Abstract

The interannual variability of the western North Pacific (WNP) summer monsoon is examined for the non-ENSO, ENSO developing, and ENSO decaying years, respectively. The ENSO developing (decaying) year is defined as the year before (after) the mature phase of ENSO, and the non-ENSO year is defined as the year that is neither the ENSO developing year nor the ENSO decaying year. A strong (weak) WNP summer monsoon tends to occur during the El Niño (La Niña) developing year and a weak (strong) WNP summer monsoon tends to occur during the El Niño (La Niña) decaying year. In all non-ENSO, ENSO developing, and ENSO decaying years, the strong (weak) WNP summer monsoon is associated with the positive (negative) rainfall anomalies, cold (warm) sea surface temperature anomalies, warm (cold) upper-tropospheric temperature anomalies, low (high) surface pressure anomalies, and a low-level cyclonic (anticyclonic) circulation anomaly over the subtropical WNP. The 850-hPa wave train associated with the WNP and east Asian (EA) summer monsoons in the non-ENSO, ENSO developing, and ENSO decaying years extends northward and suggests a possible teleconnection between the WNP summer monsoon and the North American climate. The wave train extended into the Southern Hemisphere in the non-ENSO and ENSO developing years implies a teleconnection between the WNP summer monsoon and the Australian winter climate. The anomalous WNP monsoon in the non-ENSO and ENSO developing years exists only in summer, while the anomalous WNP monsoon in the ENSO decaying year persists from the beginning of the year to the summer season. The anomalous WNP summer monsoon exhibits a strong ocean–atmosphere interaction, especially in the ENSO decaying year. This study suggests that the anomalous WNP summer monsoon in the non-ENSO year is associated with the variation of the meridional temperature gradient in the upper troposphere, while the anomalous WNP summer monsoon in the ENSO developing and decaying years is associated with ENSO-related SST anomalies.

1. Introduction

The summer monsoon system in east Asia (EA) and the western North Pacific (WNP) are parts of the broadscale Asian summer monsoon system, but with very distinguished features of their own (Lau et al. 2000; Wang and Fan 1999; Wang et al. 2001; Wang and LinHo 2002). The EA summer monsoon over central and northern China, Korea, and Japan has been discussed in numerous studies (e.g., Chang et al. 1998; Chen et al. 1992; Huang and Sun 1992; Lau 1992; Nitta 1987; Tao and Chen 1987; Weng et al. 1999; Zhang et al. 1996). The characteristics of the WNP summer monsoon were first discussed by Tanaka (1997) and Wu and Wang (2000), and a monsoon index was defined by Wang and Fan (1999) to represent the variations of the WNP summer monsoon over the vicinity of the Philippine Sea. The variations of the WNP summer monsoon tend to be out of phase with the EA summer monsoon associated with the mei-yu/baiu front (Chen et al. 2000; Lau et al. 2000; Wang et al. 2001).

The WNP and EA summer monsoons show a strong signal of interannual variability with a biennial cycle (e.g., Clarke et al. 1998; Shen and Lau 1995; Tian and Yasunari 1992; Wang et al. 2001), which is similar to the tropospheric biennial oscillation involved with the Asian–Australian summer monsoon (e.g., Barnett 1991; Chang and Li 2000; Meehl 1994a, 1997; Nicholls 1978; Yasunari 1990). The tropical Pacific sea surface temperature (SST) associated with the WNP and EA summer monsoons also shows a biennial signal (Chang et al. 2000a; Wang et al. 2001). A strong WNP summer monsoon (weak EA summer monsoon) tends to be led by cold tropical eastern Pacific SST anomalies and followed by warm tropical eastern Pacific SST anomalies. A coupled land–atmosphere–ocean interaction involved in midlatitude circulation and the shift of convection–SST anomalies has been proposed to explain the mechanism of the tropospheric biennial oscillation (Meehl 1997). During a strong Asian summer monsoon season, the corresponding strong convection reduces the local SST in the Indian Ocean and enhances the SST over the western Pacific through air–sea interaction. When the convective heating moves eastward to the warm SST over the western Pacific in fall, the corresponding anomalous trough and the persistence of the cold and wet landmass enhance snow cover in winter, which keeps the surface temperature over the Asian continent cold in spring. The cold surface temperature induces a weak meridional tropospheric temperature gradient that is associated with a weak summer monsoon (Chou 2003, hereafter CHOU; Li and Yanai 1996; Meehl 1994b).

The interannual variability of the WNP and EA summer monsoons often links to the El Niño–Southern Oscillation (ENSO; e.g., Chang et al. 2000a,b; Chen et al. 1992; Huang and Sun 1992; Lau and Bua 1998; Lau et al. 2000; Lau and Wu 2001; Nitta 1987; Wang et al. 2001; Weng et al. 1999; Wu and Wang 2000). In a mature phase of El Niño (La Niña), a low-level anticyclonic (cyclonic) circulation over the subtropical WNP is associated with a Rossby wave in response to the suppressed (enhanced) convection in the western Pacific, which is induced by both the local SST cooling (warming) and the remote forcing associated with the weakened (strengthened) Walker circulation (Su et al. 2001; Wang et al. 2000). The anticyclonic (cyclonic) circulation can persist through the following spring and early summer by coupling with the ocean and causes a weak (strong) WNP summer monsoon (Wang et al. 2000). Kawamura et al. (2001) emphasized the importance of the wind–evaporation–SST feedback associated with the anticyclonic (cyclonic) circulation in the Asian summer monsoon and ENSO coupling. The anticyclonic circulation also blocks the mei-yu/baiu front from moving southward and thus enhances the EA summer monsoon rainfall (Chang et al. 2000a). Therefore, a weak (strong) WNP summer monsoon and a strong (weak) EA summer monsoon are often preceded by El Niño (La Niña) in the previous winter (e.g., Wang et al. 2001).

Besides the strong interannual variability in those years of ENSO in the preceding winter, the WNP and EA summer monsoons also exhibit significant interannual variability in those years without ENSO in the preceding winter. This interannual variability might be a part of the tropospheric biennial oscillation or an independent event. For instance, an extremely cold and wet summer associated with a strong EA summer monsoon and a warm and dry summer associated with a weak EA summer monsoon are found in 1993 and 1994, respectively (Geng et al. 2000; Park and Schubert 1997). Thus, it is important to isolate the structure and effects of the WNP summer monsoon by identifying the years other than the mature phases of ENSO, and compare the atmospheric and sea surface anomalies between the ENSO and non-ENSO years. This study focuses on understanding the spatial and temporal patterns of the interannual variability of the WNP summer monsoon with and without the ENSO influence. The definitions of the strong and weak WNP summer monsoon are described in section 2, along with a description of the data used in this study. The relationship between the WNP summer monsoon and ENSO are discussed in section 3, so is the reason for dividing the data into non-ENSO, ENSO developing, and ENSO decaying years. This is followed by discussions on the spatial and temporal patterns of the interannual variability of the WNP summer monsoon. The conclusions are in section 6.

