Symmetry and Asymmetry of the Asian and Australian Summer Monsoons

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1 15 JUNE 2004 HUNG ET AL Symmetry and Asymmetry of the Asian and Australian Summer Monsoons CHIH-WEN HUNG * Department of Atmospheric Sciences, University of California, Los Angeles, Los Angeles, California XIAODONG LIU Institute of Earth Environment, Chinese Academy of Sciences, Xi an, China MICHIO YANAI Department of Atmospheric Sciences, University of California, Los Angeles, Los Angeles, California (Manuscript received 21 January 2003, in final form 5 December 2003) ABSTRACT The rainfalls associated with the Asian summer monsoon have significant correlation with succeeding Australian summer monsoon rainfalls. This is partly due to the typical life cycle of the El Niño Southern Oscillation (ENSO) phase locked with the annual cycle within an Asian monsoon year (May April). This feature is referred to as the symmetric behavior of the Asian Australian monsoon system. However, preceding Australian summer monsoon rainfalls have no significant impact on succeeding Asian summer monsoon rainfalls. Thus, there is a one-way (asymmetric) relationship between the two monsoons and a communication gap exists in boreal spring prior to the onset of the Asian summer monsoon. The annual cycles of precipitation associated with the intertropical convergence zone (ITCZ) and the sea surface temperature (SST) anomaly in the Indonesia sector ( E) that links the two monsoon regions are compared. The continuous propagation of the ITCZ from the Asian to the Australian summer monsoon regions in Northern Hemisphere (NH) fall provides a one-way link between them. However, the SST anomaly shows a sudden change of sign in November, suggesting that the anomaly is mainly a response to the ITCZ movement rather than its cause. From January to May, the ITCZ is confined at the latitudes south of 5 N due to the subsidence in oceanic regions adjacent to the southern coast of Asia. This subsidence prevents the ITCZ from propagating northward until the abrupt onset of the Asian summer monsoon with the reversal of meridional temperature gradient. The discontinuity of the ITCZ movement and the seasonal heating over the Asian continent result in the communication gap in NH spring. 1. Introduction Among the monsoons of the world, the Asian and Australian summer monsoons are the most conspicuous (e.g., Webster et al. 1998). The Asian summer monsoon usually begins in May and ends in September. There are regions of large precipitation associated with seasonal reversals of wind directions in the Asian summer monsoon, such as Southeast Asia, the Bay of Bengal, and the west coast of India. The Australian summer monsoon, on the other hand, usually begins in December and ends in March. Its mature stage is characterized by * Current affiliation: Department of Atmospheric Sciences, National Taiwan University, Taipei, Taiwan. Corresponding author address: Dr. Michio Yanai, Department of Atmospheric Sciences, University of California, Los Angeles, 405 Hilgard Avenue, Los Angeles, CA heavy precipitation and low-level westerly wind in northern Australia and the Arafura Sea (e.g., McBride 1987; Manton and McBride 1992). Although these two monsoons occur in adjacent longitudinal sectors on opposite sides of the equator in the respective summer months, there are similarities as well as notable differences between them. The year-to-year variations of the rainfalls in southern Asia and northern Australia have generally negative relationships with the sea surface temperature (SST) anomaly in the equatorial eastern Pacific Ocean (e.g., Rasmusson and Carpenter 1983; Mooley and Shukla 1987; Nicholls 1992). While the effect of the El Niño Southern Oscillation (ENSO) on the Australian summer monsoon is pronounced (e.g., Holland 1986; Drosdowsky 1996), interactions of the ENSO and the Asian summer monsoon are more complicated (e.g., Yasunari 1990, 1991; Webster and Yang 1992; Torrence and Webster 1999; Slingo and Annamalai 2000). It has been reported that the correlation between the ENSO and the 2004 American Meteorological Society

2 2414 JOURNAL OF CLIMATE VOLUME 17 summer monsoon rainfall over the Indian subcontinent became weak by the late 1980s (e.g., Krishna Kumar et al. 1999). Several studies, on the other hand, related the climate regime shift over the entire North Pacific Ocean around 1976 to the change in the relationship between the Asian monsoon and ENSO (e.g., Kinter et al. 2002). The relationship between the Asian and Australian monsoons has been studied by several authors. Meehl (1987) noted that a region of active convection moves from northwest to southeast (from the Asian summer monsoon to the Australian summer monsoon regions) as the annual cycle proceeds from Northern Hemisphere (NH) summer to Southern Hemisphere (SH) summer. Joseph et al. (1991) found that the date of Australian summer monsoon onset is correlated with the monsoon rainfall in India during the preceding NH summer. Analyzing the seasonal variations of the wind field and outgoing longwave radiation (OLR), Matsumoto (1992) showed that the advances and retreats of the Asian and Australian monsoons are closely related. Gregory (1991), on the other hand, studied the station precipitation data in the two monsoon regions during and showed that the rainfall amount in the Australian summer monsoon is correlated with that in the preceding Asian summer monsoon. He found, however, that Australian summer monsoon rainfalls have no significant effects on succeeding Asian summer monsoon rainfalls. This result suggests that the Asian and Australian summer monsoons are not simple mirror images of each other and there is a one-way or asymmetric relationship between them. The ENSO creates a nearly symmetric impact on both monsoons through the east west Walker circulation. However, because of the differences in land sea distributions (Asia/Indian Ocean versus Australia/western Pacific) and continental size and topography, there is a great spatial asymmetry between the heat source distribution and circulation patterns associated with the Asian and Australian monsoons. The spatial asymmetry of the two monsoons reflects in the differences between spring and fall circulation patterns, thus creating a temporal asymmetry of the seasonal cycles in the two monsoon regions. This paper examines further how the spatial and temporal asymmetries in the Asian and Australian monsoons translate into the one-way relationship between the associated rainfall amounts. In section 2, the datasets used in this study including the estimated heat sources and moisture sinks are described. In section 3, the relationship between the Asian and Australian summer monsoon rainfalls is reexamined. Nearly symmetric impact of the ENSO on the Asian and Australian summer monsoons as well as asymmetric behavior of the two monsoon systems are discussed. In section 4, the spatial and temporal asymmetries in heat source distribution and circulation patterns associated with the two monsoons are presented. In section 5, the annual cycles of the intertropical convergence zone (ITCZ) and SST anomaly in a sector connecting the two monsoon regions are compared. In section 6, the movement of the ITCZ is examined in relation to the time evolutions of the heat source and induced subsidence in this sector. Finally, section 7 presents a summary. 2. Data The primary data source used for this study is the 15- yr ( ) European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA; Gibson et al. 1997). The ERA dataset is chosen, despite its relative short record length, because the moisture fields of the National Centers for Environmental Prediction National Center for Atmospheric Research (NCEP NCAR) reanalysis (Kalney et al. 1996) contain large and significant biases in the Tropics (Trenberth and Guillemot 1998; Yanai and Tomita 1998; Annamalai et al. 1999). The monthly Climate Prediction Center (CPC) Merged Analysis of Precipitation (CMAP; Xie and Arkin 1997), the Comprehensive Ocean Atmosphere Data Set (COADS) provided by the National Oceanic and Atmospheric Administration (NOAA) Climate Diagnostics Center (CDC), and the Global Sea Ice and Sea Surface Temperature dataset (GISST 2.3b) created by the Hadley Centre (Rayner et al. 1996) for the same 15 years or a period extending to recent years are also used. In addition, the apparent heat source Q 1 and the apparent moisture sink Q 2 (Yanai et al. 1973) are computed from large-scale heat and moisture budget: p p t p o Q c v and (1) 1 p q q Q2 L v q, (2) t p where is the potential temperature, q the mixing ratio, v the horizontal velocity, the vertical p velocity; L is the specific latent heat of vaporization, R/c p with R the gas constant and c p the specific heat capacity at constant pressure of dry air, and p hpa. The overbar denotes the running horizontal average over a large-scale area. The horizontal wind components, potential temperature, and mixing ratio are directly obtained from ERA products; however, for the accuracy of Q 1 and Q 2, the vertical p velocity is recomputed following the method described in Tung et al. (1999) and Tung and Yanai (2002). In order to express Q 1 and Q 2 in units of equivalent warming/cooling rate (K day 1 ), a constant number, 1 c p, is multiplied for both of them. Integrating Eqs. (3) and (4) from tropopause pressure p t to the surface pressure p s, we can obtain

3 15 JUNE 2004 HUNG ET AL in a region is mainly contributed by the condensation process, the values of Q 1 and Q 2 should be similar. On the other hand, if there is strong sensible heat flux or evaporation from the surface, the horizontal distributions of Q 1 and Q 2 are significantly different (e.g., Luo and Yanai 1984; Yanai and Tomita 1998). FIG averaged 850-hPa streamlines and precipitation rate (mm day 1, shaded) for (a) JJA and (b) DJF. The northern and southern rectangular boxes show the regions of southern Asia (SAS) and northern Australia (NAU) respectively (see text). The thick solid line is the smoothed topographic contour of 4000 m showing the outline of the Tibetan Plateau. where Q Q LP S and (3) 1 R Q L(P E), (4) 2 p 1 s ()dp; (5) g P, S, and E are the precipitation rate, the sensible heat flux, and the evaporation rate, respectively, per unit area at the surface (Yanai et al. 1973). The left-hand terms Q 1 and Q 2 are useful for physical interpretation of the budget results. If the