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  • View in gallery
    Fig. 1.

    (a) The East–Southeast Asian summer monsoon life cycle [active phase (1/5–15/6), break phase (16/6–15/7), and revival phase (16/7–15/9)], depicted with the climatological 15-day mean rainfall for Taiwan from April to October [adopted from Chen et al. (2004)], and (b) the rainfall for Taiwan contributed by different weather systems [including diurnal variation (localized convection), typhoons, rainstorms, and fronts] for different phases of the summer monsoon life cycle, measured with Wang and Chen’s (2008) approach and averaged only over the southwestern region (21.7°–24.3°N, 120°–120.8°E) of Taiwan.

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    Fig. 2.

    A typical rainstorm case identified at 1200 UTC 18 May 2008 is shown: (a) 700-hPa streamline chart superimposed with vorticity, (b) surface streamline chart superimposed with precipitation of TRMM, (c) JMA surface analysis, and corresponding (d) satellite image of Multifunctional Transport Satellite (MTSAT) and (e) surface wind vectors of the Quick Scatterometer (QuikSCAT). (top right) Scales of vorticity and precipitation are shown in (a) and (b), respectively. The genesis location of this rainstorm is pointed by an arrow attached to symbol R1 (the first rainstorm identified in the mei-yu season of 2008). The arrow with R1 in (e) is pointing to the location of this rainstorm at 1100 UTC 19 May 2008.

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    Fig. 3.

    (a) Trajectories of rainstorms are depicted by lines with color corresponding to genesis locations. Genesis locations of rainstorms for Indochina (the South China Sea) during May–June over the 1979–2008 period are marked by red (blue) dots, (b) rainfall contributed by rainstorms (PRS), (c) distribution of the climatological rainfall (PT) ensembled with TRMM, CMORPH, and Global Precipitation Climate Project (GPCP) for May–June, and (d) ratio PRS/PT. The contour interval of (b),(c),(d) is 1 mm day−1, 1 mm day−1, and 5%, respectively.

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    Fig. 4.

    (a) Occurrence frequency of rainstorms NRS over the region of PRS ≥ 6 mm day−1 during May–June for the 1979–2008 period, superimposed with the ΔSST (Niño-3.4) index averaged over May–June and precipitation histograms of (b) rainstorm PRS, (c) total rainfall PT, and (d) the ratio, PRS/PT, over the region of PRS ≥ 6 mm day−1. Averaged values of all variables are also presented in each panel. The thermal condition of SST over the NOAA Niño-3.4 area (5°S–5°N, 170°–120°W) is determined by the following criteria: warm ΔSST (Niño-3.4) ≥ 0.5°C and cold ΔSST (Niño-3.4) ≤ −0.5°C. The warm and cold thermal conditions are colored red and blue, respectively, on histograms of all variables shown in this figure.

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    Fig. 5.

    As in Fig. 4, but for (left) Hong Kong and (right) Taiwan.

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    Fig. 6.

    (a) Genesis locations (red/blue dots) and track (red/blue lines) during warm May–June for rainstorms and (b) PRS(warm) produced by rainstorms superimposed with average daily locations of rainstorm centers during warm May–June. (c),(d) As in (a),(b), respectively, but for cold May–June, histograms of average daily PRS for (e) all identified rainstorms, (f) warm and (g) cold May–June, and (h),(i),(j) as in (e),(f),(g), but for convergence of water vapor flux. Note that the average tracks in (b),(c) are portrayed by a thick red line connecting the average locations (red dots) of rainstorm centers over the composite life cycle of these rainstorms, which is depicted by the day number attached to those average locations. The center location of rainstorms included in the averaging process is covered by red-dashed ellipse and its peripheral area. The axes of these ellipses are one standard deviation of rainstorm distances from the average center parallel and perpendicular to the averaged track.

  • View in gallery
    Fig. 7.

    Streamline charts at 700 hPa superimposed with (a) isotach |V|, and (b) variance of U(700 hPa) during May–June for the 1979–2008 period, (c) time series of u(700 hPa) averaged over the area of (15°–25°N, 100°–115°E) (solid line with dots) surrounding the maximum of u(700 hPa) variance at (20°N, 104°E) and ΔSST (Niño-3.4) (dash line with open circles), (d) the height–time cross section of u averaged over the area of (15°–25°N, 100°–115°E), and (e) the latitude–time cross section of u(700 hPa) over the longitudinal zone of (100°–115°E). (top right) Scales of |V| and Var(U) are shown in (a),(b), respectively. Measured in terms of ΔSST (Niño-3.4), the warm ENSO events [ΔSST (Niño-3.4) ≥ 0.5°C] are marked as W in (d),(e). The contour interval of (d),(e) is 0.5 m s−1, and (bottom right) the scale is shown in (d). The monsoon trough in (a),(b) is indicated by a thick, dashed blue line.

  • View in gallery
    Fig. 8.

    (a) Eddy of streamfunction at 700 hPa, ψE(700 hPa), superimposed with isotachs (red) averaged over May–June for the 1979–2008 period. Based on the index of ΔSST (Niño-3.4, MJ) shown in Fig. 7c, composite ψE(7000 hPa) anomalies [ΔψE(700 hPa)] superimposed with ΔSST anomalies for (b) warm and (c) cold events determined by ΔSST (Niño-3.4) anomalies averaged over May–June. Contour intervals of ψE in (a) and ΔψE in (b),(c) are 106 m2 s−1 and 3 × 105 m s−1, respectively. (bottom right) Scales of |V|(700 hPa) and ΔSST are shown in (a) and (b),(c), respectively. Arrows are added on ψE(700 hPa) and ΔψE(700 hPa) to indicate the flow directions.

  • View in gallery
    Fig. 9.

    Latitude–height cross sections of the meridional gradient of potential vorticity added with the meridional gradient of 925-hPa potential temperature (925 hPa) at (left) 104°E and (right) 114°E. (a),(d) Strong monsoon westerlies, (b),(e) weak monsoon westerlies, and (c),(f) the difference between strong and weak monsoon westerlies are shown. The contour interval of and is 10−6 m2 s−1 K kg−1.

  • View in gallery
    Fig. 10.

    As in Fig. 8, but for (a) , (b) (warm), and (c) (cold). Contour intervals of and are 107 kg s−1 and 4 × 106 kg s−1, respectively.

  • View in gallery
    Fig. 11.

    Histograms of (a) convergence of total water vapor flux [] averaged over the region of PRS ≥ 6 mm day−1, superimposed with PT histograms, and (b) convergence of water vapor flux [] associated with PRS averaged over the region of PRS ≥ 6 mm day−1, superimposed with PRS histograms, too. The seasons with ΔSST (Niño-3.4) ≥ 0.5°C (≤−0.5°C) are colored by pink (light blue), while the corresponding PT and PRS histograms are colored by red and dark blue, respectively. Histograms of and during normal years [i.e. 0.5°C ≥ ΔSST(Niño-3.4) ≥ −0.5°C] are not colored, but those of corresponding PT and PRS are shaded.