2. Data and a monsoon index for the western North Pacific monsoon

The monthly air temperature, surface pressure, and winds used in the study are derived from the 50-yr (1951–2000) National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis (Kalnay et al. 1996) with a resolution of 2.5° × 2.5°. Monthly SST data are from NCEP using the empirical orthogonal function (EOF) reconstructed SST analysis of Smith et al. (1996) with a resolution of 2° × 2°. Monthly precipitation data are from January 1979 to December 2000 derived by the Climate Prediction Center (CPC) Merged Analysis of Precipitation (CMAP; Xie and Arkin 1997) with a resolution of 2.5° × 2.5°. Note that the period for the precipitation, 1979–2000, is inconsistent with the period of 1951–2000 for other fields, such as the SST. This discrepancy of the period for different data might induce caveats on the interdecadal variation of the Asian summer monsoon–ENSO relation (e.g., Chang et al. 2000a,b; Torrence and Webster 1999; Wang et al. 2001; Weng et al. 1999). However, the comparisons of the nonprecipitation fields, such as the SST and winds, between the period of 1979–2000 and the period of 1951–2000 show that the features of the anomalous WNP summer monsoon, discussed in the next three sections, are robust. The CMAP precipitation associated with the anomalous WNP summer monsoon from the period of 1979–2000 should be compatible with the fields, such as the SST, from the period of 1951–2000. A dynamical monsoon index defined by Wang et al. (2001) is used to measure the variability of the WNP monsoon. The WNP monsoon index similar to Wang and Fan (1999) is defined as the difference of 850-hPa westerly anomalies between the region (5°–15°N, 100°–130°E) and the region (20°–30°N, 110°–140°E). The WNP monsoon with a positive WNP monsoon index averaged over summer is defined as a strong summer monsoon, and vice versa for a weak WNP summer monsoon with a negative WNP monsoon index over the same period. The strength of the WNP summer monsoon for each year is shown in Table 1.

Table 1.

Classification of the data. The definition of the classification is discussed in the text. The * denotes either a year with both ENSO developing and decaying phases in the same calendar year or a year with ENSO persisting from the previous year to the following year

Classification of the data. The definition of the classification is discussed in the text. The * denotes either a year with both ENSO developing and decaying phases in the same calendar year or a year with ENSO persisting from the previous year to the following year
Classification of the data. The definition of the classification is discussed in the text. The * denotes either a year with both ENSO developing and decaying phases in the same calendar year or a year with ENSO persisting from the previous year to the following year

To determine an ENSO event, numerous attempts have been made (e.g., Larkin and Harrison 2001; Neelin et al. 2000; Trenberth 1997; Wang et al. 2000; Xu and Chan 2001, and references therein). However, there is no consensus on the definition of ENSO because of the complexity of the evolution of ENSO (Wallace et al. 1998). ENSO tends to have a phase-locking behavior (e.g., Larkin and Harrison 2002; Wang et al. 2000) that the mature phase of ENSO tends to occur toward the end of the calendar year, but the ENSO phase-locking behavior is more complicated and scattered (Neelin et al. 2000). In this study, we focus on the relation of the WNP summer monsoon with a typical ENSO event that has the phase-locking behavior with the mature phase in winter. The Niño-3.4 (5°N–5°S, 170°–120°W) SST anomalies are used as an ENSO index. A mature phase of the warm event of ENSO (El Niño) is defined when the ENSO index averaged over winter is larger than 0.9°C and a mature phase of the cold event of ENSO (La Niña) is defined when the ENSO index averaged over the same period is less than −0.9°C. An ENSO developing year (Year 0) is defined as the year before the mature phase of ENSO and an ENSO decaying year (Year 1) is defined as the year after the mature phase of ENSO. The labels of years follow the notation of Rasmusson and Carpenter (1982). The summer season is June–August (JJA), and the winter season is December–February (DJF). All ENSO events defined in this study are included within the definition of Trenberth (1997), but not all ENSO events defined by Trenberth (1997) are included within this definition. Compared to another definition of ENSO (Larkin and Harison 2002), three ENSO periods (1984/85, 1994/95, and 2000) defined in this study are not included within their definition. The discrepancy implies caveats on the discussion of the ENSO–monsoon relation. Nevertheless, the overall conclusions are fairly robust based on the sensitivity tests (not shown).

3. Relationships between the western North Pacific summer monsoon and ENSO

The WNP summer monsoon has a strong relation with the tropical eastern Pacific SST. Figure 1a shows the lag correlation of the seasonal Niño-3.4 SST anomalies to the summer WNP monsoon index in the period of 1951–2000. A clear biennial signal is found and this biennial signal is consistent with the finding of Wang et al. (2001). A strong WNP summer monsoon tends to be preceded by cold SST anomalies over the tropical eastern Pacific, and also tends to be followed by warm SST anomalies over the same region. Both the cold and warm SST anomalies that are preceded and followed by the strong WNP summer monsoon can persist for at least three seasons. However, further examinations are needed to determine whether the corresponding warm (cold) SST anomalies, shown in Fig. 1a, are a warm (cold) event of ENSO. First, we divide the data into three categories: ENSO developing years (Year 0), ENSO decaying years (Year 1), and non-ENSO years. Four out of the 18 ENSO-related years are either a year with both developing and decaying phases of ENSO in the same calendar year (1973 and 1998) or a year with ENSO persisting from the previous year to the following year (1955 and 1999). In order to have distinct ENSO developing and decaying years, these four years are excluded. The non-ENSO year is defined as the year that is neither an ENSO developing year nor an ENSO decaying year. Note that in the category of the non-ENSO years, “Year 0” denotes the reference year and “Year 1 (−1)” denotes the year after (before) the reference year. Table 1 shows the classification for each year.

Fig. 1.

Lag correlation of the seasonal Niño-3.4 index with respect to the summer WNP monsoon index for (a) the total period (1951–2000), (b) the ENSO developing years, (c) the ENSO decaying years, and (d) the non-ENSO years. The dashed lines indicate the Pearson's t test at 5% significance level

Fig. 1.