heating pt 3. Symmetry and asymmetry in Asian and Australian monsoon rainfalls Precipitation is the most important meteorological element to measure monsoon activities. Therefore, we first inspect the climatological distributions of precipitation over the Asian Australian monsoon area. Figures 1a and 1b show the distributions of mean precipitation rates and 850-hPa streamlines in NH summer (June August, hereafter JJA) and SH summer (December February, hereafter DJF), respectively. An area in southern Asia (8 23 N, E, hereafter SAS) and an area including northern Australia (2 15 S, E, hereafter NAU) of a similar size are chosen to represent the major rainy regions of the Asian and Australian monsoons. The monthly precipitation rates averaged over SAS and NAU in are shown in Fig. 2. The yearto-year variation of rainfalls in NAU is larger than that in SAS, although the mean values in both regions are similar. The monsoon seasons in SAS and NAU are well separated in time, and the precipitation in both regions is concentrated in different half years (May October for SAS; November April for NAU). From Fig. 2, we readily see the disadvantage of using the calendar year (January December) as a unit to describe the annual Asian Australian monsoon activities because an Australian summer monsoon season is always spread over two consecutive calendar years. If a new unit of year started from May or November is used, there is no such a disadvantage. The concept monsoon year was first proposed by Yasunari (1991) as a unit year that starts just before the northern summer monsoon season. Following Yasunari, we first define the Asian monsoon year beginning with May and ending with April. Figure 3a shows the standardized annual rainfall amounts over SAS and NAU and the Niño-3.4 SST index (SST anomalies averaged over 5 S 5 N, W), measured in the Asian monsoon year. For later discussion, similar figures, but for the calendar year and the Australian monsoon year (November October), are shown in Figs. 3b and 3c, respectively. The interannual variability of the rainfalls in the SAS and NAU regions have negative correlations with the SST anomalies in the equatorial eastern Pacific Ocean [e.g., Mooley and Shukla (1987) for the Asian monsoon; Nicholls (1992) for the Australian monsoon]. The correlation coefficients between the Niño-3.4 (5 S 5 N, W) SST index and the annual SAS as well as NAU rainfalls, measured in the Asian monsoon year,

4 2416 JOURNAL OF CLIMATE VOLUME 17 FIG. 2. Monthly precipitation rate (mm day 1 ) averaged over SAS (solid lines) and NAU (dashed lines) in The time axis starts from May and ends in Apr. the calendar year, and the Australian monsoon year in are shown in Table 1. For all three different units of a year, the Niño-3.4 SST index has larger negative correlation coefficients with the NAU rainfalls than those with the SAS rainfalls. The magnitudes of correlation coefficients between the yearly rainfalls (in SAS or NAU) and the Niño-3.4 SST index measured in the Asian monsoon year are only slightly larger than those measured in the Australian monsoon year. This indicates that the impacts of the ENSO on the two monsoons are qualitatively similar regardless of the choice of monsoon years. We refer to this feature as the symmetric behavior of the Asian Australian monsoon system. The standardized annual rainfall amounts measured in the Asian monsoon year over SAS and NAU are similar (Fig. 3a), but the yearly SAS rainfalls exhibit clearer quasi-biennial tendency, also known as the tropospheric biennial oscillation (e.g., Trenberth 1975; Meehl 1987, 1993, 1997; Yasunari 1990, 1991) than the yearly NAU rainfalls. We average the yearly standardized rainfalls over SAS and NAU to define an Asian Australian monsoon index to determine strong and weak Asian Australian monsoon years in The years with the Asian Australian monsoon index larger (smaller) than zero are called strong (weak) monsoon years. The monsoon year of 2000 (from May 2000 to April 2001) is not used, because the CMAP data ends in December With this index, 1980, 1981, 1983, 1985, 1988, 1990, 1995, 1996, 1998, and 1999 are strong monsoon years and 1979, 1982, 1984, 1986, 1987, 1989, 1991, 1992, 1993, 1994, and 1997 are weak monsoon years. The averages of the monthly Niño-3.4 SST index (the mean annual cycle is removed) are then calculated for the strong and weak Asian Australian monsoon years and shown in Fig. 4. The time series in this figure indicates that the strong/weak Asian Australian monsoon is phase locked with the seasonal variation of the SST anomaly over the eastern Pacific Ocean. For the strong (weak) Asian Australian monsoon cases, the Niño-3.4 SST is undergoing a cold (warm) excursion within an Asian monsoon year. Because a typical warm or cold event in the ENSO cycle lasts for about one year starting from NH spring and reaches the maximum magnitude in December, it provides a similar impact on the Asian and Australian monsoon regions within an Asian monsoon year. This shows the symmetric impact of the ENSO on the Asian Australian monsoon system. The correlation coefficients between the SAS and NAU rainfalls measured in the Asian monsoon year, the calendar year, and the Australian monsoon year are calculated and shown in Table 2. The correlation coefficient between the annual rainfalls in SAS and NAU measured in the Asian monsoon year (Fig. 3a) for the period is 0.61 (Table 2). On the other hand, the correlation coefficient is only 0.36 when the calendar years are used (Fig. 3b). When we use the Australian monsoon year (Fig. 3c), the correlation coefficient between the annual rainfall amounts in SAS and NAU becomes even smaller (0.32) than that using the calendar year. These confirm that the preceding Asian summer monsoon rainfall has