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Interannual Variation of the Late Spring–Early Summer Monsoon Rainfall in the Northern Part of the South China Sea

Tsing-Chang ChenDepartment of Geological and Atmospheric Sciences, Iowa State University, Ames, Iowa

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Wan-Ru HuangDepartment of Geological and Atmospheric Sciences, Iowa State University, Ames, Iowa

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Ming-Cheng YenDepartment of Atmospheric Sciences, National Central University, Chung-Li, Taiwan

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Abstract

Major rainfall (≥60%) in the northern part of the South China Sea (between North Vietnam and Taiwan) during May–June (the mei-yu season—the first phase of the Southeast–East Asian monsoon) is produced by rainstorms originating over the northern Vietnam–southwestern China region and the northern part of the South China Sea. As observed in this study, the occurrence frequency of rainstorms and rainfall contribution by these rainstorms undergoes a distinct interannual variation, in-phase with those of monsoon westerlies in northern Indochina and sea surface temperature (SST) anomalies over the NOAA Niño-3.4 region ΔSST (Niño-3.4). This in-phase relationship between monsoon westerlies and the ΔSST (Niño-3.4) anomalies is a result of the filling (deepening) of the subtropical Asian continental thermal low in response to the ΔSST (Niño-3.4) warm (cold) anomalies. Accompanied with this response is a slight southward (northward) shift of the North Pacific convergence zone (NPCZ), which extends from southern China to the North Pacific east of Japan. Thus, a favorable environment that meets the Charney–Stern instability criterion in initiating rainstorm genesis is enhanced (suppressed) by the intensification (weakening) of the monsoon shear flow formed by the midtropospheric northwesterly flow around the northeast periphery of the Tibetan Plateau and the monsoon westerlies. The meridional shift of the NPCZ established an elongated anomalous convergence (divergence) zone of water vapor flux along rainstorm tracks to increase (reduce) the rain-producing efficiency of rainstorms. Consequently, this interannual rainfall variation between northern Vietnam and Taiwan is primarily caused by rainstorm genesis and rain-producing efficiency.

Current affiliation: Research Fellow, Guy Carpenter Asia-Pacific Climate Impact Centre, School of Energy and Environment, City University of Hong Kong, Hong Kong, China.

Corresponding author address: Tsing-Chang (Mike) Chen, Atmospheric Science Program, Department of Geological and Atmospheric Sciences, 3010 Agronomy Hall, Iowa State University, Ames, IA 50011. E-mail: tmchen@iastate.edu

Abstract

Major rainfall (≥60%) in the northern part of the South China Sea (between North Vietnam and Taiwan) during May–June (the mei-yu season—the first phase of the Southeast–East Asian monsoon) is produced by rainstorms originating over the northern Vietnam–southwestern China region and the northern part of the South China Sea. As observed in this study, the occurrence frequency of rainstorms and rainfall contribution by these rainstorms undergoes a distinct interannual variation, in-phase with those of monsoon westerlies in northern Indochina and sea surface temperature (SST) anomalies over the NOAA Niño-3.4 region ΔSST (Niño-3.4). This in-phase relationship between monsoon westerlies and the ΔSST (Niño-3.4) anomalies is a result of the filling (deepening) of the subtropical Asian continental thermal low in response to the ΔSST (Niño-3.4) warm (cold) anomalies. Accompanied with this response is a slight southward (northward) shift of the North Pacific convergence zone (NPCZ), which extends from southern China to the North Pacific east of Japan. Thus, a favorable environment that meets the Charney–Stern instability criterion in initiating rainstorm genesis is enhanced (suppressed) by the intensification (weakening) of the monsoon shear flow formed by the midtropospheric northwesterly flow around the northeast periphery of the Tibetan Plateau and the monsoon westerlies. The meridional shift of the NPCZ established an elongated anomalous convergence (divergence) zone of water vapor flux along rainstorm tracks to increase (reduce) the rain-producing efficiency of rainstorms. Consequently, this interannual rainfall variation between northern Vietnam and Taiwan is primarily caused by rainstorm genesis and rain-producing efficiency.

Current affiliation: Research Fellow, Guy Carpenter Asia-Pacific Climate Impact Centre, School of Energy and Environment, City University of Hong Kong, Hong Kong, China.

Corresponding author address: Tsing-Chang (Mike) Chen, Atmospheric Science Program, Department of Geological and Atmospheric Sciences, 3010 Agronomy Hall, Iowa State University, Ames, IA 50011. E-mail: tmchen@iastate.edu

1. Introduction

The summer monsoon rainfall in East–Southeast Asia is characterized by the mei-yu rainband around the northwestern rim and the monsoon trough rainband (stretching from northern Indochina, across the northern part of the South China Sea, to the Philippine Sea) around the southwestern rim of the western Pacific subtropic high (e.g., Chen and Murakami 1988). Analyzing station observations, Ramage (1952) suggested the monsoon life cycle in the northern part of Southeast Asia and the southern part of East Asia is established by a transition from the early summer mei-yu regime into the late summer tropical cyclone season through a break in monsoon rains during late June and early July. A depiction of this monsoon life cycle with Taiwan rainfall is presented in Fig. 1a [a modification of Chen et al.’s (2004) Fig. 6a]. The transition of monsoon rainfall regimes is caused by the sequential passage of the mei-yu rainband in early summer, the western Pacific subtropic high in midsummer, and the tropical cyclone activity in late summer. Dynamically, the northward migration of these monsoon elements is caused by the coupling of the Asian monsoon circulation with the eastward propagation of the global intraseasonal mode (Chen and Murakami 1988). This coupling is reflected by a distinct out-of-phase intraseasonal oscillation between the mei-yu rainband along the Yangtze River Valley and the monsoon rainband across the South China Sea (Chen et al. 2000). In addition to this intraseasonal oscillation, it was observed by Samel et al. (1995) that interannual variation of rainfall in East Asia is centered over the Yangtze River Valley and southeastern China. In the former region, rainfall intensifies when the interaction between major elements of the eastern Asian monsoon circulation causes the temperature gradient across the mei-yu rainbelt to increase. The rainfall in southeastern China intensifies, when the monsoon continental thermal low moves to the north.

Fig. 1.
Fig. 1.

(a) The East–Southeast Asian summer monsoon life cycle [active phase (1/5–15/6), break phase (16/6–15/7), and revival phase (16/7–15/9)], depicted with the climatological 15-day mean rainfall for Taiwan from April to October [adopted from Chen et al. (2004)], and (b) the rainfall for Taiwan contributed by different weather systems [including diurnal variation (localized convection), typhoons, rainstorms, and fronts] for different phases of the summer monsoon life cycle, measured with Wang and Chen’s (2008) approach and averaged only over the southwestern region (21.7°–24.3°N, 120°–120.8°E) of Taiwan.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

According to Ramage’s (1952) observation, the monsoon rainfall in Southeast–East Asia is primarily produced by fronts during the first phase of the monsoon life cycle (the mei-yu season) and by tropical cyclones during the revival phase. Using the global analyzed precipitation, including the Climate Prediction Center (CPC) Merged Analysis of Precipitation (CMAP) Project (Xie and Arkin 1997), the Global Historical Climatological Network (GHCN) station rainfall (Easterling et al. 1996), and Japan Meteorological Agency (JMA) 6-h surface analysis maps, Chen et al. (2004) showed the mei-yu rainband coincides with major frontal activity from southern China to southern Japan during the mei-yu season. Following the identification approach of different weather systems across Taiwan adopted by Wang and Chen (2008), rainfall produced by these weather systems was measured by the rain gauge networks available—the Automatic Rainfall and Meteorological Telemetry system (Chen et al. 1999) and conventional surface stations of the Weather Bureau and the Environmental Protection Administration in Taiwan. A histogram of rainfall averaged over the southwestern region (21.7°–24.3°N, 120°–120.8°E) of Taiwan is shown in Fig. 1b. Over half of the rainfall is produced by rainstorms during the active monsoon phase, while slightly less than half of the rainfall is produced by tropical cyclones during the revival monsoon phase.