Lag correlation of the seasonal Niño-3.4 index with respect to the summer WNP monsoon index for (a) the total period (1951–2000), (b) the ENSO developing years, (c) the ENSO decaying years, and (d) the non-ENSO years. The dashed lines indicate the Pearson's t test at 5% significance level

In the ENSO developing year, a biennial signal of the lag correlation of the seasonal Niño-3.4 SST anomalies to the summer WNP monsoon index similar to Fig. 1a is found in Fig. 1b. The biennial signal indicates that a strong (weak) WNP summer monsoon tends to occur during the developing year (Year 0) of El Niño (La Niña) and weaker-than-normal ENSO cold (warm) SST anomalies over the tropical eastern Pacific tend to occur in the previous year (Year −1). Note that the ENSO SST anomalies over the Niño-3.4 region persist one season longer than the weaker-than-normal ENSO SST anomalies. In the 14 ENSO developing years, the positive correlation of the ENSO SST anomalies over the Niño-3.4 region to the summer WNP monsoon index is found in 11 yr except for 1957, 1968, and 1984. However, only 8 yr (1954, 1965, 1970, 1972, 1984, 1986, 1988, and 1997) are consistent with the negative correlation of the weaker-than-normal ENSO Niño-3.4 SST anomalies to the summer WNP monsoon index in the ENSO developing year (Year 0). The WNP summer monsoon has a statistically significant instantaneous positive correlation (0.63) with the Niño-3.4 SST anomalies in JJA (Year 0). The similar biennial signal is not found in Figs. 1c and 1d. Figure 1c shows that a strong WNP summer monsoon tends to occur during the decaying year of La Niña. In the 14 ENSO decaying years, only 3 yr (1956, 1971, and 1992) are the exception. Unlike in the ENSO developing year, there is no statistically significant instantaneous correlation of the Niño-3.4 SST anomalies in JJA (Year 1) to the summer WNP monsoon index in the ENSO decaying year. During the non-ENSO year (Fig. 1d), no significant relation between the WNP summer monsoon and the tropical eastern Pacific SST anomalies is found. Thus, the relation between the WNP summer monsoon and the winter Niño-3.4 SST anomalies is good only when ENSO occurs. Based on the results of Figs. 1b–d, the biennial signal found in Fig. 1a does not indicate consecutive warm and cold events of ENSO except for 1973 and 1998 (the El Niño decaying phase and the La Niña developing phase in the same calendar year, see Table 1). The biennial signal is rather a combination of separate warm and cold events of ENSO.

Figure 1 indicates that the analysis that uses the entire data without classifying them cannot simply represent the complete features of the interannual variability of the WNP summer monsoon. Thus, it is necessary to divide the data into ENSO developing years, ENSO decaying years, and non-ENSO years when studying the characteristics of the interannual variability of the WNP summer monsoon. The following two sections will focus on the spatial and temporal structures of the interannual variability of the WNP summer monsoon in these three categories, respectively.

4. Spatial structures of the interannual variability of the western North Pacific summer monsoon

Figure 2 shows the composite differences of summer rainfall between strong and weak WNP summer monsoons in the non-ENSO years (Fig. 2a), the ENSO developing years (Fig. 2b), and the ENSO decaying years (Fig. 2c). According to Fig. 2a, which is the case without any ENSO influence, the positive rainfall anomalies over the region of 5°–25°N and 110°E–180° correspond to the strong WNP summer monsoon defined by 850-hPa zonal winds. The rainfall anomalies associated with the anomalous WNP summer monsoon in Fig. 2a are modified during the ENSO developing and decaying years. No statistically significant positive rainfall anomalies are found over the subtropical WNP in the ENSO developing year, but the rainfall anomalies over this region do show a positive tendency. The positive rainfall anomalies associated with the horseshoe pattern of the ENSO rainfall anomalies (Wallace et al. 1998) are located over the subtropical WNP in the ENSO decaying year. Comparing the ENSO decaying and non-ENSO years, in which statistically significant rainfall anomalies are found over the subtropical WNP, some slight differences of the WNP summer monsoon rainfall anomalies might be important for the local climate. For instance, the region of the maximum rainfall anomalies in the ENSO decaying year is a little more eastward than in the non-ENSO year. The summer precipitation associated with the strong WNP summer monsoon over the South China Sea and Taiwan is less in the ENSO decaying year than in the non-ENSO year. Examining the rainfall anomalies over the whole Pacific, the results in Figs. 2b and 2c are consistent with the results in Fig. 1: the strong WNP summer monsoon in the ENSO developing year is associated with El Niño, while the strong WNP summer monsoon in the ENSO decaying year is associated with La Niña. During the strong WNP summer monsoon, relatively weak negative rainfall anomalies associated with the weak EA summer monsoon are found over central and northern China and extend to Korea and Japan in the non-ENSO, ENSO developing, and ENSO decaying years (Fig. 2). It indicates an out-of-phase interannual variation of the rainfall anomalies between the WNP and EA summer monsoons, in which the latter is associated with the mei-yu/baiu front (Wang et al. 2001). Note that the relatively weak signal of the EA summer monsoon rainfall anomalies is due to the climatologically weak rainfall amount over this region. The signal of the EA summer monsoon rainfall anomalies is enhanced when using the correlation of the summer monsoon rainfall anomalies to the summer WNP monsoon index (not shown).

Fig. 2.

Composite difference of the summer rainfall anomalies between the strong and weak WNP summer monsoons for (a) the non-ENSO years, (b) the ENSO developing years, and (c) the ENSO decaying years. The contour interval is 1 mm day−1. Areas with significance level at 5% by the two-sample t test are shaded

Fig. 2.

Composite difference of the summer rainfall anomalies between the strong and weak WNP summer monsoons for (a) the non-ENSO years, (b) the ENSO developing years, and (c) the ENSO decaying years. The contour interval is 1 mm day−1. Areas with significance level at 5% by the two-sample t test are shaded

The SST anomalies associated with the strong WNP summer monsoon in the non-ENSO, ENSO developing, and ENSO decaying years are shown in Fig. 3. In the non-ENSO year, a dipole pattern of the SST anomalies is found over the WNP along the east coast of the Asian continent (Fig. 3a). During the strong WNP summer monsoon, negative SST anomalies centered at the South China Sea extend westward to the Bay of Bengal. The corresponding positive SST anomalies with stronger magnitude are around 40°N and between 120°E–180°. The dipole pattern of the SST anomalies over the WNP roughly parallels the dipole pattern of the WNP and EA summer monsoon rainfall anomalies shown in Fig. 2a, but with a reversed sign of the anomalies. The weaker negative SST anomalies associate with the stronger positive rainfall anomalies in the south and the stronger positive SST anomalies are related to the weaker negative rainfall anomalies in the north. The alignment of the SST and rainfall anomalies implies strong coupling between the atmosphere and the ocean. Over the region with positive rainfall anomalies, strong evaporation is associated with the strong WNP summer monsoon and solar radiation that reaches the ocean surface is reduced by the increase of cloud amount, so the SST becomes colder. Over the region with negative rainfall anomalies, the weaker evaporation and the increase of solar radiation at the ocean surface due to less clouds warm up the local SST (Lau et al. 2000; Wang et al. 2001). The strong WNP summer monsoon over the negative SST anomalies cannot be simply simulated by an atmospheric model with the prescribed SST anomalies, so a coupled ocean–atmosphere model is needed for simulating the correct WNP summer monsoon.