5 15 JUNE 2004 HUNG ET AL FIG. 4. The composites of monthly Niño-3.4 SST index for the strong (black line) and weak (gray line) Asian Australian monsoon (see text). The big rectangular box indicates the Asian monsoon year (May Apr). FIG. 3. Yearly variations of standardized annual mean precipitation rates over SAS (solid line) and NAU (dashed line), and the standardized Niño-3.4 SST index (gray solid line) measured by (a) the Asian monsoon year, (b) the calendar year, and (c) Australian monsoon years (see text). a significant correlation with the succeeding Australian rainfall, but the preceding Australian monsoon rainfall has no significant correlation with the succeeding Asian summer monsoon rainfall (Gregory 1991). We further calculate correlation coefficents of the annual (May April) precipitation amount in SAS with those in all grid boxes ( ) for the years in (Fig. 5a). The regions with high positive correlation coefficients are connected all the way from the India Bay of Bengal region, to Indonesia and to TABLE 1. Correlation coefficients of the yearly SAS and NAU rainfalls with the Niño-3.4 SST index in when the Asian monsoon year (May Apr), calendar year, and Australian monsoon year (Nov Oct) are used. The 95% confidence level is Asian monsoon year Calendar year Australian monsoon year SAS vs Niño-3.4 SST NAU vs Niño-3.4 SST northern Australia, forming a corridor, although the rain seasons are concentrated in different months in an Asian monsoon year for different locations. When the correlation coefficients are calculated between the annual precipitation amounts at NAU and those in all other grid boxes using the Australian monsoon year (November October), we find no significant positive correlation coefficients in SAS (Fig. 5b). The correlation corridor in Fig. 5b is discontinued in the oceanic area near southern Asia. In other words, the corridor shown in Fig. 5a is a one-way street. Meehl (1987) found similar facts using OLR, precipitation, and sea level pressure. It is worth noting that the correlation coefficients between the Niño-3.4 SST index and rainfalls in SAS or NAU do not have large differences when the yearly values are measured in either the Asian or Australian monsoon years (Table 1). However, the remarkable difference between the SAS versus NAU rainfall correlation measured in the Asian monsoon years and that in the Australian monsoon years (Table 2) suggests that a temporal asymmetry exists in the annual cycles of the Asian and Australian monsoons. The poor correlation between the Australian summer monsoon rainfalls and the succeeding Asian summer monsoon rainfalls is partly due to the phase-locked life cycle of the ENSO, which usually terminates its warm or cold excursion in NH spring (Fig. 4). This typical life cycle of the ENSO results in the predictability barrier of the SST over the eastern Pacific Ocean in the boreal spring and creates a similar barrier in the Asian Australian monsoon system (e.g., Yasunari 1990, 1991; Yasunari and Seki 1992; Webster and Yang 1992; Lau and Yang 1996). However, the year-to-year variation of the Asian Australian monsoon is not always related to the ENSO. The observed TABLE 2. Correlation coefficients between the yearly SAS and NAU rainfalls in when the Asian monsoon year (May Apr), calendar year, and Australian monsoon year (Nov Oct) are used. The 95% confidence level is Asian monsoon year Calendar year Australian monsoon year SAS vs NAU

6 2418 JOURNAL OF CLIMATE VOLUME 17 rainfall in the Asian Australian monsoon system is correlated with the ENSO in a symmetric way, this fact alone does not explain the one-way (northwest to southeast) corridor of high correlations from the Asian to the Australian summer monsoon regions. This motivates us to examine whether asymmetry of the mean annual cycle in the Asian and Australian monsoons contributes to this gap. FIG. 5. Correlation coefficients (only positive values larger than 0.43, the 95% confidence level, are shown) between (a) the annual mean precipitation (measured in the Asian monsoon year) over the SAS region (shown in a rectangular box) and all grid boxes, and (b) the annual mean precipitation (measured in the Australian monsoon year) over the NAU region (shown in another rectangular box) and all grid boxes in biennial tendency (Fig. 3a) in the two monsoons is not related to the ENSO in some years (e.g., Lau and Wu 2001). The results presented in this section demonstrate that the Asian and Australian summer monsoons can be considered as one system that varies together when the Asian monsoon year (May April) is used, but there is a communication gap between the two monsoon regions in NH spring prior to