Performing an empirical orthogonal function analysis of station rainfall over China, Samel et al. (1995) identified two regions of large interannual variability—the Yangtze River Valley and southeastern China. Can this interannual variation of monsoon rainfall be reflected by the activity of rain-producing weather systems? Although some extensive analysis is lacking, a severe heavy rainfall event and a wet mei-yu season that occurred along the Yangtze River Valley in 1991 (Wang et al. 2000) were observed, following the development of a major warm ENSO episode. This unusual rainfall season over the Yangtze River Valley seems to suggest an answer to the question raised above. In contrast, the possible interannual variations of rain-producing weather systems in the northern part of the South China Sea (SCS) and southeastern China and their causes have not been explored so far. As inferred from contributions of different weather systems during the mei-yu season (shown in Fig. 1b), the interannual variation of monsoon rainfall during the mei-yu season is likely attributed to the interannual variations of rainstorm populations and rain-producing efficiency between the southwestern China–northern Indochina region and southern Japan.

An effort was made in this study to explore the interannual variation of this rainstorm activity and its contribution to the monsoon rainfall during the mei-yu season over the northern part of Southeast Asia and the southern part of East Asia for the period of 1979–2008. Because the analysis period covers three decades, the rainfall and reanalyses used were derived from multiple data sources described in section 2. The interannual variations of rainfall and population and rain-producing efficiency of rainstorms over the concerned region are presented in section 3. The cause of these interannual variations and its link with interannual variations of monsoon westerlies in this region and the associated anomalous monsoon circulation are analyzed in section 4. The interannual variation of the Asian monsoon is often depicted by the change in rainfall. Thus, the rainstorm contribution to the monsoon rainfall in Southeast–East Asia during the mei-yu season adds a new perspective to the interannual variation of the Asian monsoon. The mechanism of this monsoon variation is often tested with global climate models (Sperber et al. 2001). The role played by rainstorms in this monsoon variation offers a new dimension to this effort. A summary of this study and remarks about the climate simulations during the Asian monsoon region are provided in section 5.

2. Identification and composite rainfall produced from rainstorm and data

Interannual variation of monsoon rainfall in the northern Southeast Asia–southern East Asia region contributed by rainstorms for the 1979–2008 period is the focus of the present study. Thus, the data are essentially determined by how rainstorms are identified during late spring–early summer. The criteria to identify a rainstorm are outlined below.

  1. A rainstorm is synoptically coupled with (or ahead of) a midtropospheric (700–600 hPa) subsynoptic-scale trough (indicated by a thick short line in Fig. 2a), slightly west of this rainstorm’s genesis location, identified by a clear IR center in a Multifunctional Transport Satellite (MTSAT) image (Fig. 2d).

  2. Vorticity of a rainstorm at its genesis location, usually in the midtroposphere, should have a value ≥ 2 × 10−5 s−1. This location is determined by backtracking from the identified rainstorm.

  3. Rainfall in a rainstorm should reach an amount ≥ 50 mm·6 h−1 at the mature stage of its life cycle.

  4. Underneath the midtropospheric subsynoptic-scale trough, a cyclonic vortex in the lower troposphere can be identified either with the surface analysis (Fig. 2b), the JMA surface analysis (Fig. 2c), or the QuikSCAT image (Fig. 2e) close to the date of rainstorm genesis.

  5. Regardless of its genesis location (either in the northern Indochina–southwestern China region or the northern South China Sea), the trajectory of a rainstorm must reach the vicinity or east of Taiwan, and its life span should be ≥ 2 days.

  6. As measured by the subsynoptic-scale disturbance in the midtroposphere or by the cyclonic vortex in the lower troposphere, the horizontal scale of a rainstorm is ≤ O(103 km) (synoptic-scale disturbance), but ≥ O(102 km) (mesoscale convective system).

The identification genesis of a rainstorm at 1200 UTC 18 May 2008 is shown in Fig. 2 as an example.
Fig. 2.
Fig. 2.

A typical rainstorm case identified at 1200 UTC 18 May 2008 is shown: (a) 700-hPa streamline chart superimposed with vorticity, (b) surface streamline chart superimposed with precipitation of TRMM, (c) JMA surface analysis, and corresponding (d) satellite image of Multifunctional Transport Satellite (MTSAT) and (e) surface wind vectors of the Quick Scatterometer (QuikSCAT). (top right) Scales of vorticity and precipitation are shown in (a) and (b), respectively. The genesis location of this rainstorm is pointed by an arrow attached to symbol R1 (the first rainstorm identified in the mei-yu season of 2008). The arrow with R1 in (e) is pointing to the location of this rainstorm at 1100 UTC 19 May 2008.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

Criteria (ii)–(vi) should be identified with analyses of surface and upper-air winds, while all criteria [(i)–(vi)] should be assessed by the following rainfall data sources and rainfall proxy—surface [World Meteorological Organization (WMO)/GHCN] station and satellite [e.g., Tropical Rainfall Measuring Mission (TRMM)] rainfall measurements, global precipitation analyses [e.g., CPC morphing technique (CMORPH)], rainfall proxy [e.g., Geostationary Operational Environmental Satellite (GOES) precipitation index (GPI)], and satellite images generated with IR or blackbody brightness temperature (TBB) measurements. The data sources used for these two groups of criteria are designated as dataset I and II, respectively. Information regarding acronyms, spatial and temporal resolutions, spatial domain, available time period, and references of analyzed datasets are provided in Tables 1 and 2. Because the analysis period covers three decades, it is difficult (if not impossible) to compile all available data sources in a uniform manner to facilitate the identification of rainstorms. Therefore, it was required that the horizontal scale of a rainstorm occur between O(103 km) and O(102 km). Thus, the identification and depiction of a rainstorm need the data of horizontal resolution ≤O(102 km). To properly use these available datasets to identify rainstorms over three decades, the manual identification procedure is adopted, which is a time-consuming task.

Table 1.

Dataset-I data: surface and upper-air wind fields.

Table 1.
Table 2.

Dataset-II precipitation: station rainfall, analyzed global precipitation, and rainfall proxy derived from satellite observations.

Table 2.

Since rainfall produced by a rainstorm (PRS) is crucial information needed to examine the major theme of this study, the PRS distribution is prepared using the following steps:

  1. The identified rainstorm may start to merge with an East Asian front 2~3 days on average after its genesis. As long as the vortex of this rainstorm can be recognized, the rainfall associated with this vortex is included in the PRS distribution.

  2. The time for most rainstorms to propagate from their genesis locations either over northern Vietnam or the northern part of the South China Sea to Okinawa is about 5 days. The daily location of an average rainstorm center on a given day of its life cycle is the average of all rainstorms (belonging to the same genesis time window of either evening or early morning) with a given day in their life cycles.

  3. Daily rainfalls produced by all rainstorms with their centers located around the daily-averaged center (red dot) on the same given day of their life cycles are summed up.

  4. The PRS distribution is then obtained by piecing together the composite daily rainfall of all rainstorms obtained in step 3. The process of this step will be illustrated later by examples shown in section 3b(2).

  5. The daily rainfall histogram of the composite rainstorm on a given daily is the area average of the daily ensemble rainfall coverage of this day obtained in step 3. This daily rainfall histogram is used to define the intensity of the composite rainstorm life cycle.