Fig. 3.

As in Fig. 2, except for the summer SST anomalies. The contour interval is (a) 0.1°C and (b),(c) 0.2°C. Areas with significant levels at 5% are shaded

Fig. 3.

As in Fig. 2, except for the summer SST anomalies. The contour interval is (a) 0.1°C and (b),(c) 0.2°C. Areas with significant levels at 5% are shaded

In the ENSO developing year, a typical SST anomaly pattern associated with El Niño is found in Fig. 3b. Strong warm SST anomalies over the central and eastern Pacific are flanked by relatively weak cold SST anomalies (Su et al. 2001; Wallace et al. 1998). Focusing on the WNP, a similar dipole pattern of the SST anomalies are also found. During the strong WNP summer monsoon, the southern part of the dipole with the cold SST anomalies is located farther eastward in the ENSO developing year than in the non-ENSO year. These cold SST anomalies are associated with the northern branch of the horseshoe pattern of the El Niño-associated SST anomalies. The northern part of the dipole with the warm SST anomalies is much weaker in the ENSO developing year than in the non-ENSO year. In the ENSO decaying year, the cold SST anomalies associated with the southern part of the dipole pattern are found over the South China Sea during the strong WNP summer monsoon, but the associated warm SST anomalies are found only over the Sea of Japan (Fig. 3c). Concluding from Fig. 3, the cold SST anomalies over the South China Sea are a common feature for all the cases of strong WNP summer monsoon. The tropical eastern Pacific SST anomalies in the non-ENSO, ENSO developing, and ENSO decaying years are significantly different, so the SST anomalies over the tropical eastern Pacific only play a modest role in causing the anomalous WNP summer monsoon.

Figure 4 shows the upper-tropospheric temperature anomalies averaged over 200–500 hPa. In the non-ENSO year, strong upper-tropospheric warming is found over the Asian continent between 25° and 50°N. The enhanced meridional gradient of the upper-tropospheric temperature over Asia favors a strong WNP summer monsoon (CHOU; Li and Yanai 1996). The positive upper-tropospheric temperature anomalies extend to the northwestern part of North America, evidence of the teleconnection between the WNP summer monsoon and the North American summer climate. This teleconnection has been discussed by Lau et al. (2000) and Wang et al. (2001), in which a strong (weak) WNP summer monsoon tends to associate with positive (negative) rainfall anomalies over the U.S. Great Plains. When ENSO occurs, the upper-tropospheric temperature anomalies change signs. In the ENSO developing year, the cold upper-tropospheric temperature anomalies dominate the eastern part of Asia during the strong WNP summer monsoon, so the meridional upper-tropospheric temperature gradient is weakened and disfavors a strong WNP summer monsoon. Thus, the meridional upper-tropospheric temperature gradient is no longer a forcing for the anomalous WNP summer monsoon in the ENSO developing year. In the ENSO decaying year, the upper troposphere in the Tropics is dominated by the typical La Niña–induced cold temperature anomalies, as described by Wallace et al. (1998), during the strong WNP summer monsoon. The meridional upper-tropospheric temperature gradient averaged over Asia varies little (not shown), so the meridional upper-tropospheric temperature gradient is also not a forcing factor for the anomalous WNP summer monsoon in the ENSO decaying year. In conclusion, the meridional upper-tropospheric temperature gradient could be a forcing for the anomalous WNP summer monsoon only in the non-ENSO years. In the ENSO developing and decaying years, there must be different forcings that induce the anomalous WNP summer monsoon.

Fig. 4.

As in Fig. 2, except for the summer upper-tropospheric (200–500 hPa) temperature anomalies. The contour interval is 0.2°C. Areas with significant levels at 5% are shaded

Fig. 4.

As in Fig. 2, except for the summer upper-tropospheric (200–500 hPa) temperature anomalies. The contour interval is 0.2°C. Areas with significant levels at 5% are shaded

Corresponding to the convective heating of the strong WNP summer monsoon, positive upper-tropospheric temperature anomalies are found over the subtropical WNP in the non-ENSO and ENSO developing years. The positive upper-tropospheric temperature anomalies in the ENSO developing year provide further evidence for the tendency of the positive rainfall anomalies over the subtropical WNP as shown in Fig. 2b. In the ENSO decaying year, there are no statistically significant positive upper-tropospheric temperature anomalies over the subtropical WNP since the Tropics is dominated by the La Niña–induced cold temperature anomalies (Wallace et al. 1998). However, the tendency of the weaker cold upper-tropospheric temperature anomalies over the subtropical WNP implies that the convective heating associated with the strong WNP summer monsoon does have an impact on this region. More positive upper-tropospheric temperature anomalies are found over the Australian continent in the non-ENSO and ENSO developing years with the stronger anomalies in the non-ENSO year. This implies a possible teleconnection between the WNP summer monsoon and the Australian winter climate.

In Fig. 5, low surface pressure anomalies and a cyclonic circulation at 850 hPa associated with the strong WNP summer monsoon are found in all of the non-ENSO, ENSO developing, and ENSO decaying years. This implies the weakening of the WNP subtropical ridge that is associated with a weak EA summer monsoon (Chang et al. 2000a). Beginning with the low surface pressure anomalies in the subtropical WNP, a wave train of the surface pressure anomalies along the east coast of Asia is found in all three categories with a slight difference and extends to North America. Besides the low surface pressure anomalies in the subtropical WNP, the wave train includes the high surface pressure anomalies over Korea and Japan, the low surface pressure anomalies from northeastern Asia to the Bering Sea, and the high surface pressure anomalies over the Gulf of Alaska and northern North America. The 850-hPa wind anomalies are consistent with the surface pressure anomalies, with a wave train extending from the subtropical WNP to North America. The high surface pressure anomalies and the anticyclonic circulation over Korea and Japan are associated with a weak EA summer monsoon. The southern branch of the cyclonic circulation (westerlies) over the Philippine Sea enhances the local evaporation and brings moist air from the South China Sea and the Indian Ocean; thus, it has a positive feedback on the WNP summer monsoon. The southern branch of the anticyclonic circulation (easterlies) over Korea and Japan reduces the local evaporation and transports dry air into the region; thus, it induces a weak EA summer monsoon (Lau et al. 2000).

Fig. 5.

As in Fig. 2, except for the summer surface pressure anomalies and 850-hPa wind anomalies (m s−1). The contour interval is 0.5 hPa and areas with significant levels for the surface pressure anomalies at 5% are shaded

Fig. 5.