the onset of the Asian summer monsoon. Although the interannual variability of the 4. Asymmetries of heat source distribution and circulation patterns The differences in land sea distribution, topography, and continental size in the Asian and Australian regions introduce spatial asymmetries of the heating distributions and the circulation patterns related to the two monsoon systems. Land has two major effects on the largescale atmospheric motion through (i) mechanical forcing due to topography and (ii) thermal forcing due to land sea contrast. The Asian continent contains the Tibetan Plateau (the average height is greater than 4000 m), causing airflow going around it (e.g., Yin 1949; Trenberth and Chen 1988) and acting as an elevated heat source (e.g., Flohn 1957; Yanai et al. 1992; Chou 2003) for the Asian summer monsoon. However, the Australian continent has no high land features to provide the elevated heat source and produce large impacts on its surrounding regions (e.g., Webster et al. 1998). Figure 6 shows the vertical distributions of Q 1 and Q 2 averaged between 70 and 100 E (the same longitudes as the SAS region defined previously) in NH spring (March May, hereafter MAM) and summer (JJA). The large values of Q 1 and Q 2 seen near the equator are associated with the migrating ITCZ, which contains deep convection throughout the year. In addition, during MAM, large values of Q 1 are seen from the surface to 250 hpa above the Tibetan Plateau (Fig. 6a). In contrast, the values of Q 2 there are small (Fig. 6c). This shows that the heating above the Tibetan Plateau in NH spring is mainly contributed by sensible heat flux from the surface (e.g., Yanai et al. 1992; Yanai and Li 1994). After the onset of the Asian summer monsoon, heating from condensation in the ITCZ and the heating over the land surface merge together. In JJA, large heating between the surface and 200 hpa is centered over the southern edge of the Tibetan Plateau, although weak peaks of Q 1 and Q 2 are still observed between 10 S and the equator (Figs. 6b and 6d). Figure 7 displays the vertical distributions of Q 1 and Q 2 averaged between 115 and 150 E (the same longitudes as the NAU region). The ITCZ represented by large Q 1 and Q 2 migrates between the two hemispheres. Shallow sensible heating over the Australian continent is observed in every season, but it is the most intense in DJF (Fig. 7b). When the ITCZ is still located in the NH in September November (SON), the sensible heating from the continental surface has already increased (Fig. 7a). Although the land sea thermal contrast in this

7 15 JUNE 2004 HUNG ET AL FIG. 6. Latitude height sections of the mean (a) MAM Q 1, (b) JJA Q 1, (c) MAM Q 2, and (d) JJA Q 2 averaged over E. Contour interval is 0.5 K day 1. region provides an environment for the monsoon to occur (Hung and Yanai 2004), the sensible heating from land surface and the heating resulted from deep convection of the ITCZ are well separated. In Figs. 8 and 9, the upper-tropospheric ( hpa) mean temperature and 200-hPa mean streamlines for JJA and DJF are shown, respectively. The horizontal distribution of hPa mean temperature in JJA (Fig. 8a) shows a huge area of warm air centered over the southern edge of the Tibetan Plateau. This warm region corresponds to the large Q 1 observed in Fig. 6b and results in a strong temperature gradient between the Asian continent and the Indian Ocean, which maintains the monsoon system (e.g., Flohn 1957; Yanai et al. 1992; Li and Yanai 1996). On the other hand, the flat Australia continent does not elevate the sensible heat to higher levels, so the upper-level mean temperature does not show a warm core (Fig. 8b). The streamlines at 200 hpa also display this difference between the two hemispheres. In Fig. 9a, a huge anticyclone (the so-called South Asian or Tibetan high) is centered at the southern edge of the Tibetan Plateau (Krishnamurti 1971a, b). The circulation associated with this anticyclone extends almost one-third of the globe in the subtropical region of the NH. In contrast, the circulation pattern in DJF (Fig. 9b) does not show an impact of the Australian

8 2420 JOURNAL OF CLIMATE VOLUME 17 FIG. 7. Latitude height sections of the mean (a) SON Q 1, (b) DJF Q 1, (c) SON Q 2, and (d) DJF Q 2 averaged over E. Contour interval is 0.5 K day 1. continent. Instead, the so-called Matsuno Gill pattern (Matsuno 1966; Gill 1980), with one pair of anticyclones and an axis of diffluence along the equator, is observed in the 200-hPa streamlines. The impacts of the Asian and Australian summer monsoon systems on the surrounding regions are evidently different. In SON, which follows the Asian summer monsoon season (JJA), major features of the uppertropospheric temperature and the circulation patterns in JJA are still recognizable, although centers of the warm area and the South Asian anticyclone moves southeastward to near the Philippines (not shown). In contrast, during MAM, which is after the Australian summer monsoon season (DJF), the patterns remain symmetric about the equator. Thus, there is a notable asymmetry in the time evolution of the Asian and Australian summer monsoons. 5. Annual cycles of precipitation and SST anomalies In this section, we first examine the annual cycles of the ITCZ and the SST in the meridional strip E (named the Indonesia sector) that links the Asian and Australian summer monsoon regions. Figures 10a d show the mean monthly anomalies of precip-

