A possible mechanism effecting the earlier onset of southwesterly monsoon in the South China Sea compared to the Indian monsoon

Meteorol. Atmos. Phys. 76, 237±249 (2001) Department of Atmospheric Sciences, Nanjing University, Nanjing China, 210093 A possible mechanism effecting the earlier onset of southwesterly monsoon in the South China Sea compared to the Indian monsoon Y.-F. Qian, S.-Y. Wang, and H. Shao With 10 Figures Received June 13, 2000 Revised October 18, 2000 Summary A possible mechanism effecting the earlier onset of the South China Sea (SCS) southwesterly (summer) monsoon compared to the Indian monsoon is revealed by use of the NCAR/NCEP reanalysis data during the 1998 South China Sea Monsoon Experiment (SCSMEX). It is found that the fundamental cause might be the different spatial distributions and time variations of the surface sensible heat (SH) uxes in the two areas. In the SCS area, positive surface sensible heating exists all the time in both the tropics and the subtropics with the latter being dominant after March. While in the India Peninsula (IP) area, the tropical heating is the dominant one at the early period before May and the subtropical heating is smaller and even negative. It is further found that the west-east contrasts of the SH uxes at middle latitudes are also favorable to the earlier onset of the SCS monsoon. The time evolutions of circulation expose the dynamic cause. A strong cyclone existing between the equator and 10 N at the south tip of the IP on May 13 may be taken as an indicator of the monsoon developing process. The circulation of the cyclone and the blocking effect of the Tibetan Plateau (TP) combine together to make the southwesterly wind easily reach the SCS area. When and only when the cyclone moves towards the TP, becomes weaker and weaker and forms a shallow monsoon trough at last, the Indian monsoon bursts out. Therefore, it takes almost a month to occupy the entire IP area and is later than the SCS monsoon onset by about one month. Jointly sponsored by the National Natural Science Foundation of China under the grant ``Analyses and Mechanism Studies of Regional Climate Changes in China'' (No. 49735170) and the SCSMEX. 1. Introduction The Indian summer monsoon has been considered as the Asian monsoon for a long time. However, the Chinese scientists pointed out that the Asian summer monsoon consists of the East Asian summer monsoon (EASM) and the South Asian summer monsoon (SASM) which is commonly called the Indian summer monsoon (see Chen et al., 1991). The former consists again of the tropical and the subtropical Asian summer monsoons. The Indian and the South China Sea (SCS) summer monsoons belong to the tropical monsoons due to the southwesterly winds coming from the tropics over the two areas in the monsoon period. There are quite a lot of studies and reviews on monsoons made by Ramage (1971), Krishnamurti (1985), Tao and Chen (1987), Webster and Yang (1992), Ding and Murakami (1994), Lau and Yang (1996), Shen and Lau (1995) and others. Ma and Ding (1997) summarized the studies on the EASM in recent years in China. The tropical monsoon in the SCS and the surrounding areas is also the Southeast Asian monsoon de ned by Lau and Yang (1997). Wang and Wu (1997) de ned a western North Paci c monsoon and considered the SCS monsoon as the junction of the three major Northern Hemisphere summer monsoons, that is, the Indian monsoon,

238 Y.-F. Qian et al. the western North Paci c monsoon and the East Asia subtropical monsoon. Therefore, the summer monsoons in various areas of Asia are not uniquely de ned. In this paper we take the SCS monsoon as a component of the EASM as de ned by most Chinese meteorologists. It is well known that the EASM begins rst in the SCS, and then propagates northward to the mainland of China. The onset period of the SCS monsoon is usually in the middle of May, while the onset of the Indian monsoon over the entire India subcontinent is in the middle of June. The time lag of the latter is about one month. As it is commonly known that the fundamental mechanism of monsoon is the thermal contrast between land and sea. The large scale thermal contrast between the Eurasian and the African continent and the Western Paci c and the Indian Ocean results in the most obvious monsoon phenomena over the regions. This kind of thermal contrast induces the large-scale monsoon circulation. The thermal contrast between the plain land and the high land, such as the Tibetan Plateau (TP), should also have evident in uences on monsoon phenomena. The so-called ``Plateau monsoon'' is produced by this kind contrast (see Ye and Gao, 1979). The third thermal contrast comes from local geographical conditions. For example, in the tropics there are underlying water surfaces such as the SCS, the Bay of Bengal (BB) and the Arabian Sea (AS), as well as underlying land surfaces such as the India Peninsula (IP), the Indochina Peninsula (ICP) and some marine islands. As pointed out above, the different monsoon subsystems onset at different periods and should be resulted from the different local contrasts between land and sea. However, the local conditions of land-sea contrast in the SCS area and the IP area are somewhat similar to each other, why does the SCS summer monsoon onset almost one month earlier than that of the Indian monsoon? This problem is worth studying for deeper understanding of the monsoon mechanisms. In this paper, we are going to study the possible mechanism that causes the earlier onset of the SCS monsoon than that of the Indian monsoon from two aspects. One is the thermal contrast that may be different in the SCS and the IP areas. The other is the circulation pattern which can be affected by both the TP blocking and its elevated heating effects. Therefore, the NCAR/NCEP reanalysis daily and pentad data of surface heat uxes and circulation elds in the 1998 SCSMEX are used. The data are believed reliable because of the more accurate and intensive observations made in the SCSMEX and of carefully reanalyzing for convenient use. The heat uxes are computed in terms of the four dimensional assimilation by use of both observational and model results. 2. Onset dates of the SCS and the Indian summer monsoons The onset dates of the summer monsoon can be found from the time variations of some basic meteorological element elds, such as the wind components, the cloudiness, the precipitation, the surface air temperature at 2 m and the humidity (He et al., 1987). Figure 1 reveals the time evolutions of the above elements averaged in the SCS area (7.5 N 20 N, 110 E 120 E) in 1998 from January to August. The ordinates are the values of the elements and the abscissas the pentads. It is seen from Fig. 1 that there are abrupt Fig. 1. Time evolutions of the pentad mean wind components a the cloudiness, b the precipitation, c the surface air temperature at 2 m, d and the humidity e. The solid curve refers to the u-component and the dashed the v-component

A possible mechanism effecting the earlier onset of southwesterly monsoon 239 changes of wind components (especially the u-component), cloudiness and precipitation in the 28th pentad (May 16 20), the time changes of the other elements are not so abrupt though, they show large changes also around the 28th pentad. Therefore, the SCS summer monsoon onsets about in the 28th pentad in 1998 that is a little later than the normal. However, in Sect. 4 it will be found that the onset date of the SCS monsoon is May 23. The inconsistency between the two onset dates is due to the area difference. In 1998 the SCS monsoon onsets rst in the north part and then propagates southward, therefore, the onset date for the entire SCS is later than that in the region north of 7.5 N where the averages are made. The onset dates can be found also from the time evolutions of the differences of divergence elds between the higher and the lower levels. Fig. 2 shows the divergence variations with time in 1998 from May 1 to August 31 at the 850 hpa (solid lines) and the 200 hpa (dashed lines) levels averaged over the SCS (a), the IP (b), the BB (c), the ICP (d), and the AS (e). The rst maximum difference of the elds with divergence at the higher level and convergence at the lower level can be de ned as an index of summer monsoon onset. Then the onset date is around May 17 over the ICP and the BB, around May 23 over the SCS, around June 5 over the AS and around June 23 over the IP. However, the rst maximum of divergence difference over the BB is smaller compared with those in other places and much smaller than the second one in the same area. So the onset date over the BB should be around June 10 which is close to the onset date over the AS. It is seen from the above analysis that the summer monsoon onset begins earliest over the ICP and latest over the IP. The onset date is the second earliest over the SCS. It is earlier than that in the IP by one month. Nevertheless, Chinese meteorologists nd that the westerly wind components at lower levels before the middle May usually come from subtropics instead from tropics. Therefore, the onset of the EASM is not considered