3. Interannual variation of rainfall: Contribution of rainstorms

After the rainstorms were identified by criteria presented in section 2, the backtracking procedure was used to track/locate genesis locations of all rainstorms from 1979 to 2008 (Fig. 3a). They are marked by red dots over the northern Indochina–southwestern China region and by blue dots over the northern part of the South China Sea. The rainstorm genesis exhibits a location/time preference basically regulated by the day–night alternation of the land–sea thermal contrast. Land is warmer than water in the afternoon and evening but cooler than water in the early morning. Thus, most rainstorm genesis occurs over land in the former time period and over sea in the later time period. The purpose for keeping this distinction of genesis location/time is because it will be used to illustrate the environmental conditions of the rainstorm genesis over land and sea later in section 4. Trajectories of these rainstorms are depicted by color lines corresponding to their genesis locations. As indicated by their trajectories, all rainstorms move from genesis locations, across the northern South China Sea, South China, and Taiwan, to southern Japan. The primary rain-producing weather system in the northern part of Southeast Asia and the southern part of East Asia is a rainstorm during late spring–early summer [May–June (MJ)]. Following the procedure outline in section 2, the PRS distribution during this phase of the monsoon life cycle is shown in Fig. 3b. It is no surprise to see that the PRS distribution coincides with rainstorm trajectories.

Fig. 3.
Fig. 3.

(a) Trajectories of rainstorms are depicted by lines with color corresponding to genesis locations. Genesis locations of rainstorms for Indochina (the South China Sea) during May–June over the 1979–2008 period are marked by red (blue) dots, (b) rainfall contributed by rainstorms (PRS), (c) distribution of the climatological rainfall (PT) ensembled with TRMM, CMORPH, and Global Precipitation Climate Project (GPCP) for May–June, and (d) ratio PRS/PT. The contour interval of (b),(c),(d) is 1 mm day−1, 1 mm day−1, and 5%, respectively.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

The major contribution t the total amount of rainfall (PT) during the late spring–early summer along the mei-yu rainband (Fig. 3c) has long been considered by the East Asian monsoon community to be produced by mei-yu fronts/mesoscale convective systems (MCSs) (e.g., Ding and Sikka 2006). This rainband stretches from northern Vietnam, across southern China and Taiwan, to southern Japan. The distribution of the ratio PRS/PT is shown in Fig. 3d. The patterns of PRS ≥ 6 mm day−1 and PRS/PT ≥ 50% bear a close resemblance. This pattern resemblance between PRS and PRS/PT indicates the primary rainfall contribution to the mei-yu rainband in its western section during the mei-yu season (May–June) is derived from rainstorms, instead of fronts.

a. Population and rainfall statistics of rainstorms

Monsoon rainfall in southern China exhibits interannual variability, out-of-phase with that along the Yangtze River Valley (Samel et al. 1999). Because part of the southern China rainfall in summer is the result of the mei-yu rainband, the monsoon rainfall in the western section of the mei-yu rainband during the first phase of the East–Southeast Asian summer monsoon life cycle likely undergoes an interannual variation too. As the major rain producer in the western section of the mei-yu rainband, the rainstorm’s population and rainfall may also exhibit interannual variation in accord with the mei-yu rainband. To substantiate this inference, histograms of the rainstorm population (NRS) over the region with PRS ≥ 6 mm day−1, superimposed with ΔSST (Niño-3.4) [the SST (Niño-3.4) (MJ) departure from its average over the period of 1979–2008], are shown in Fig. 4a. Following the interannual variation of ΔSST (Niño-3.4), NRS is generally larger (smaller) than its long-term mean value (≃8 per season), when ΔSST (Niño-3.4) ≥ 0.5°C (≤−0.5°C) during late spring–early summer. The coincident interannual variations of NRS and ΔSST (Niño-3.4) are substantiated by the large correlation coefficient, 0.82 (with a significant confidence level of 90%), between these two variables. The former thermal condition of SST is defined as warm late spring–early summer, while the latter one is defined as cold late spring–early summer. The possible mechanism causing this NRS interannual variation will be presented in section 4. The area-average rainfall produced by rainstorms during May–June over the region of PRS ≥ 6 mm day−1 is shown in Fig. 4b. The three-decadal average of PRS over this defined region is 7.1 mm day−1. It is clearly revealed from the contrast between the NRS and PRS histograms that the interannual variation of PRS closely follows that of NRS.

Fig. 4.
Fig. 4.

(a) Occurrence frequency of rainstorms NRS over the region of PRS ≥ 6 mm day−1 during May–June for the 1979–2008 period, superimposed with the ΔSST (Niño-3.4) index averaged over May–June and precipitation histograms of (b) rainstorm PRS, (c) total rainfall PT, and (d) the ratio, PRS/PT, over the region of PRS ≥ 6 mm day−1. Averaged values of all variables are also presented in each panel. The thermal condition of SST over the NOAA Niño-3.4 area (5°S–5°N, 170°–120°W) is determined by the following criteria: warm ΔSST (Niño-3.4) ≥ 0.5°C and cold ΔSST (Niño-3.4) ≤ −0.5°C. The warm and cold thermal conditions are colored red and blue, respectively, on histograms of all variables shown in this figure.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

If rainstorms are the major rainfall producer along the western part of the mei-yu rainband, how much PT is contributed by rainstorms? To answer this question quantitatively, a histogram of PT averaged over the region with PRS ≥ 6 mm day−1 during May–June is shown in Fig. 4c; the interannual variations of the PRS and PT histograms are coincident. The long-term average of PT is 11.5 mm day−1, about 4 mm day−1 larger than PRS, but the long-term average ratio PRS/PT (Fig. 4d) indicates that PRS contributes slightly over 62% of PT—most of the monsoon rain along the western part of the mei-yu rainband. However, our major concern is whether the interannual variation of PT along this part of the mei-yu rainband is primarily caused by that of rainstorms, and, in turn, interannual variations of rainstorm population and rain production. It is clearly revealed from Fig. 4d that the ratio PRS/PT always exceeds its average value of 62% during warm years and drops below this average value during cold years. These statistics indicate the interannual variation of PT along the western section of the mei-yu rainband is heavily affected by rainstorms.

Since there are few surface stations available over the ocean, the rainfall data used in this study are derived from several different sources, as described in section 2. Using different data sources and analysis algorithms, the analyzed global precipitation generated by different agencies and projects may have analysis bias, particularly over the ocean, because of a lack of station measurement. However, to verify results of the analysis shown in Fig. 4, surface observations made in Hong Kong and Taiwan are analyzed and presented in Fig. 5. Both islands are located in the major path of rainstorms, but stations on these two islands can only detect part of the rainstorm population because of their size. Recall the horizontal scale of rainstorm varies between O(102 km) and O(103 km). Despite the small sizes of these two islands, the numerical numbers of rainstorm passage across them include not only those with their centers moving directly through, but also those whose rain can reach these islands. Thus, observations from these two islands may not provide interannual variations of NRS, PRS, and PT during May–June for all rainstorms. However, distinct signals of these interannual variations are well reflected by observations made by surface stations on these two islands.

Fig. 5.
Fig. 5.