9 15 JUNE 2004 HUNG ET AL FIG. 8. The mean upper-tropospheric ( hpa) temperature (K) for (a) JJA and (b) DJF. The thick solid line is the smoothed topographic contour of 4000 m showing the outline of the Tibetan Plateau. itation rate (from CMAP), SST (from GISST), wind speed (from COADS), and latent heat parameter (from COADS), respectively. The annual means are removed for each month, and the same latitude time section is plotted twice to show the Australian summer monsoon season continuously. The thick curve is the solar declination angle to indicate seasons. Dots are plotted in May and November to indicate major periods for discussion. The ITCZ, by definition, is located in a region with heavy rainfall and low-level convergence (thus small wind speed). In the latitude time section of the precipitation anomaly (Fig. 10a), the positive values between 20 S and 30 N show the rain belt of the ITCZ moving between the two hemispheres. The COADS surface wind speed (Fig. 10c) shows that its negative anomalies from September to February coincide with the positive anomalies of rainfall (Fig. 10a). During NH fall and winter, the rain belt of the ITCZ propagates smoothly from north to south with some lag from the seasonal solar forcing, consistent with the corridor of positive correlation of annual rainfalls seen in Fig. 5a. However, there is no bridge extending from the Australian summer monsoon to the following Asian summer monsoon, and FIG. 9. Similar to Fig. 8 except for the streamlines at 200 hpa. a discontinuity in the ITCZ movement is seen in NH spring and early summer (Fig. 10a). The SST anomaly distribution (Fig. 10b) shows a propagation characteristic quite different from that of precipitation. The warm SST anomaly continuously propagates from the SH to the NH from March to June. The change of sign of SST anomalies from negative north (positive south) to positive north (negative south) (marked by a dot in May) is rather gradual. Because the region has clear sky (from COADS cloudiness data; not shown), the SST anomalies closely follow the seasonal solar forcing during this period. In contrast to the smooth southward movement of the ITCZ in NH fall (Fig. 10a), a sudden change of the sign of SST anomaly is observed in November. The pattern of positive north (negative south) with respect to the equator quickly changes to that of negative north (positive south). The sudden change of the sign of SST anomalies in November (Fig. 10b) may be explained by the wind evaporation SST (WES) feedback mechanism (Xie and Philander 1994). In the WES theory, the bulk aerodynamic formula is commonly used to describe the evaporation rate at the ocean surface. The evaporation rate is calculated by a term multiplied from latent heat parameter, air density, and a dimensionless bulk transfer coefficient C e for water vapor (Liu et al. 1979). The latent heat parameter is a value calculated by multiplying the wind speed and the air sea humidity difference

10 2422 JOURNAL OF CLIMATE VOLUME 17 the wind speed (Fig. 10c) and the evaporation rate (Fig. 10d) are small between 25 S and 25 N. The sudden change of sign of the SST anomaly does not occur during this period and the warm SST continuously propagates from the SH to the NH (Fig. 10b). Although the SST anomaly continuously propagates from NAU to SAS in January to May (Fig. 10b), the ITCZ does not follow the warm SST anomaly or the gradient of the total SST (not shown). This fact, along with the sudden change of sign of the SST anomaly in November, suggests that the behavior of SST in the region is a result of the ITCZ movement. The SST anomaly plays a passive role that reflects the movement of the ITCZ in the annual cycle and it does not explain the observed movement of the ITCZ. FIG. 10. Time latitude sections of the mean monthly anomalies of (a) precipitation (contour interval is 1 mm day 1 ), (b) SST (contour interval is 1 K), (c) surface wind speed (contour interval is 0.5 m s 1 ), and (d) latent heat parameter (contour interval is 2gkg 1 ms 1 ) averaged over E. The annual means are removed, and the same latitude time section is plotted twice to show the Australian summer monsoon season continuously. The thick curve is the solar declination angle. Dots are plotted in May and Nov to indicate major periods for discussion in text. (the difference between the saturation specific humidity at the value of SST and the specific humidity at the air temperature). Because the variations of the air density and air specific humidity are very small, the evaporation rate can be considered as a nonlinear combination of SST and wind speed. The variation of the latent heat parameter approximately represents the variation of the evaporation rate. In November, the ITCZ is located at the equator when it is moving southward from the NH to the SH. The region north of the ITCZ has strong wind speeds as expected (Fig. 10c). The COADS latent heat parameter (Fig. 10d) has a pattern similar to that of the wind speed because