beginning from the ICP, and the onset of the SCS summer monsoon is the earliest. As it is commonly known that the ICP and the IP are both embedded in oceans, with the former by the SCS and the BB and the latter by the BB and the AS. Why are the onset dates so different? It is the core problem we shall discuss next. Fig. 2. Time evolution of the regionally averaged daily mean divergences at the 850 hpa (solid line) and the 200 hpa (dashed line) levels, respectively, a in the SCS (7.5 N 20 N, 110 E 120 E), b in the India Peninsula (7.5 N 20 N, 72.5 E 82.5 E), c in the Bay of Bengal (7.5 N 20 N, 82.5 E 97.5 E), d in the Indochina Peninsula (7.5 N 20 N, 97.5 E 110 E), and e in the Arabian Sea (7.5 N 20 N, 57.5 E 72.5 E) 3. The thermal mechanism As it is well known, the sensible heat (SH) and the latent heat (LH) uxes (unit: W/m 2 )atthe surface are the basic heat sources. The SH ux can directly heat the atmosphere above and result in the air temperature near the surface going up. The LH ux represents evaporation from the surface and moistens the air. However, the LH ux comes to effect only when condensation happens. If there is no strong moisture divergence and upward motion dominates, then the LH ux can be released and heat the local atmosphere just as the SH ux does. Therefore, in order to nd the

240 Y.-F. Qian et al. thermal mechanism, we need to analyze the spatial distributions and the time evolutions of the surface SH and LH uxes. 3.1 EOF Analysis of the surface heat elds The Empirical Orthogonal Function (EOF) can be used to decompose a time series of spatial elds to eigenvectors of certain modes and corresponding time coef cients. The former represents spatial distributive patterns of the modes and the latter the corresponding temporal variations. Therefore by use of the EOF decomposition method, we can get spatial modes and correspondent time coef cients of the SH and LH ux elds during 243 days from Jan. 1 to Aug. 31. The variances of the rst ve modes are 43%, 24%, 7%, 2% and 2% for SH ux and 79%, 7%, 2%, 1% and 1% for LH ux, respectively. It is seen that the sum of variances of the rst two modes is 67% for SH and 86% for LH, respectively. So the rst two spatial vectors can represent the basic features of the surface heating elds. Figure 3 shows the spatial distributions of the rst mode vectors and the correspondent time coef cients. The values in Fig. 3a, b are not the surface uxes themselves, however, they represent the relative magnitudes of the uxes by taking into account the correspondent time coefficients. Hence we still use SH and LH to denote them next for convenience. It is found from Fig. 3a that in land, one of the maximum SH ux areas is located around 40 N and between 90 Eand 110 E, another maximum SH ux area at 60 E, 30 N which southeastward expands to the India subcontinent. South to the rst area and north to the second area there are minimum SH ux areas. Hence, over the area, east of 90 E, the SH ux is larger in the north than in the south, while over the area, west of 90 E, the case is reversed. In ocean, the SH ux is smaller than that in land and distributes uniformly. Figure 3b shows that the large LH ux areas locate in ocean with maximum uxes over the north AS, the BB, the SCS and the western Paci c. In land, the LH uxes are smaller with minimum areas in coincidence with the maximum areas of SH. From time coef cients (Fig. 3c, d) we can see that all SH and LH uxes are positive, that means the SH and LH uxes are Fig. 3. The distributions of the SH and LH rst spatial vectors and the correspondent time coef cients; a SH vector, b LH vector, c SH time coef cient, d LH time coef cient

A possible mechanism effecting the earlier onset of southwesterly monsoon 241 both from the surface to the atmosphere all the time from Jan. to Aug. 1998. Therefore the spatial distribution of the rst SH mode is generally favorable for the temperature over land to increase. While the rst spatial mode of the LH ux is generally favorable for the temperature over ocean to increase. The SH uxes increase from Jan. to June and then decrease, while the LH uxesdecreasefromfeb.tomayandthenincrease with small amplitude low frequencies. The period from Feb. to May is the key brewing stage of monsoon because the summer monsoon does not onset yet. After the late April and the early May, the deep convection starts over the ICP. The fact of the small LH amount in the brewing stage and the large LH amount after reveals that the LH uxes are not as important as the SH uxes and are only the companions of the summer monsoon