As in Fig. 4, but for (left) Hong Kong and (right) Taiwan.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

Numerical values of NRS detected at Hong Kong (Fig. 5a) and Taiwan (Fig. 5e) are 5.8 and 5.4, respectively. These values are smaller than NRS over the region with PRS ≥ 6 mm day−1 (Fig. 4a). This numerical difference in the rainstorm population between these two islands and the western part of the mei-yu rainband is attributed to the limited propagation of rainstorms across these two islands. For convenience, designate MR, HK, and TW as the mei-yu rainband, Hong Kong, and Taiwan, respectively. Interannual variations of NRS(HK) and NRS(TW) closely follow ΔSST (Niño-3.4). Thus, PRS(HK) (Fig. 5b) and PRS(TW) (Fig. 5f) should undergo a coincident interannual variation with PRS(MR). In addition, the average values of PRS(HK) and PRS(TW) are 8.3 and 10.2 mm day−1, respectively, and larger than that of PRS(MR), despite the fact that numerical values of both NRS(HK) and NRS(TW) are smaller than NRS(MR). The averaged values of PT (HK) (Fig. 5c) and PT(TW) (Fig. 5g) are 13.7 and 15.7 mm day−1, respectively, larger than PT(MR), but both PT(HK) and PT(TW) exhibit an interannual variation coincident with the ΔSST(Niño-3.4) index, as PRS(HK) and PRS(TW) do.

Average ratios PRS(HK)/PT(HK) and PRS(TW)/PT(TW) shown in Figs. 5d and 5h, respectively, are slightly over 60%, close to that of PRS(MR)/PT(MR) (Fig. 4d). During the warm (cold) ΔSST (Niño-3.4) May–June, PRS/PT at both Hong Kong and Taiwan, PRS(MR)/PT(MR), are always larger (smaller) than their respective average values. This contrast shows the interannual variation of PT at these two islands is primarily caused by the rainstorms. Interannual variations of NRS, PRS, PT, and PRS/PT analyzed with surface observations at Hong Kong and Taiwan are consistent with those over the western part of the mei-yu rainband (Fig. 4). As the primary rainfall contributor over this region, interannual variations of the rainstorm population and the rainfall produced by rainstorms result in the interannual variation of monsoon rain in this part of the mei-yu rainband during May–June.

b. Rain-producing efficiency of rainstorms

One may infer from the coincident interannual variations of NRS, PRS, and PT in the northern part of the South China Sea (Figs. 4 and 5) that interannual variations of the latter two variables follow rainstorm population. However, it is observed from previous studies (Nakamura et al. 2002; Chang et al. 2002) that precipitation along the North Pacific storm track undergoes an interannual variation following the jet stream in response to the tropical Pacific SST anomalies. These studies lead to the following concern. In addition to the change in the rainstorm population, are the interannual PRS and PT variations also related to some basic features of rainstorm activity in response to the change of the large-scale monsoon circulation? To clarify this concern, several features of rainstorm activity are presented in Fig. 6.

1) Genesis location and track

During both warm and cold May–June, locations of rainstorm genesis and the ensuing tracks are shown in Figs. 6a and 6c, respectively. The average tracks of rainstorms in these two extreme climate conditions are portrayed by thick red lines in (b) and (d). One can find that rainstorm tracks during warm years are, on average, somewhat north of those during cold years. The northeastern propagation of rainstorms is caused by the advection of monsoon southwesterlies. This minor northward migration of rainstorm tracks is a reflection of the change in the Southeast–East Asian monsoon circulation in response to the tropical Pacific SST anomalies during May–June, mainly through a dynamic process instead of a hydrological outcome.

Fig. 6.
Fig. 6.

(a) Genesis locations (red/blue dots) and track (red/blue lines) during warm May–June for rainstorms and (b) PRS(warm) produced by rainstorms superimposed with average daily locations of rainstorm centers during warm May–June. (c),(d) As in (a),(b), respectively, but for cold May–June, histograms of average daily PRS for (e) all identified rainstorms, (f) warm and (g) cold May–June, and (h),(i),(j) as in (e),(f),(g), but for convergence of water vapor flux. Note that the average tracks in (b),(c) are portrayed by a thick red line connecting the average locations (red dots) of rainstorm centers over the composite life cycle of these rainstorms, which is depicted by the day number attached to those average locations. The center location of rainstorms included in the averaging process is covered by red-dashed ellipse and its peripheral area. The axes of these ellipses are one standard deviation of rainstorm distances from the average center parallel and perpendicular to the averaged track.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

2) Rain-producing efficiency

Following the procedure to prepare the rainfall distribution contributed by all identified rainstorms PRS, the PRS distribution during warm and cold MayJune, PRS(warm) and PRS(cold), are shown in Figs. 6b and 6d, respectively. Distributions of both PRS(warm) and PRS(cold) coincide with the corresponding rainstorm tracks. These rainfall distributions indicate the magnitude of PRS(warm) is larger than that of PRS(cold). More quantitatively, the daily rainfall amounts of PRS, PRS(warm), and PRS(cold) over the composite life cycle of rainstorms are shown in Figs. 6e–g. Even on a daily basis, one can easily find that PRS(warm) > PRS(total) > PRS(cold). It is clearly indicated by this comparison of daily rainfall amounts that the rain-producing efficiency is larger (smaller) during warm (cold) year. To summarize this comparison with average daily rainfall amount over the composite life cycle of rainstorm, PRS(total), PRS(warm), and PRS(cold) are 7.1, 8.9, and 5.4 mm day−1, respectively. The rain-producing efficiency can be measured with the following ratios: PRS(warm)/PRS ≃ 125% and PRS(cold)/PRS ≃ 84%. The difference in rain-producing efficiency between warm and cold years is about 49%. The cause of this efficiency difference can be inferred from the water vapor budget analysis.

Rainfall P is maintained primarily by the convergence of water vapor flux ,
e1
where Q (water vapor flux) , g, ps, V, q, and p are gravity, surface pressure, velocity vector, specific humidity, and pressure, respectively. Evaporation is neglected in this approximated water vapor budget equation because its contribution in maintaining rainstorm rainfall is relatively minor. Note that P and are localized hydrological processes, but water vapor flux is driven by the large-scale circulation embedded by the rainstorm. Apparently, the rain-producing efficiency of a rainstorm is vitally affected by the water vapor supply from the large-scale environmental flow.

Following the procedure in computing PRS histograms (Figs. 6e,f), the corresponding daily ()RS histograms are shown in Figs. 6h–j. The chronicle evolution of the rainstorm hydrological cycle over its composite life cycle is revealed from the daily histograms of PRS and −()RS over the composite life cycle of the rainstorm. The daily magnitude of −()RS is somewhat smaller than that of corresponding PRS. This discrepancy may be contributed by evaporation, which is part of the water vapor source neglected in the approximated water vapor budget. Comparing histograms of −()RS shown in Figs. 6e–g, one finds −()RS(warm) > −()RS > −()RS(cold). According to the approximated water vapor budget and the contrast between daily −()RS(warm) and −()RS(cold) histograms, it is inferred that the rain-producing efficiency of rainstorms is related to the water vapor supply by the environment. This inference will be substantiated further in section 4b.

4. Possible mechanism

A summer cross-Pacific short-wave train around the North Pacific rim can be induced by warm sea surface temperature (SST) anomalies over the western tropical Pacific (Nitta 1987). An anomalous cyclonic cell adjacent to the western part of these SST anomalies is juxtaposed with an anomalous anticyclonic cell centered over northeastern Asia to form the so-called Pacific–Japan oscillation (PJO). The monsoon westerlies over northern Indochina may be intensified (weakened) by the PJO, while those along the Yangtze River Valley may be weakened (intensity) when the positive (negative) SST anomalies appear over the western tropical Pacific. Samel et al.’s (1999) out-of-phase interannual variation of rainfall between southern China and the Yangtze River Valley during summer is likely linked to the interannual variation of monsoon westerlies coupled with the PJO. The major rain-producing weather system before the break is the rainstorm, while those after the break are related to typhoons and the diurnal cycle. Thus, interannual variations of rainfall before and after the break should be related to different weather systems. Therefore, we shall focus on the search for the possible mechanism causing interannual variations of 1) rainstorm genesis and 2) rain-producing efficiency of a rainstorm.

a. Rainstorm genesis

The genesis mechanism of rainstorms during May–June was explored/illustrated from two perspectives: 1) synoptic environment and 2) dynamic instability. Therefore, the possible mechanism of the interannual variation in rainstorm genesis may be demonstrated through interannual variations of synoptic environment and dynamic intensity.