the wind speed dominates the variation of the evaporation over the warm pool (Zhang and McPhaden 1995). The strongest evaporation occurs north of the equator in this sector in November and results in cooling of the SST (Fig. 10b). Therefore, the sudden north south exchange of SST signs in November is consistent with the ITCZ movement. In contrast, during NH spring 6. Movement of ITCZ and effect of continental heating The results presented in the previous section show that the SST in the Indonesia sector ( E) is not responsible for the movement of the ITCZ and it does not explain the existence of a communication gap in NH spring prior to the onset of the Asian summer monsoon. In order to identify the cause of the communication gap, we closely examine the movement of the ITCZ with the time evolution of heat sources in this sector. Figure 11a d show the latitude time sections of Q 1, Q 2, precipitation rate, and the mean upper-tropospheric ( hpa) temperature averaged in the same sector from 1979 to The daily Q 1 and Q 2 are averaged from 6-hourly data, the precipitation rate is from the pentad CMAP, and the upper-tropospheric mean temperature is from the ERA. The thick curves in the figures indicate the solar declination angle. We note a large region of positive values of Q 1 between 40 and 70 N from April to September in Fig. 11a without corresponding features in Q 2 (Fig. 11b) and rainfall (Fig. 11c). This region of positive Q 1 at high latitudes shows the sensible heating over the Asian continent. Its intensity and the influenced area closely follow the seasonal solar forcing without significant time lags. On the other hand, large positive values in Q 1 between 20 S and 30 N are with large Q 2 and rainfall. Weak rainfall (less than 4 mm day 1 ) observed between 20 and 30 N from late February to April are the spring rains over China (e.g., Tian and Yasunari 1998). A more prominent rain belt is associated with the ITCZ moving between the two hemispheres. During SON, the ITCZ propagates smoothly from north to south with some delay from the seasonal forcing. However, the ITCZ stays between 15 S and 5 N from January to May without any north south movement. Then, it shows a sudden northward jump in mid May at the onset of the Asian summer monsoon. The onset is concurrent with the reversal of the meridional temperature gradient in the upper troposphere over the Bay of Bengal region (e.g., Flohn 1957; Li and Yanai 1996). Although the

11 15 JUNE 2004 HUNG ET AL FIG. 11. Time latitude sections of the mean (a) daily Q 1, (b) daily Q 2, (c) pentad precipitation rate, and (d) upper-tropospheric ( hpa) temperature averaged over E. The contour interval for (a) and (b) is 50 W m 2, and for (c) is 2 mm day 1. For (d), the contour interval is 0.5 K (2 K) when the values are larger (smaller) than 248 K. The thick curve is the solar declination angle. upper-tropospheric temperature shown in Fig. 11d is not over the longitudes of the SAS region, it exhibits a similar reversal near the southern coast of the Asian continent between 90 and 120 E. Figures 12a and 12b show the mean (vertical p velocity) at 850 hpa in March and April, respectively. During these months, strong subsidence appears along the peripheries of continental heat sources (Tibetan Plateau, India, and the Indochina peninsula), including the oceanic regions adjacent to the southern coast of Asia. Although subsidence over the Arabian Sea and the Bay of Bengal is partly due to the ITCZ, the upward motions over land (India and Indochina) and the strong descending motions over the oceanic regions at the same latitudes suggest the effect of land heating. The subsidence occurred before the monsoon season and acted against the northward movement of the ITCZ. In SON and MAM, the ITCZ migrates between the Asian and Australian continents over the ocean (Figs. FIG. 12. Horizontal distributions of mean (vertical p velocity; contour interval is 0.01 Pa s 1 ) at 850 hpa in (a) Mar and (b) Apr. The shading corresponds to regions of ascent. 11a c). During SON, convection associated with the ITCZ continuously moves from the Asian summer monsoon region to the Australian summer monsoon region. This connected path of deep convection links the previous Asian summer monsoon to the following Australian summer monsoon. Therefore, a corridor of high correlation is generated when the annual rainfalls in SAS and NAU are measured in the Asian monsoon year (Fig. 5a). However, the ITCZ is confined to the latitudes south of 5 N from January to May because the subsidence over the Arabian Sea and the Bay of Bengal acts against its northward movement. At the same time, the seasonal heating over the Asian continent begins (Figs. 11a and 11b). When the ITCZ abruptly shifts northward and merges with the monsoon heat low induced by the continental heating, a new Asian summer monsoon season starts without much influence of the previous Australian summer monsoon. Thus, no significant relationship between the preceding Australian summer monsoon rainfall and succeeding Asian summer monsoon rainfall is found. Similar subsidence is found around the coast of Australia in September and October (Fig. 13). The subsi-