process. Hence, in order to discuss the mechanism of monsoon onset, it is necessary to pay more attention to the SH uxes and their time variations than to the LH uxes. The second modes explain much smaller variances of the total heat uxes than the rst ones, especially for the LH ux. Therefore, they can be referred as the additional turbulence and complements of the rst modes. Figure 4 indicates the distributions of the second spatial vectors and the correspondent time coef cients. It is seen from the spatial patterns of SH (a), and LH (b) uxes that they are indeed complementary patterns to the rst ones (see Fig. 3a, b). The patterns re ect the obvious contrast between land and sea just as in the rst mode patterns. The values of the SH mode are negative roughly south of 30 Nandpositive north of 30 N. The values of the LH mode are negative over the tropical oceans and positive over the land. The correspondent time coef cient of the SH ux is negative before the beginning of May (Fig. 4c). So the sensible heat over the tropical subcontinents and the western Paci c has positive contribution to atmospheric heating, and so does the latent heat over the AS, the BB, the SCS and the western Paci c before the early April (Fig. 4d). Therefore, the atmosphere is heated mainly with the sensible heat of the rst mode at the middle and the higher latitudes as well as with the sensible and the latent heat in tropics and subtropics to some extent before the onset stage. At the beginning of May, the time coef cient of the second SH mode changes from negative to positive. The composite effect of the rst and Fig. 4. The distributions of the SH and LH second spatial vectors and the correspondent time coef cients. a SH vector, b LH vector, c SH time coef cient, d LH time coef cient

242 Y.-F. Qian et al. thesecondshmodeistoenlargethethermal contrast between the middle and the low latitudes. The north-south heating gradients enhance, especially, over the region from 90 E to 120 E. From the early April the time coef cient of the second LH mode also changes from negative to positive. So the effect of the second LH mode is to offset the heating effect of the rst LH mode resulting in reduced north-south gradients of the LH ux. From above discussion we may conclude again that the summer monsoon onset mainly depends on the sensible heating and its north-south gradient. 3.2 Time-longitude and time-latitude evolutions of the surface sensible heat uxes The above analyses indicate that the large-scale thermal contrast between land and sea result in differential atmospheric heating and therefore induce circulation changes. Both the tropical and the subtropical heating elds exert important effects on the summer monsoon onsets. However, the rst two EOF decomposed modes are only the approximate representatives of the elds. Therefore, in this subsection we are going to analyze time evolutions of the entire surface sensible heat uxes in different latitudinal and different longitudinal belts in order to further reveal the possible mechanism of earlier onset of the SCS monsoon. Figure 5 shows the time-longitudinal (a, b), and the time-latitudinal (c, d) variations of the SH uxes. The SH uxes in Fig. 5a, b are the averages in the 10 N 20 N (a), and the 27.5 N 40 N (b) latitudinal belts, respectively, with the former representing the tropical heating and the latter the subtropical and mid-latitudinal heating. The uxes in Fig. 5c, d are averages in the 80 E 100 E (c), and the 110 E 120 E (d) longitudinal belts with the former representing the heating in the IP and the BB areas and the latter in the SCS area, respectively. It is seen from the time-longitudinal evolution of the SH ux in Fig. 5a that in the tropics, the large values of the SH are located at 70 E 80 E (the south part of the IP) and 95 E 110 E (the ICP and the west part of the SCS). Other places are occupied by small SH uxes. Therefore, in the tropics the local land-sea thermal contrast is also quite evident as it is in the subtropics. From Fig. Fig. 5. Time-longitudinal (a, b) and time-latitudinal (c, d) variations of the SH uxes averaged; a in the 10 N 20 N latitudinal belt, b in the 27.5 N 40 N belt, c in the 80 E 100 E longitudinal belt, and d in the 110 E 120 E belt