1) Synoptic environment

During May–June, the large-scale monsoon circulation over Southeast–East Asia consists of the Tibetan high in the upper troposphere and the continental thermal low in the lower troposphere surrounded by monsoon westerlies over Indochina and the ocean around the eastern seaboard of southern China. The synoptic environment favorable for the genesis of rainstorms is the formation of a strong shear by the midtropospheric northwesterly flow around the northeastern periphery of the Tibetan Plateau and lower-tropospheric monsoon westerlies across northern Indochina or the northern part of South China, as inferred from Fig. 7a. The northwesterly flow around the northeastern periphery of the Tibetan Plateau is generally the cold surge–like flow straddling a high–low couplet, coupled with an eastward-propagating synoptic-scale short wave in the upper troposphere. In contrast, after the onset of the Southeast–East Asian summer monsoon, monsoon westerlies persistently exist across Indochina, more stationary than the midtropospheric cold surge–like flow. Thus, the interannual variation of the environmental flow favorable for the formation of the monsoon shear line from northern Indochina to southern China may well be reflected by the interannual variation of monsoon westerlies.

Fig. 7.
Fig. 7.

Streamline charts at 700 hPa superimposed with (a) isotach |V|, and (b) variance of U(700 hPa) during May–June for the 1979–2008 period, (c) time series of u(700 hPa) averaged over the area of (15°–25°N, 100°–115°E) (solid line with dots) surrounding the maximum of u(700 hPa) variance at (20°N, 104°E) and ΔSST (Niño-3.4) (dash line with open circles), (d) the height–time cross section of u averaged over the area of (15°–25°N, 100°–115°E), and (e) the latitude–time cross section of u(700 hPa) over the longitudinal zone of (100°–115°E). (top right) Scales of |V| and Var(U) are shown in (a),(b), respectively. Measured in terms of ΔSST (Niño-3.4), the warm ENSO events [ΔSST (Niño-3.4) ≥ 0.5°C] are marked as W in (d),(e). The contour interval of (d),(e) is 0.5 m s−1, and (bottom right) the scale is shown in (d). The monsoon trough in (a),(b) is indicated by a thick, dashed blue line.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

As shown in Fig. 7a, strong 700-hPa monsoon westerlies appear over northern Indochina and extend northeastward to Japan. The variance of u(700 hPa), Var[u(700 hPa)], in Fig. 7b exhibits a region of significant values over northern Vietnam. This variance maximum indicates the location of the u(700 hPa) interannual variation. The time series of u(700 hPa) (MJ) averaged over an area (15°–25°N, 100°–115°E) around this variance maximum is shown in Fig. 7c by a dotted, thick solid line, along with ΔSST(Niño-3.4), open circle, dashed thin line. Surprisingly, the interannual variation of u(700 hPa) (MJ) over northern Vietnam is in phase with ΔSST(Niño-3.4). The contrast between [u(700 hPa, 15°–25°N, 100°–115°E), ΔSST(Niño-3.4)] (Fig. 7c) versus NRS (Fig. 4a) clearly shows that rainstorm genesis occurs more (less) frequently when monsoon westerlies are strong (weak). The following two issues are raised from the coherent interannual variations between these three variables:

  1. Why do interannual variations of NRS and monsoon westerlies coincide?

  2. How is the interannual variation of monsoon westerlies in the northern Indochina–southern China region coherently linked to ΔSST(Niño-3.4)?

(i) Issue 1: NRS versus monsoon westerlies

The relationship of interannual variations between NRS and u(700 hPa) in northern Vietnam is illustrated by the contrast between the pressure–time and latitude–time cross sections in Figs. 7d and 7e, and time series of [u(700 hPa), ΔSST(Niño-3.4)] in Fig. 7c. Monsoon westerlies strengthen (weaken) and deepen (shallow) when ΔSST(Niño-3.4) is ≥ 0.5°C (≤−0.5°C).1 The relationship between monsoon westerlies and ΔSST(Niño-3.4) will be illustrated further in our answer to issue II, but the relationship between NRS and monsoon westerlies is discussed here. When monsoon westerlies strengthen (weaken) and deepen (shallow), the monsoon shear flow formed by monsoon westerlies and midtropospheric northwesterlies around the northeastern periphery of the Tibetan Plateau is intensified (weakened). Consequently, the intensification (weakening) of this shear north of monsoon westerlies facilitates (hinders) rainstorm genesis and leads to the increase (decrease) of NRS.

(ii) Issue 2: Interannual variation of the monsoon circulation is reflected by monsoon westerlies

The interannual variation of monsoon westerlies should be an indicator of the interannual variation of monsoon circulation. Thus, the relationship between interannual variations of Δu[700 hPa, (15°–25°N, 100°–115°E)] and ΔSST(Niño-3.4) is explored in terms of warm and cold composite charts of eddy streamfunction anomalies at 700 hPa, ΔψE(700 hPa), shown in Figs. 8b and 8c. As indicated by a heavy arrow, Δu(700 hPa) in northern Vietnam and southern China increases(decreases) when ΔSST(Niño-3.4) ≥ 0.5°C (≤−0.5°C). It is inferred from the coincident interannual variations of Δu[700 hPa, (15°–25°N, 100°–115°E)] and ΔSST(Niño 3.4) time series shown in Fig. 7c that the anomalous monsoon circulations depicted by Figs. 8b and 8c are the response of the monsoon circulation to ΔSST anomalies.

Fig. 8.
Fig. 8.

(a) Eddy of streamfunction at 700 hPa, ψE(700 hPa), superimposed with isotachs (red) averaged over May–June for the 1979–2008 period. Based on the index of ΔSST (Niño-3.4, MJ) shown in Fig. 7c, composite ψE(7000 hPa) anomalies [ΔψE(700 hPa)] superimposed with ΔSST anomalies for (b) warm and (c) cold events determined by ΔSST (Niño-3.4) anomalies averaged over May–June. Contour intervals of ψE in (a) and ΔψE in (b),(c) are 106 m2 s−1 and 3 × 105 m s−1, respectively. (bottom right) Scales of |V|(700 hPa) and ΔSST are shown in (a) and (b),(c), respectively. Arrows are added on ψE(700 hPa) and ΔψE(700 hPa) to indicate the flow directions.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

To substantiate the aforementioned inference, the long-term-averaged (ψE, |V|) charts covering major parts of the northern Indian Ocean and the Pacific are shown in Fig. 8a. The 700-hPa summer circulation over the Pacific and Indian oceans is characterized by the North Pacific anticyclone and the Asian continental thermal low north of the equator and the South Pacific anticyclone and the Indian Oceanic anticyclone south of the equator. The South Asian monsoon westerlies, the East Asian monsoon southwesterlies, and the Pacific trade winds are indicated by arrows. The warm composite ΔSST anomalies (Fig. 8b) are signified by the eastern tropical warm tongue surrounded by cold SST anomalies in the North and South Pacific and juxtaposed in the west with the warm SST anomalies over a large part of the tropical Indian Ocean, particularly in the western section. Compared to ψE (700 hPa) (Fig. 8a), the response of the lower-tropospheric circulation to SST anomalies during the warm phase depicted by the composite ΔψE (700 hPa) anomalies is the poleward migration of oceanic anticyclones in the Indian Ocean and in both the North and South Pacific. Portrayed with the lower-tropospheric flow, the response is reflected by the tropical westerly anomalies in the Pacific and the tropical easterly anomalies in the Indian Ocean, a divergence of anomalous tropical flows over the Maritime Continent, as the surface tropical wind changes during the warm ENSO events (Philander 1990). The monsoon westerlies in South Asia move northward and the monsoon southwesterlies east of Japan are intensified. The spatial pattern of composite ΔSST anomalies during the cold phase is reversed (Fig. 8c); the response of the lower-tropospheric circulation to the SST change during the cold phase is opposite to the warm phase. This response is reflected by the strengthening of the Pacific trade easterlies in the east, the South Asian monsoon westerlies in the west, and the convergence between monsoon westerlies and the Pacific trade easterlies over the tropical western Pacific. The monsoon southwesterlies in East Asia also weaken, as indicated by the u[700 hPa, (15°–25°N, 100°–115°E)].