As in Fig. 2, except for the summer surface pressure anomalies and 850-hPa wind anomalies (m s−1). The contour interval is 0.5 hPa and areas with significant levels for the surface pressure anomalies at 5% are shaded

In the non-ENSO year, the low surface pressure anomalies over the subtropical WNP associated with the strong WNP summer monsoon tend to link to the low surface pressure anomalies over the Asian continent, which extend to the Bering Sea (Fig. 5a). The high surface pressure anomalies over Korea and Japan tend to separate from the system of the surface pressure over the Asian continent. The low-level cyclonic circulation over the Asian continent, associated with the low surface pressure anomalies over the same region, and the subtropical WNP cyclonic circulation, are consistent with an enhanced monsoon circulation that is induced by the strong meridional temperature gradient, as shown in Fig. 4a (CHOU). In the ENSO developing year, the low surface pressure anomalies associated with the strong WNP summer monsoon are an extension of the low surface pressure anomalies over the central and eastern Pacific. The southern branch of the subtropical WNP cyclonic circulation merges with the anomalous westerlies over the equatorial Pacific that respond to the warm Niño-3.4 SST anomalies. Both provide evidence of the association between the WNP summer monsoon and ENSO. In contrast to the surface pressure anomalies over the Asian continent in the non-ENSO year, the high surface pressure anomalies dominate the Asian continent in the ENSO developing and decaying years. Thus, the low surface pressure anomalies associated with the strong WNP summer monsoon tend to separate from the system of the surface pressure over the Asian continent. It implies that the strong WNP summer monsoon in the ENSO developing and decaying years is not associated with the variations of the system over the Asian continent.

To examine the vertical structure of the anomalous WNP summer monsoon, the zonal component of the 850-hPa wind anomalies averaged over 120°–160°E is shown in Fig. 6. A wave train over the Northern Hemisphere is found for all three categories but is meridionally wider in the ENSO decaying year. During the strong WNP summer monsoon, the low-level westerlies in the Tropics exhibit a baroclinic structure with easterlies on top. The anomalous tropical westerlies in the ENSO developing year extend to the higher troposphere, while the anomalous tropical westerlies in the non-ENSO and ENSO decaying years are confined to the lower troposphere. The zonal wind anomalies at higher latitudes have the same sign in the vertical and these barotropic responses to the tropical heating have been discussed in numerous studies (e.g., Hoskins and Karoly 1981; Lim and Chang 1986; Webster 1981). A sign of the wave train is also found in the Southern Hemisphere in the non-ENSO and ENSO developing years, which is consistent with the upper-tropospheric temperature anomalies over the Australian continent, as shown in Figs. 4a and 4b. Note that the wave train amplitude in the Southern Hemisphere is mostly confined to the upper troposphere.

Fig. 6.

As in Fig. 2, except for the summer zonal wind anomalies averaged over 120°–160°E. The contour interval is 0.5 m s−1. Areas with significant levels at 5% are shaded

Fig. 6.

As in Fig. 2, except for the summer zonal wind anomalies averaged over 120°–160°E. The contour interval is 0.5 m s−1. Areas with significant levels at 5% are shaded

5. Temporal structures of the interannual variability of the western North Pacific summer monsoon

To study the temporal variation of the anomalous WNP monsoon, the variables associated with the anomalous WNP monsoon are averaged over 120°–160°E, a region dominated by the WNP monsoon and correlated with the summer WNP monsoon index. Bearing in mind the caveat that not all anomalies occur exactly between 120° and 160°E, the averaged anomalies over 120°–160°E might not be statistically significant to identify the temporal variations of the anomalies shown in Figs. 25. Nevertheless, the tendency of the temporal variations of the features associated with the anomalous WNP monsoon found in Figs. 710 is good indication based on the sensitivity tests with varied regions for the different fields and categories (not shown). Figure 7 shows the lag correlation of the monthly rainfall anomalies to the summer WNP monsoon index. In the non-ENSO year, the positive rainfall anomalies associated with the strong WNP summer monsoon (Fig. 2a) persist for three months (JJA). No statistically significant signal is found during northern summer in the ENSO developing year, as expected according to Fig. 2b. Two negative rainfall anomalies associated with the negative correlation (Fig. 7b) at 20° on both sides of the equator from August to December correspond to a horseshoe pattern of the rainfall anomalies of a typical El Niño (Su et al. 2001; Wallace et al. 1998). In the ENSO decaying year, the positive rainfall anomalies associated with the strong WNP summer monsoon exist not only in summer, but they tend to persist from the previous winter when La Niña is mature. This is consistent with the findings of Wang et al. (2000). In contrast to the rainfall anomalies of the WNP monsoon, the corresponding EA monsoon rainfall anomalies do not have a statistically significant signal in the non-ENSO, ENSO developing, and ENSO decaying years. It might be due to the smaller region of the corresponding EA monsoon rainfall anomalies.

Fig. 7.

Lag correlation of the monthly rainfall anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index (Wang et al. 2001) for (a) the non-ENSO years, (b) the ENSO developing years, and (c) the ENSO decaying years. The contour interval is 0.2. Areas with significant levels for the Pearson's t test at 5% are shaded

Fig. 7.

Lag correlation of the monthly rainfall anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index (Wang et al. 2001) for (a) the non-ENSO years, (b) the ENSO developing years, and (c) the ENSO decaying years. The contour interval is 0.2. Areas with significant levels for the Pearson's t test at 5% are shaded

Fig. 10.

As in Fig. 7, except for lag correlation of the monthly 850-hPa zonal wind anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index. The contour interval is 0.1. Areas with significant levels at 5% are shaded

Fig. 10.

As in Fig. 7, except for lag correlation of the monthly 850-hPa zonal wind anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index. The contour interval is 0.1. Areas with significant levels at 5% are shaded

Figure 8 shows the lag correlation of the averaged SST anomalies to the summer WNP monsoon index. The dipole pattern of the SST anomalies over the WNP in the non-ENSO year shown in Fig. 3a is not statistically significant in the averaged SST anomalies (Fig. 8a) because the cold SST anomalies associated with the strong WNP summer monsoon are confined to the east of 140°E. Several sensitivity tests have been done using different regions to represent the temporal variation of the cold SST anomalies over the subtropical WNP. The results show that the warm SST anomalies between 30° and 45°N shown in Fig. 8a, the northern part of the dipole, can well represent the temporal variation of the SST anomalies associated with the southern part of the dipole. According to the positive correlation between 30° and 45°N, the SST anomalies associated with the anomalous WNP summer monsoon occur only in summer, same as the rainfall anomalies (Fig. 7a). In the ENSO developing year, the cold SST anomalies associated with the strong WNP summer monsoon are more profound and can persist from May to December. No statistically significant positive correlation is found over the north of the cold SST anomalies, due to the weak SST anomalies associated with the northern part of the SST dipole pattern (Fig. 3b). The cold SST anomalies over the subtropical WNP in the ENSO developing year are associated with the northern branch of the horseshoe pattern of the El Niño–related SST anomalies (Fig. 3b). In the ENSO decaying year, the pattern of the SST anomalies over the subtropical WNP is very interesting. The cold SST anomalies associated with the strong WNP summer monsoon are found between 20° and 30°N during July–October, while warm SST anomalies are found between 0° and 20°N during January–April. Over the region of the cold and warm SST anomalies from January to October, the corresponding rainfall anomalies stay positive (Fig. 7c) no matter what the sign of the SST anomalies. The positive WNP monsoon rainfall anomalies with the change of sign for the SST anomalies beneath imply two different mechanisms involved in the complicated ocean–atmosphere interaction. When ENSO is mature, the SST anomalies over the western Pacific associate with ENSO and force the anomalous WNP monsoon circulation (Su et al. 2001; Wang et al. 2000). When ENSO is decaying, the SST anomalies over the western Pacific are weakening and the local surface heat flux anomalies induced by the anomalous WNP summer monsoon become dominant in determining the SST anomalies over the western Pacific (Lau et al. 2000; Wang et al. 2001). How the strong WNP monsoon with positive rainfall anomalies persists from the winter into the summer when the forcing of the warm SST anomalies over the western Pacific associated with La Niña disappears is an interesting question and should be further investigated.