A possible mechanism effecting the earlier onset of southwesterly monsoon 243 5b it is found that between 70 E and 100 Ethe SH uxes are small, even negative before the end of April, while between 100 E and 120 Ethe uxes are large and get larger and larger day by day. Along 100 E there are large west-east gradients of the SH ux which re ect the thermal contrast between plains and high lands as pointed in the introduction. Along 120 E there are also evident west-east gradients of the SH ux re ecting the thermal contrast between land and ocean. Combined with Fig. 5a, we can see that in the longitudinal belt containing the IP large positive SH uxes exist before mid June in the south and small even negative SH uxes before early May in the north. While in the longitudinal belt containing the ICP and the SCS the SH uxes are both large in the south and in the north. Those different spatial north-south distributions and time variations of the SH uxes in the two areas may be the main cause resulting in the earlier onset of monsoon in the SCS than in the IP. Besides, it is also seen from Fig. 5a, b that the monsoons onset both in the IP area and in the SCS area after the SH ux becomes larger in the mid latitudes than that in the tropical regions. In the SCS longitudinal belt, it is after mid May and in the IP belt, it is about in late June. From Fig. 5c we can see that in 80 E 100 E longitudinal belt the main heating takes place before mid May in the tropics with small even negative SH uxes north of 30 N representing less heating in the TP and larger heating south of it. While the main heating area moves to the middle latitudes around 40 Ngraduallyaftermid April due to the increasing heating over the TP and north of it. In late June the subtropical heating becomes dominant and the north-south gradient along 30 N of the SH uxes almost reaches its maximum value, and just at that time the Indian monsoon onsets. In 110 E 120 E belt (Fig. 5d) the case is somewhat similar, but the main heating in the tropics is not so evident and important as that in 80 E 100 E. Only before March it is larger than that in the middle latitudes. After then till early May a large heating area appears between 30 Nand40 N. It abruptly moves northward to the latitudes centered at 40 N. The northsouth gradient along 30 N of the SH uxes reaches its maximum value after mid May, and the SCS monsoon onsets. The above features of SH uxes can be more clearly seen from their time evolutions in some typical regions. Figure 6a shows the SH time variations averaged in 110 E 120 E, 30 N Fig. 6. Time variations of the regional mean surface sensible heat uxes; a in 110 E 120 E, 30 N 50 N (solid curve) and 10 N 30 N (dashed curve), b in 80 E 100 E, 30 N 50 N (solid curve) and 10 N 30 N (dashed curve), c in 27.5 N 40 N, 110 E 120 E (solid curve) and 120 E 130 E (dashed curve), and d in 27.5 N 40 N, 80 E 100 E (solid curve) and 100 E 110 E (dashed curve)

244 Y.-F. Qian et al. 50 N (solid curve) and 10 N 30 N (dashed curve), the north-south difference of the SH in the SCS area may be seen from this diagram. Figure 6b is the same as Figure 6a but averaged in 80 E 100 E, showing the north-south difference of the SH in the IP area. The west-east differences of the SH at mid latitudes (27.5 N 40 N) are shown in Fig. 6c, 6d. The former indicates the local thermal contrast between land (110 E 120 E, solid curve) and sea (120 E 130 E, dashed curve) and the latter the contrast between the Tibetan Plateau (80 E 100 E, solid curve) and plains to its east (100 E 110 E, dashed curve). By comparison of Fig. 6a with Fig. 6b, it is seen that the subtropical heating becomes dominant much earlier in the SCS area than in the IP area. A comparison of Fig. 6c with Fig. 6d reveals that the west-east gradient of the SH changes sign much earlier along 120 E than along 100 E. From above analyses we may conclude that the fundamental mechanism resulting in the earlier onset of the SCS monsoon than that of the Indian monsoon is the different spatial distribution and the time evolution of the SH ux in the above two areas. In the SCS area the subtropical SH uxes are the dominant ones almost at all the time except in Jan. and Feb., while in the IP area the tropical heating is the dominant one and the subtropical heating is small even negative before May. The north-south gradient of the SH uxes along 30 N in the SCS area changes its sign much earlier than that in the IP area. Therefore, although the SCS and the Indian summer monsoons are both the tropical monsoons, the heating effects of the SH uxes at the middle latitudes and their north-south gradients along 30 N are very contributive to them. The thermal contrasts between the TP and the plain land east to it as well as between the east China plain lands and the western Paci c are also important. This conclusion can be also gotten from the spatial distributions of the SH uxes