature, 289, 271–273, doi:10.1038/289271a0.
Randel, W. J., and M. Park, 2006: Deep convective influence on the Asian summer monsoon anticyclone and associated tracer variability observed with Atmospheric Infrared Sounder (AIRS). J. Geophys. Res., 111, D12314, doi:10.1029/2005JD006490.
Ren, R., Y. Liu, and G. Wu, 2007: Impact of South Asian high on the short term variation of the subtropical anticyclone over western Pacific in July 1998 (in Chinese). Acta Meteor. Sin., 65 (2), 183–197.
Ren, X., X.-Q. Yang, and X. Sun, 2013: Zonal oscillation of western Pacific subtropical high and subseasonal SST variations during Yangtze persistent heavy rainfall events. J. Climate, 26, 8929–8946, doi:10.1175/JCLI-D-12-00861.1.
Seo, K.-H., and E.-J. Song, 2012: Initiation of boreal summer intraseasonal oscillation: Dynamic contribution by potential vorticity. Mon. Wea. Rev., 140, 1748–1760, doi:10.1175/MWR-D-11-00105.1.
Takaya, K., and H. Nakamura, 2001: A formulation of a phase-independent wave-activity flux for stationary and migratory quasigeostrophic eddies on a zonally varying basic flow. J. Atmos. Sci., 58, 608–627, doi:10.1175/1520-0469(2001)058<0608:AFOAPI>2.0.CO;2.
Tao, S.-Y., and F. Zhu, 1964: The 100-mb flow patterns in southern Asia in summer and its relation to the advance and retreat of the West-Pacific subtropical anticyclone over the Far East (in Chinese). Acta Meteor. Sin., 34, 385–396.
Tao, S.-Y., and L.-X. Chen, 1987: A review of recent research on the East Asian summer monsoon in China. Monsoon Meteorology, C.-P. Chang and T. N. Krishnamurti, Eds., Oxford University Press, 60–92.
Tao, S.-Y., and J. Wei, 2006: The westward, northward advance of the subtropical high over the west Pacific in summer (in Chinese). J. Appl. Meteor. Sci, 17 (5), 513–525.
Tory, K. J., J. D. Kepert, J. A. Sippel, and C. M. Nguyen, 2012: On the use of potential vorticity tendency equations for diagnosing atmospheric dynamics in numerical models. J. Atmos. Sci., 69, 942–960, doi:10.1175/JAS-D-10-05005.1.
Wang, Z., A. Duan, and G. Wu, 2014: Time-lagged impact of spring sensible heat over the Tibetan Plateau on the summer rainfall anomaly in East China: Case studies using the WRF model. Climate Dyn., 42, 2885–2898, doi:10.1007/s00382-013-1800-2.
Watanabe, T., and K. Yamazaki, 2012: Influence of the anticyclonic anomaly in the subtropical jet over the western Tibetan Plateau on the intraseasonal variability of the summer Asian monsoon in early summer. J. Climate, 25, 1291–1303, doi:10.1175/JCLI-D-11-00036.1.
Watanabe, T., and K. Yamazaki, 2014: The upper-level circulation anomaly over central Asia and its relationship to the Asian monsoon and mid-latitude wave train in early summer. Climate Dyn., 42, 2477–2489, doi:10.1007/s00382-013-1888-4.
Wei, W., R. Zhang, M. Wen, X. Rong, and T. Li, 2014: Impact of Indian summer monsoon on the South Asian high and its influence on summer rainfall over China. Climate Dyn., 43, 1257–1269, doi:10.1007/s00382-013-1938-y.
Wu, G., Y. Liu, B. He, Q. Bao, A. Duan, and F. Jin, 2012: Thermal controls on the Asian summer monsoon. Sci. Rep., 2, 404, doi:10.1038/srep00404.
Wu, L., and B. Wang, 2000: A potential vorticity tendency diagnostic approach for tropical cyclone motion. Mon. Wea. Rev., 128, 1899–1911, doi:10.1175/1520-0493(2000)128<1899:APVTDA>2.0.CO;2.
Xie, P., A. Yatagai, M. Chen, T. Hayasaka, Y. Fukushima, C. Liu, and S. Yang, 2007: A gauge-based analysis of daily precipitation over East Asia. J. Hydrometeor., 8, 607–627, doi:10.1175/JHM583.1.