Fig. 8.

As in Fig. 7, except for lag correlation of the monthly SST anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index. The contour interval is 0.1. Areas with significant levels at 5% are shaded

Fig. 8.

As in Fig. 7, except for lag correlation of the monthly SST anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index. The contour interval is 0.1. Areas with significant levels at 5% are shaded

Figure 9 shows the lag correlation of the surface pressure anomalies averaged over 120°–160°E to the summer WNP monsoon index. During the strong WNP summer monsoon, the corresponding low surface pressure anomalies between 10° and 20°N in the non-ENSO and ENSO developing years can only exist in summer (Figs. 9a and 9b). In the ENSO decaying year, the corresponding low surface pressure anomalies persist from the previous winter when La Niña is mature. In the mature phase of La Niña, the western Pacific is dominated by the low surface pressure anomalies (Wallace et al. 1998). When La Niña is decaying, the region dominated by the low surface pressure anomalies becomes more meridionally confined to the subtropical WNP. In the 850-hPa wind fields (Fig. 10), the cyclonic circulation over the subtropical WNP associated with the strong WNP summer monsoon also only exists in summer in the non-ENSO and ENSO developing years, as does the associated wave train. In the ENSO decaying year, the corresponding cyclonic circulation tends to persist from January to August (Fig. 10c). Similar to the surface pressure anomalies in Fig. 9c, the persistence of the cyclonic circulation is also evidence that ENSO may affect the WNP summer monsoon in the following summer.

Fig. 9.

As in Fig. 7, except for lag correlation of the monthly surface pressure anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index. The contour interval is 0.1. Areas with significant levels at 5% are shaded

Fig. 9.

As in Fig. 7, except for lag correlation of the monthly surface pressure anomalies averaged over 120°–160°E with respect to the summer WNP monsoon index. The contour interval is 0.1. Areas with significant levels at 5% are shaded

6. Discussion and conclusions

The WNP summer monsoon has a profound relationship with ENSO, while the WNP summer monsoon also exhibits an interannual variability of its own. A strong (weak) WNP summer monsoon tends to occur in the El Niño (La Niña) developing year and the La Niña (El Niño) decaying year (Figs. 1b and 1c). Without the ENSO influence, the interannual variability of the WNP summer monsoon does not link to the tropical eastern Pacific SST (Fig. 1d). Because of the distinctly different interannual variability of the WNP summer monsoon with and without ENSO influence, a 50-yr NCEP–NCAR reanalysis data including SST and a 22-yr CMAP precipitation data are divided into three categories: non-ENSO, ENSO developing, and ENSO decaying years to examine the characteristics of the interannual variability of the WNP summer monsoon.

The spatial pattern of the interannual variability of a typical strong (weak) WNP summer monsoon in the non-ENSO, ENSO developing, and ENSO decaying years includes the positive (negative) rainfall anomalies, cold (warm) SST anomalies, warm (cold) upper-tropospheric temperature anomalies, low (high) surface pressure anomalies, and a low-level cyclonic (anticyclonic) circulation over the subtropical WNP. The convective heating over the subtropical WNP associated with the strong WNP summer monsoon cools down the SST beneath by reducing downward solar radiation and enhancing evaporation. The convective heating also warms up the upper troposphere and induces low surface pressure anomalies and a low-level cyclonic circulation anomaly. A wave train including the WNP and EA summer monsoons is also found along the east coast of Asia, in which the WNP summer monsoon tends to have a negative correlation with the EA summer monsoon. The wave train extends to North America, which implies a teleconnection between the WNP summer monsoon and the North American summer climate (Lau et al. 2000; Wang et al. 2001). The vertical structure of the wave train is barotropic except for the southern branch of the anomalous WNP summer monsoon circulation.

By comparing the results shown in Figs. 26, the differences of the interannual variability of the anomalous WNP summer monsoon with and without ENSO influence are established. The positive rainfall anomalies associated with the strong WNP summer monsoon in the ENSO developing year are not as statistically significant as in the non-ENSO year. The corresponding rainfall anomalies over the subtropical WNP in the ENSO decaying year are associated with the northern branch of the horseshoe rainfall anomalies that relate to La Niña. The cold SST anomalies associated with the strong WNP summer monsoon are stronger and wider in the ENSO developing year than in the non-ENSO and ENSO decaying years. This cold SST anomalies in the ENSO developing year are related to the horseshoe pattern of the El Niño SST anomalies. The warm upper-tropospheric temperature anomalies over Asia during the strong WNP summer monsoon in the non-ENSO year enhance the meridional temperature gradient and favor the strong summer monsoon. In contrast to the non-ENSO year, cold upper-tropospheric temperature anomalies are found over Asia during the strong WNP summer monsoon in the ENSO developing year. They disfavor the strong summer monsoon due to the weakening of the meridional temperature gradient. In the ENSO decaying year, the meridional upper-tropospheric temperature gradient does not vary much, so the temperature gradient is also not a forcing for the strong WNP summer monsoon. Thus, the strong WNP summer monsoon in the ENSO developing and decaying years should be induced by different mechanisms. The low surface pressure anomalies associated with the strong WNP summer monsoon tend to connect with the low surface pressure anomalies over the Asian continent in the non-ENSO year, but are disjointed from the high surface pressure anomalies over the Asian continent in the ENSO developing and decaying years. The change of the sign for the Asian surface pressure anomalies between the non-ENSO year and the ENSO developing and decaying years provides further evidence that there are different forcings for the interannual variability of the WNP summer monsoon in the non-ENSO, ENSO developing, and ENSO decaying years. During the strong WNP summer monsoon in the ENSO developing year, the low surface pressure anomalies over the subtropical WNP are associated with the El Niño–induced low surface pressure anomalies over the central and eastern Pacific (Fig. 5b). The wave train of the zonal wind anomalies is found at both hemispheres in the non-ENSO and ENSO developing years, but the wave train in the ENSO decaying year is confined to the Northern Hemisphere.