12 2424 JOURNAL OF CLIMATE VOLUME 17 FIG. 13. Similar to Fig. 12, except for (a) Sep and (b) Oct. dence along the northern coast of Australia in SH spring is weaker and shallower than that formed over the oceanic region adjacent to the southern coast of Asia in NH spring (Fig. 12). Therefore, the belt of convection moving down from SAS smoothly reaches NAU. 7. Summary and conclusions Using the unit of the Asian monsoon year (May April), the correlation between the preceding Asian summer monsoon and the succeeding Australian summer monsoon in rainfall is clearly recognized, and the two monsoons are considered as one system that varies together. This feature is referred to as the symmetric behavior of the Asian and Australian monsoons. It is partly due to (i) the generally negative correlation between rainfalls associated with both monsoons and the SST of the eastern Pacific Ocean and (ii) the typical life cycle of the ENSO, which is phase locked to the annual cycle within an Asian monsoon year. In contrast, there is no significant correlation between the preceding Australian summer monsoon rainfall and succeeding Asian summer monsoon rainfall. Correlation analysis of the annual rainfall of SAS and other regions shows that a corridor with high positive correlation coefficients extends from SAS to NAU when the Asian monsoon year is used. However, the correlation corridor is discontinued in the oceanic area near southern Asia when the yearly values are measured in the Australian monsoon year (November October). This one-way or asymmetric relationship between the Asian and Australian monsoon rainfalls suggests the existence of a communication gap in NH spring. The differences between the two monsoon regions in land sea distribution, continental size, and topography introduce spatial asymmetries in the heat source distributions and the circulation patterns associated with the two monsoons. A notable feature is the sensible heating over the Tibetan Plateau in NH spring summer. The Australian continent, on the other hand, produces only a shallow heating layer near the surface. These differences reflect in the upper-tropospheric temperature distributions and flow patterns during the respective monsoon and following seasons. The propagation characteristics of the precipitation and SST anomalies in the Indonesia sector ( E) connecting the two monsoon regions are quite different. In NH spring and early summer, the SST anomaly in this sector continuously moves from NAU to SAS following the seasonal solar forcing, but the ITCZ does not propagate with the warm SST anomaly and stays at the latitudes south of 5 N. In NH fall and winter, the ITCZ moves smoothly from the NH to the SH, but a sudden change of sign of SST anomalies occurs in November. This sudden change is consistent with the wind evaporation SST relationship, suggesting that the SST anomaly is a response to the movement of the ITCZ in the region. The continuously propagating ITCZ in the Indonesia sector during NH fall provides a one-way link between the Asian and Australian summer monsoons. However, the ITCZ is confined south of 5 N from January to May due to the subsidence located along the southern periphery of the continental heat sources (e.g., India and Indochina) until the abrupt onset of the Asian summer monsoon. This onset is concurrent with the reversal of the meridional temperature gradient in the upper troposphere resulting from the seasonal heating over the Asian continent. Because the northward movement of the ITCZ is prevented in NH spring, a new Asian summer monsoon season starts without significant influence of the previous Australian summer monsoon. This asymmetric feature between the Asian and Australian monsoons, seen in the mean annual cycle, provides the basic relationship between the two monsoons. When the interactions with other factors such as the ENSO (e.g., Yasunari 1991; Webster and Yang 1992) and the Indian Ocean dipole (e.g., Saji et al. 1999; Ashok et al. 2001) are considered, the mean annual cycles of the two monsoons are modified. Thus, interannual variability such as the tropospheric biennial oscillation in the Asian Australian monsoon system may appear (e.g., Meehl and Arblaster 2002).

13 15 JUNE 2004 HUNG ET AL Acknowledgments. The authors thank two anonymous reviewers and the editor M. P. Hoerling for their useful comments on the manuscript. The suggestions from W. C. Chao, A. Hall, J. D. Neelin, M. Raphael, H. Su, S.- P. Weng, and J.-Y. Yu are appreciated. Special thanks are due to W.-w. Tung for her help with the calculations of Q 1 and Q 2. The computations were performed at the NCAR Scientific Computing Division and UCLA Department of Atmospheric Sciences. This work was supported by the National Oceanic and Atmospheric Administration under Grant NA96GP0331 and by the National Science Foundation under Grant ATM REFERENCES Annamalai, H., J. M. Slingo, K. R. Sperber, and K. Hodges, 1999: The mean evolution and variability of the Asian summer monsoon: Comparison of ECMWF and NCEP NCAR reanalyses. Mon. Wea. Rev., 127, Ashok, K., Z. Y. Guan, and T. Yamagata, 2001: Impact of the Indian Ocean Dipole on the relationship between the Indian monsoon rainfall and ENSO. 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