2) Dynamic instability

Rainstorm genesis generally occurs in the midtroposphere (700–600 hPa) over northern Vietnam–southwestern China and the northern part of the South China Sea (Fig. 3) because the environmental flow over these two regions usually meets the Charney–Stern instability criterion (Charney and Stern 1962) during May–June. As inferred from Fig. 7, the strengthening (weakening) and deepening (shallowing) of monsoon westerlies facilitate (hinder) the formation of the monsoon shear line and rainstorm genesis. In view of this possible relationship between rainstorm genesis and the environmental flow embedded by the monsoon shear line, it is strongly suggested that the Charney–Stern instability criterion is much easier to satisfy when monsoon westerlies are strong.

The Charney–Stern criterion includes two elements:

  1. a sign change in the meridional gradient of potential vorticity (q) , and

  2. a maximum gradient of potential temperature θ north of the location of the sign change and south of the region of the negative .

The composite latitude–height cross sections of and the latitudinal distribution of at the surface at two longitudes cutting through northern Vietnam (at 104°E) and the northern part of the South China Sea (at 114°E) are shown in Figs. 9a and 9d, respectively, when monsoon westerlies are strong. The strength of monsoon westerlies is defined by their magnitude indicated by the u[700 hPa, (15°–25°N, 100°–115°E)] time series. Monsoon westerlies are defined as strong (weak) when the u[700 hPa, (20°N, 104°E)] magnitude is larger (smaller) than its mean value, plus one positive (negative) standard deviation. It is clearly revealed from Figs. 9a and 9d that a sign change of occurs at 600 hPa close to 20°N, when monsoon westerlies are strong and a maximum appears slightly north of the location where changes signs. Apparently, the Charney–Stern instability criterion is met by the environmental flow over northern Vietnam (104°E) and the northern South China Sea (114°E), when monsoon westerlies are strong (Figs. 9a,d). In contrast, when monsoon westerlies are weak, the Charney–Stern instability criterion can be met by the composite latitude–height cross section of and the latitudinal distribution of at the surface (Figs. 9b,e), but these meridional gradients are much smaller in magnitude.

Fig. 9.
Fig. 9.

Latitude–height cross sections of the meridional gradient of potential vorticity added with the meridional gradient of 925-hPa potential temperature (925 hPa) at (left) 104°E and (right) 114°E. (a),(d) Strong monsoon westerlies, (b),(e) weak monsoon westerlies, and (c),(f) the difference between strong and weak monsoon westerlies are shown. The contour interval of and is 10−6 m2 s−1 K kg−1.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

The comparison of the Charney–Stern instability criterion between strong and weak monsoon conditions can be quantitatively assessed by the difference of this instability criterion in Figs. 9c and 9f. It becomes clear that the strong (weak) monsoon westerlies in northern Indochina strengthens (weaken) the monsoon shear so that the Charney–Stern instability criterion of the environmental flow can (cannot) be met and rainstorm genesis can occur more (less) frequently.

b. Environment enhancing the rain-producing efficiency

According to Chen (1985), the water vapor flux may be separated into divergent () and rotational () components,
eq1
These two components of water vapor flux can be expressed with the potential function () and streamfunction () of water vapor flux,
eq2
Thus, the divergence of water vapor flux may be written as
eq3
The major content of water vapor resides in the lowest layer of the atmosphere. Thus, the water vapor transport is driven primarily by the lower-tropospheric circulation. Because water vapor is a scalar variable, spatial patterns for both and resemble those of velocity potential () and streamfunction (ψ) in the lower troposphere.

The onset of the Asian monsoon and the active monsoon phase in the South, Southeast, and part of East Asia occur during May–June. During this time period, the global water vapor converges toward the Asian monsoon hemisphere so the global field exhibits a wave-1 structure around a latitude circle (not shown), but its Asian monsoon hemisphere portion is shown in Fig. 10a. The convergent center of is located over the Indochina–South China Sea–Philippine Sea region. The distribution of rainfall radiates from the Asian monsoon region eastward along the North Pacific convergence zone, the South Pacific convergence zone, and the intertropical convergence zone.

Fig. 10.
Fig. 10.

As in Fig. 8, but for (a) , (b) (warm), and (c) (cold). Contour intervals of and are 107 kg s−1 and 4 × 106 kg s−1, respectively.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

During warm (cold) May–June, the tropical SST anomalies exhibit positive (negative) anomalies over the tropical Indian ocean centered at 60°E and the National Oceanic and Atmospheric Administration (NOAA) Niño-3.4 area as shown in Fig. 8b (Fig. 8c). The response of the global divergent circulation in the lower troposphere to these two tropical SST anomaly centers during warm May–June is reflected by convergent centers of over the tropical Indian Ocean and the central-eastern tropical Pacific and a divergent region covering the Asian monsoon and Australia (Fig. 10b). As inferred from this low-level divergent circulation, the coupled east–west circulation in the tropics possesses the upward branches over the positive SST anomaly regions and a downward branch over the monsoon region as the tropical east–west circulation depicted by Streten and Zillman (1984). The downward branch of this east–west circulation suppresses weather development and convective activity and reduces rainfall. The reversed spatial structure of global divergent circulation and tropical east–west circulation occur during cold May–June, as inferred from Fig. 10c. The rainfall increases over the Asian monsoon region because the upward branch of the tropical east–west circulation appears there.

With the large-scale divergent circulation change resembling the fields presented in Figs. 10b and 10c, in response to the SST anomalies over the tropical Indian Ocean and the NOAA Niño-3.4 region, it seems unusual to find the interannual rainfall variation ΔP over the major PRS region (Fig. 2b) is opposite to those over the rest of the Asian monsoon region. The disparity of these monsoon rainfall variations leads to two basic issues of the regional hydrological cycle:

  1. What causes the disparity of the interannual rainfall variation between the major PRS region and the remaining Asian monsoon region?