The temporal variation of the interannual variability of the WNP monsoon has also been studied. The anomalous WNP monsoon in the non-ENSO and ENSO developing years occurs only in summer, while the anomalous WNP monsoon in the ENSO decaying year persists from the beginning of the year, when ENSO is mature, to the summer. The persistence of the anomalous WNP monsoon in the ENSO decaying year provides evidence that ENSO affects the WNP summer monsoon in the following year (Wang et al. 2000). However, the mechanism of how ENSO affects the interannual variability of the following WNP summer monsoon involves a complicated ocean–atmosphere interaction since the sign of the SST anomalies over the subtropical WNP changes when ENSO is decaying. The warm (cold) SST anomalies over the western Pacific during the mature phase of La Niña (El Niño) induce positive (negative) rainfall anomalies over the western Pacific (Su et al. 2001) and can last for a few months even when the SST anomalies have changed the sign. In summer, the subtropical WNP is dominated by the cold (warm) SST anomalies that are induced by the strong (weak) evaporation and weak (strong) downward solar radiation (Lau et al. 2000; Wang et al. 2001).

The analysis suggests that the anomalous WNP summer monsoon is related to different processes for the non-ENSO, ENSO developing, and ENSO decaying years. In the non-ENSO year, the variation of the meridional upper-tropospheric temperature gradient might be responsible for the anomalous WNP summer monsoon. In the ENSO developing year, processes related to the eastward extension of the eastern Pacific SST anomalies associated with ENSO could be the reason for the anomalous WNP summer monsoon. In the ENSO decaying year, the western Pacific SST anomalies in the mature phase of ENSO might associate with the anomalous WNP summer monsoon. These processes involving the interannual variability of the WNP summer monsoon in the non-ENSO, ENSO developing, and ENSO decaying years are currently under investigation.

Acknowledgments

The authors thank Drs. C.-T. Chen, LinHo, H.-H. Hsu, Y.-A. Lee, M.-M. Lu, J. D. Neelin, H. Su, C.-H. Sui, and C.-H. Tsou for discussions. The first author also thanks Dr. Hsin-Cheng Huang at the Institute of Statistical Science, Academia Sinica, for help with the statistical methods. Comments from Drs. C.-P. Chang, S. D. Schubert, and two anonymous reviewers were particularly helpful for improving the quality of this paper. This work was supported under National Science Council Grant 90-2119-M-001-020 and the Environmental Change Research Project of Academia Sinica.