Yanai, M., and T. Tomita, 1998: Seasonal and interannual variability of atmospheric heat sources and moisture sinks as determined from NCEP–NCAR reanalysis. J. Climate, 11, 463–482, doi:10.1175/1520-0442(1998)011<0463:SAIVOA>2.0.CO;2.
Yanai, M., S. Esbensen, and J.-H. Chu, 1973: Determination of bulk properties of tropical cloud clusters from large-scale heat and moisture budgets. J. Atmos. Sci., 30, 611–627, doi:10.1175/1520-0469(1973)030<0611:DOBPOT>2.0.CO;2.
Yang, J., and Coauthors, 2014: Distinct quasi-biweekly features of the subtropical East Asian monsoon during early and late summer. Climate Dyn., 42, 1469–1486, doi:10.1007/s00382-013-1728-6.
Yang, J. L., Q. Liu, S.-P. Xie, Z. Liu, and L. Wu, 2007: Impact of the Indian Ocean SST basin mode on the Asian summer monsoon. Geophys. Res. Lett., 34, L02708, doi:10.1029/2006GL028571.
Yasui, S., and M. Watanabe, 2010: Forcing processes of the summertime circumglobal teleconnection pattern in a dry AGCM. J. Climate, 23, 2093–2114, doi:10.1175/2009JCLI3323.1.
Yatagai, A., O. Arakawa, K. Kamiguchi, H. Kawamoto, M. I. Nodzu, and A. Hamada, 2009: A 44-year daily precipitation dataset for Asia based on dense network of rain gauges. SOLA, 5, 137–140, doi:10.2151/sola.2009-035.
Yatagai, A., K. Kamiguchi, O. Arakawa, A. Hamada, N. Yasutomi, and A. Kitoh, 2012: APHRODITE: Constructing a long-term daily gridded precipitation dataset for Asia based on a dense network of rain gauges. Bull. Amer. Meteor. Soc., 93, 1401–1415, doi:10.1175/BAMS-D-11-00122.1.
Yuan, F., W. Chen, and W. Zhou, 2012: Analysis of the role played by circulation in the persistent precipitation over South China in June 2010. Adv. Atmos. Sci., 29, 769–781, doi:10.1007/s00376-012-2018-7.
Zarrin, A., H. Ghaemi, M. Azadi, and M. Farajzadeh, 2010: The spatial pattern of summertime subtropical anticyclones over Asia and Africa: A climatological review. Int. J. Climatol., 30, 159–173, doi:10.1002/joc.1879.
Zhang, C., and J. Ling, 2012: Potential vorticity of the Madden–Julian oscillation. J. Atmos. Sci., 69, 65–78, doi:10.1175/JAS-D-11-081.1.
Zhang, P., S. Yang, and V. E. Kousky, 2005: South Asian high and Asian–Pacific–American climate teleconnection. Adv. Atmos. Sci., 22, 915–923, doi:10.1007/BF02918690.
Zhang, Q., Y. Qian, and X. Zhang, 2000: Interannual and interdecadal variations of the South Asian high (in Chinese). Chin. J. Atmos. Sci., 24 (1), 67–78.
Zhang, Q., G. Wu, and Y. Qian, 2002: The bimodality of the 100 hPa South Asia high and its relationship to the climate anomaly over East Asia in summer. J. Meteor. Soc. Japan, 80, 733–744, doi:10.2151/jmsj.80.733.
Zhang, Y., X. Kuang, W. Guo, and T. Zhou, 2006: Seasonal evolution of the upper-tropospheric westerly jet core over East Asia. Geophys. Res. Lett., 33, L11708, doi:10.1029/2006GL026377.
Zhang, Y.-N., G. Wu, Y. Liu, and G. Yue, 2014: The effects of asymmetric potential vorticity forcing on the instability of South Asia high and Indian summer monsoon onset. Sci. China Earth Sci.,57, 337–350, doi:10.1007/s11430-013-4664-8.
Zhao, P., X. Zhang, Y. Li, and J. Chen, 2009: Remotely modulated tropical–North Pacific ocean–atmosphere interactions by the South Asian high. Atmos. Res., 94, 45–60, doi:10.1016/j.atmosres.2009.01.018.
Zhou, T., and R. Yu, 2005: Atmospheric water vapor transport associated with typical anomalous summer rainfall patterns in China. J. Geophys. Res.,110, D08104, doi:10.1029/2004JD005413.
Zhu, F., L. Lu, X. Chen, and W. Zhao, 1980: The South Asian High (in Chinese). Science Press, 95 pp.