  2. How does the interannual variation of facilitate or hinder the rain-producing efficiency of rainstorm?

1) Issue 1

Extending southward along 105°E from southwestern China to the eastern coast of Vietnam, a shallow monsoon trough indicated by a dashed blue line is discernable on the 700-hPa streamline charts (Figs. 7a,b). Overlapping with the favorable region of rainstorm genesis in southwestern China (Fig. 3a), this trough facilitates rainstorm genesis. Based on the tropical east–west circulation inferred from the patterns (Figs. 10b,c), it is expected that the genesis/development of synoptic disturbances in the Asian monsoon region should be suppressed (enhanced) by the downward (upward) branch of this east–west circulation. However, it was indicated by both the height–t diagram of u(15°–25°N, 100°–115°E) (Fig. 7d) and the yt diagram of u(700 hPa, 100°–115°E)(MJ) (Fig. 7e) that monsoon westerlies increase (reduce) during warm (cold) May–June in the northern Vietnam–southwestern China region. As indicated by the anomalous circulations in response to the tropical Pacific SST anomalies during May–June (Figs. 8b,c), the interannual variation of monsoon westerlies (Figs. 7d,e) is caused by anomalous anticyclonic (cyclonic) cell centered at the northern part of the South China Sea and the Philippine Sea. During warm (cold) May–June, the aforementioned anticyclone (cyclonic) cell exhibits two centers over the head Bay of Bengal and the northern part of the South China Sea; therefore, an anomalous trough (ridge) appears between these two centers. An anomalous convergent (divergent) zone of and the accompanied positive (negative) zone of ΔP are collocated ahead of the aforementioned trough (ridge). This interannual variation of the large-scale divergent circulation forms a favorable (unfavorable) environment to enhance (reduce) the rain-producing efficiency of rainstorms.

2) Issue 2

A histogram of superimposed with PT for the region of PRS ≥ 6 mm day−1 is shown in Fig. 11a, while the same for superimposed PRS is shown in Fig. 11b. These two quantities during warm and cold May–June are colored red and blue, respectively; their interannual variations are coincident with a correlation coefficient of 0.89 at a confidence level of 90%. Magnitude ratios of average /PT and /PRS are 90% and 93%, respectively. Interannual variation magnitudes of and may be measured statistically by their standard deviations—both are 1.5 mm day−1. Because is part of , it is implied by such close statistical values that the interannual variation of is largely modulated by . The latter is determined by the large-scale divergent circulation through because rainstorms are embedded in the large-scale divergent circulation and in turn the rain-producing efficiency of the rainstorm is affected by this circulation through .

Fig. 11.
Fig. 11.

Histograms of (a) convergence of total water vapor flux [] averaged over the region of PRS ≥ 6 mm day−1, superimposed with PT histograms, and (b) convergence of water vapor flux [] associated with PRS averaged over the region of PRS ≥ 6 mm day−1, superimposed with PRS histograms, too. The seasons with ΔSST (Niño-3.4) ≥ 0.5°C (≤−0.5°C) are colored by pink (light blue), while the corresponding PT and PRS histograms are colored by red and dark blue, respectively. Histograms of and during normal years [i.e. 0.5°C ≥ ΔSST(Niño-3.4) ≥ −0.5°C] are not colored, but those of corresponding PT and PRS are shaded.

Citation: Journal of Climate 24, 16; 10.1175/2011JCLI3930.1

5. Concluding remarks

The north–south extent of the eastern seaboard of Southeast–East Asia under the influence of the summer monsoon is over 4000 km, covering the tropics and midlatitudes. This long north–south distance causes the monsoon life cycle in East Asia to lag behind the monsoon life cycle in Southeast Asia. The out-of-phase intraseasonal oscillation of summer monsoon rainfall between the northern South China Sea and the Yangtze River Valley is a reflection of this phase lag (Chen et al. 2004). Thus, the rain-producing weather systems in different latitudinal zones may be different not only in the timing of their activity but also in their nature. It has long been regarded by the Asian monsoon community that rainfall along the mei-yu rainband during May–June is produced by mei-yu fronts. It was shown in this study that a major part of the May–June monsoon rainfall in the northern South China Sea, over northern Vietnam, across Taiwan, and into southern Japan is produced by rainstorms over the ocean.

It was shown by Samel et al. (1999) that the summer rainfall in South China exhibits an out-of-phase interannual variation with that in the Yangtze River Valley. An effort was made in this study to explore whether the rainfall over the northern part of the South China Sea undergoes an interannual variation during May–June, and the cause of this interannual variation. The major findings of this study are summarized as follows:

  1. Interannual variation of May–June monsoon rainfall: A distinct interannual variation emerges from the occurrence frequency of rainstorm genesis NRS rainfall produced by rainstorms PRS and total rainfall PT over the region (PRS ≥ 6 mm day−1). Interannual variations of NRS, PRS, and PT coincide with the interannual variation of SST over the NOAA Niño-3.4 region ΔSST (Niño-3.4); (NRS, PRS, PT) are larger (smaller), when ΔSST (Niño-3.4) ≥ 0.5°C (warm) [≤−0.5°C (cold)].

  2. Possible cause of (NRS, PRS, PT) interannual variations: In-phase interannual variations between (NRS, PRS, PT) and monsoon westerlies in northern Vietnam were observed. It is inferred from the in-phase interannual variations of monsoon westerlies over there and ΔSST (Niño-3.4) that interannual variations of (NRS, PRS, PT) are a reflection of the interannual variation of the monsoon circulation’s impact on the genesis occurrence frequency and rain-producing efficiency of rainstorms.

a. Genesis occurrence frequency

1) Synoptic perspective

The formation of the monsoon shear line between monsoon southwesterlies across northern Indochina and the midtropospheric cold surge–like northwesterly around the northeastern periphery of the Tibetan Plateau may be facilitated (hindered) by the intensification (weakening) and deepening (shallowing) of the monsoon southwesterlies.

2) Dynamic instability

The Charney–Stern instability criterion was shown to be the dynamic mechanism responsible for the rainstorm genesis, which consists of a sign change of meridional potential vorticity coupled with a maximum meridional gradient of potential temperature north of the sign change of . The Charney–Stern instability criterion is much easier (more difficult) to meet by the monsoon shear when monsoon westerlies are strong (weak).

b. Rain-producing efficiency

The response of the global divergent circulation in the lower troposphere to the SST anomalies in the tropical Indian and Pacific oceans leads to the development of the tropical east–west circulation with its upward (downward) branches over these two regions and a downward (upward) over the Asian monsoon region during warm (cold) May–June. On the other hand, the monsoon trough, extending southward from southwestern China to the coast of Vietnam, deepens (fills) during warm (cold) May–June, so that an anomalous convergent (divergent) zone of water vapor flux appears ahead of the deepened (filled) monsoon trough and is coincident with rainstorm tracks. The water vapor supply to this region is enhanced (depleted) so the rain-producing efficiency of rainstorms increases (decreases) during warm (cold) May–June.

These findings have two important implications:

  1. Seasonal prediction of the Southeast–East Asian climate: The ΔSST (Niño-3.4) index is operationally issued by the Climate Prediction Center, Washington, D.C. The in-phase relationship between (NRS, PRS, PT) and [u(monsoon westerlies), ΔSST (Niño-3.4)] may provide an excellent tool to monitor the interannual variation of the Southeast–East Asian monsoon.

  2. The seasonal predictions of South Asian monsoons performed by global climate model simulations have been under careful scrutiny (Sperber et al. 2001), but this research effort has not been expanded to cover the Southeast–East Asian monsoon. Findings about the interannual variations of monsoon rainfalls in Southeast–East Asia and their cause may provide a way to examine the performance of global climate models over this region.

Acknowledgments

This study is partially sponsored by the Cheney Research Fund and the NSF Grant ATM-0836220 and is greatly improved with comments and suggestions offered by two reviewers. The effort made by Ming-Cheng Yen is supported by the Grant NSC99-2111-M-008-012. Technical assistance provided by Paul Tsay is crucial to the revision of this manuscript.

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1

, long-term average value of (). The long-term average in this study covers the period of 1979–2008.

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