REFERENCES

REFERENCES
Barnett
,
T. P.
,
1991
:
The interaction of multiple timescales in the tropical climate system.
J. Climate
,
4
,
269
285
.
Chang
,
C-P.
, and
T.
Li
,
2000
:
A theory for the tropical tropospheric biennial oscillation.
J. Atmos. Sci.
,
57
,
2209
2224
.
Chang
,
C-P.
,
S. C.
Hou
,
H. C.
Kuo
, and
G. T. J.
Chen
,
1998
:
The development of an intense East Asian summer monsoon disturbance with strong vertical coupling.
Mon. Wea. Rev.
,
126
,
2692
2712
.
Chang
,
C-P.
,
Y.
Zhang
, and
T.
Li
,
2000a
:
Interannual and interdecadal variations of the East Asian summer monsoon and tropical Pacific SSTs. Part I: Roles of the subtropical ridge.
J. Climate
,
13
,
4310
4325
.
Chang
,
C-P.
,
Y.
Zhang
, and
T.
Li
,
2000b
:
Interannual and interdecadal variations of the East Asian summer monsoon and tropical Pacific SSTs. Part II: Meridional structure of the monsoon.
J. Climate
,
13
,
4326
4340
.
Chen
,
L.
,
M.
Dong
, and
Y.
Shao
,
1992
:
The characteristics of interannual variations on the East Asian monsoon.
J. Meteor. Soc. Japan
,
70
,
397
421
.
Chen
,
T-C.
,
M-C.
Yen
, and
S-P.
Weng
,
2000
:
Interaction between the summer monsoons in East Asia and the South China Sea: Intraseasonal monsoon modes.
J. Atmos. Sci.
,
57
,
1373
1392
.
Chou
,
C.
,
2003
:
Land–sea heating contrast in an idealized Asian summer monsoon.
Climate Dyn., in press
.
Clarke
,
A. J.
,
X.
Liu
, and
S.
van Gorder
,
1998
:
Dynamics of the biennial oscillation in the equatorial Indian and far western Pacific Oceans.
J. Climate
,
11
,
987
1001
.
Geng
,
Q. Z.
,
A.
Sumi
, and
A.
Numaguti
,
2000
:
Role of transients in the dynamics of East Asian summer seasonal mean circulation anomalies—A study of 1993 and 1994.
J. Climate
,
13
,
3511
3531
.
Hoskins
,
B. J.
, and
D. J.
Karoly
,
1981
:
The steady linear response of a spherical atmosphere to thermal and orographic forcing.
J. Atmos. Sci.
,
38
,
1179
1196
.
Huang
,
R.
, and
F.
Sun
,
1992
:
Impacts of the tropical western Pacific on the East Asian summer monsoon.
J. Meteor. Soc. Japan
,
70
,
243
256
.
Kalnay
,
E.
, and
Coauthors
.
1996
:
The NCEP/NCAR 40-Year Reanalysis Project.
Bull. Amer. Meteor. Soc.
,
77
,
437
471
.
Kawamura
,
R.
,
T.
Matsuura
, and
S.
Iisuka
,
2001
:
Interannual atmosphere–ocean variations in the tropical western North Pacific relevant to the Asian summer monsoon–ENSO coupling.
J. Meteor. Soc. Japan
,
79
,
883
898
.
Larkin
,
N. K.
, and
D. E.
Harrison
,
2001
:
Tropical Pacific ENSO cold events, 1946–95: SST, SLP, and surface wind composite anomalies.
J. Climate
,
14
,
3904
3931
.
Larkin
,
N. K.
, and
D. E.
Harrison
,
2002
:
ENSO warm (El Niño) and cold (La Niña) event life cycles: Ocean surface anomaly patterns, their symmetries, asymmetries, and implications.
J. Climate
,
15
,
1118
1140
.
Lau
,
K-M.
,
1992
:
East Asian summer monsoon rainfall variability and climate teleconnection.
J. Meteor. Soc. Japan
,
70
,
211
241
.
Lau
,
K-M.
, and
W.
Bua
,
1998
:
Mechanism of monsoon–Southern Oscillation coupling: Insights from GCM experiments.
Climate Dyn.
,
14
,
759
779
.
Lau
,
K-M.
, and
H. T.
Wu
,
2001
:
Principal modes of rainfall–SST variability of the Asian summer monsoon: A reassessment of the monsoon–ENSO relationship.
J. Climate
,
14
,
2880
2895
.
Lau
,
K-M.
,
K-M.
Kim
, and
S.
Yang
,
2000
:
Dynamical and boundary forcing characteristics of regional components of the Asian summer monsoon.
J. Climate
,
13
,
2461
2482
.
Li
,
C.
, and
M.
Yanai
,
1996
:
The onset and interannual variability of the Asian summer monsoon in relation to land–sea thermal contrast.
J. Climate
,
9
,
358
375
.
Lim
,
H.
, and
C-P.
Chang
,
1986
:
Generation of internal- and external-mode motions from internal heating: Effects of vertical shear and damping.
J. Atmos. Sci.
,
43
,
948
957
.
Meehl
,
G. A.
,
1994a
:
Coupled land–ocean–atmosphere processes and south Asian monsoon variability.
Science
,
266
,
263
267
.
Meehl
,
G. A.
,
1994b
:
Influence of the land surface in the Asian summer monsoon: External conditions versus internal feedbacks.
J. Climate
,
7
,
1033
1049
.
Meehl
,
G. A.
,
1997
:
The south Asian monsoon and the tropospheric biennial oscillation.
J. Climate
,
10
,
1921
1943
.
Neelin
,
J. D.
,
F-F.
Jin
, and
H-H.
Syu
,
2000
:
Variations in ENSO phase locking.
J. Climate
,
13
,
2570
2590
.
Nicholls
,
N.
,
1978
:
Air–sea interaction and the quasi-biennial oscillation.
Mon. Wea. Rev.
,
106
,
1505
1508
.
Nitta
,
T.
,
1987
:
Convective activities in the tropical western Pacific and their impacts on the Northern Hemisphere summer circulation.
J. Meteor. Soc. Japan
,
65
,
165
171
.
Park
,
C-K.
, and
S. D.
Schubert
,
1997
:
On the nature of the 1994 East Asian summer drought.
J. Climate
,
10
,
1056
1070
.
Rasmusson
,
E. M.
, and
T. H.
Carpenter
,
1982
:
Variations in tropical sea surface temperature and surface wind fields associated with the Southern Oscillation/El Niño.
Mon. Wea. Rev.
,
110
,
354
384
.
Shen
,
S.
, and
K-M.
Lau
,
1995
:
Biennial oscillation associated with the East Asian summer monsoon and tropical sea surface temperatures.
J. Meteor. Soc. Japan
,
73
,
105
124
.
Smith
,
T. M.
,
R. W.
Renolds
,
R. E.
Livezey
, and
D. C.
Stokes
,
1996
:
Reconstruction of historical sea surface temperatures using empirical orthogonal functions.
J. Climate
,
9
,
1403
1420
.
Su
,
H.
,
J. D.
Neelin
, and
C.
Chou
,
2001
:
Tropical teleconnection and local response to SST anomalies during the 1997–1998 El Niño.
J. Geophys. Res.
,
106
,
20025
20043
.
Tanaka
,
M.
,
1997
:
Interannual and interdecadal variations of the western North Pacific monsoon and the East Asian Baiu rainfall and their relationship to ENSO cycles.
J. Meteor. Soc. Japan
,
75
,
1109
1123
.
Tao
,
S.
, and
L.
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
.
Tian
,
S-F.
, and
T.
Yasunari
,
1992
:
Time and space structure of interannual variations in summer rainfall over China.
J. Meteor. Soc. Japan
,
70
,
585
596
.
Torrence
,
C.
, and
P. J.
Webster
,
1999
:
Interdecadal changes in the ENSO–monsoon system.
J. Climate
,
12
,
2679
2710
.
Trenberth
,
K. E.
,
1997
:
The definition of El Niño.
Bull. Amer. Meteor. Soc.
,
78
,
2771
2777
.
Wallace
,
J. M.
,
E. M.
Rasmusson
,
T. P.
Mitchell
,
V. E.
Kousky
,
E. S.
Sarachik
, and
H.
von Storch
,
1998
:
On the structure and evolution of ENSO-related climate variability in the tropical Pacific: Lessons from TOGA.
J. Geophys. Res.
,
103
,
14241
14260
.
Wang
,
B.
, and
Z.
Fan
,
1999
:
Choice of south Asian summer monsoon indices.
Bull. Amer. Meteor. Soc.
,
80
,
629
638
.
Wang
,
B.
, and
LinHo
,
2002
:
Rainy season of the Asian–Pacific summer monsoon.
J. Climate
,
15
,
386
398
.
Wang
,
B.
,
R.
Wu
, and
X.
Fu
,
2000
:
Pacific–East Asian teleconnection: How does ENSO affect East Asian climate?
J. Climate
,
13
,
1517
1536
.
Wang
,
B.
,
R.
Wu
, and
K-M.
Lau
,
2001
:
Interannual variability of the Asian summer monsoon: Contrasts between the Indian and the western North Pacific–East Asian monsoons.
J. Climate
,
14
,
4073
4090
.
Webster
,
P. J.
,
1981
:
Mechanisms determining the atmospheric response to sea surface temperature anomalies.
J. Atmos. Sci.
,
38
,
554
571
.
Weng
,
H.
,
K-M.
Lau
, and
K-K.
Xue
,
1999
:
Multi-scale summer rainfall variability over China and its long-term link to global sea surface temperature variability.
J. Meteor. Soc. Japan
,
77
,
845
857
.
Wu
,
R.
, and
B.
Wang
,
2000
:
Interannual variability of summer monsoon onset over the western North Pacific and the underlying processes.
J. Climate
,
13
,
2483
2501
.
Xie
,
P.
, and
P. A.
Arkin
,
1997
:
Global precipitation: A 17-year monthly analysis based on gauge observations, satellite estimates, and numerical outputs.
Bull. Amer. Meteor. Soc.
,
78
,
2539
2558
.
Xu
,
J.
, and
J. C. L.
Chan
,
2001
:
The role of the Asian–Australian monsoon system in the onset time of El Niño events.
J. Climate
,
14
,
418
433
.
Yasunari
,
T.
,
1990
:
Impact of Indian monsoon on the coupled atmosphere/ocean system in the tropical Pacific.
Meteor. Atmos. Phys.
,
44
,
29
41
.
Zhang
,
R.
,
A.
Sumi
, and
M.
Kimoto
,
1996
:
Impacts of El Niño on the East Asian monsoon: A diagnostic study of the '86/87 and '91/92 events.
J. Meteor. Soc. Japan
,
74
,
49
62
.

Footnotes

Corresponding author address: Dr. Chia Chou, Environmental Change Research Center, Academia Sinica, Taipei, 115 Taiwan. Email: chia@earth.sinica.edu.tw