during different periods and from the analyses of the atmospheric heating elds Q 1 and Q 2 (see Wang and Qian, 2000). Numerical simulations of the mean July monsoon development and diurnal changes of weather and climate made by Kuo and Qian (1981, 1982) also showed the importance of the heating effect over the TP and the thermal contrast between the TP and the plain lands. For saving space we shall not discuss this here. 4. The dynamic mechanism The dynamic mechanism can be evidently found from circulation evolutions at the low-level during different periods. Zhu et al. (1986) has already studied the circulation differences between the EASM and the Indian monsoon and their interactive impacts by use of observations in early stage. However, the thermal and the dynamic mechanisms are usually connected to and can not be separated from each other. Therefore, when we analyze the dynamic mechanism we have to mention the thermal cause sometimes, too. Figure 7 shows the streamline elds at the 850 hpa level on individual days before, during and after the SCS monsoon onset. On May 13 (Fig. 7a), the SCS monsoon has not onset yet. Over the SCS there are anticyclonic streamlines which are the edge currents of the subtropical high over the western Paci c. A strong cyclone exists between the equator and 10 N near the south tip of the IP. Due to the cyclonic circulation and the blocking effect of the TP the streams, west of the north part of the IP, are northerly or northwesterly. On May 15 (Fig. 7b), the stream pattern changes. Due to the increasing heating over the IP and the TP the cyclone moves northward to the south part of the IP. The tropical westerly coming from the south hemisphere ows to the south part of the ICP and the west part of the SCS. The anticyclonic streamlines of the subtropical high withdraw eastward. However, over the northwest part of the IP the streams are still from the north or the northwest. This day may be taken as the foregoing stage of the SCS monsoon onset. On May 23 (Fig. 7c), the case changes a lot. The cyclone has been weakened and the tropical westerly coming through the equator at 50 E and 115 E controls almost entire area of the SCS. Therefore, May 23 can be taken as the onset date of the SCS monsoon, this date is a little later than the onset date of the 28th pentad (May 16 20) in the area north of 7.5 N. At this time, however, the IP is still dominated by the northwesterly. The circulation patterns on later days continue to change (Fig. 7d± f). We notice that when and only when the cyclone diminishes, becomes a low trough and moves gradually to the southwest part of the TP, the streamlines over the IP change to the westerly and induce the onset of the Indian monsoon. The northwestward movement of the low trough is

A possible mechanism effecting the earlier onset of southwesterly monsoon 245 Fig. 7. Streamline elds at 850 hpa level on a May 13, b May 15, c May 23, d May 26, e June 5, and f June 10 resulted from the increasing heating effect of the TP undoubtedly. Without the increasing heating over the TP the heat low there can not develop and move westward. The blocking effect can not be destroyed, either. From above analyses, we may conclude that the dynamic blocking effect of the Plateau combined with its increasing heating effect also plays an important role in the later onset of the Indian monsoon. The above circulation changes can induce changes of the vertical motion elds. We average the daily vertical p-velocities over the SCS and the IP areas, and give the results in Fig. 8, showing the time variations of the daily average vertical motions at various levels. It is seen that in the SCS area (Fig. 8a) the organized ascending motions begin on May 15 and reach the rst maximum in the whole vertical atmospheric column on May 23 when the SCS monsoon onsets. From that day on the vertical upward motions have time oscillations showing recoveries Fig. 8. Time-pressure variations of the daily area mean vertical p-velocities over the SCS a, and the India Peninsula b. Unit: 10 2 Pa/s

246 Y.-F. Qian et al. and breaks of the SCS monsoon. There are evident 10 day and 30 day low frequency oscillations. The upward motion related to three monsoon recoveries lasts about one month and the downward motion related to monsoon breaks only lasts about 10 days. In the IP region (Fig. 8b) the descending motions dominate before June 16 and on that day the ascending motions abruptly take place in the whole vertical atmospheric column showing the onset of the Indian monsoon. After June 16 the upward motions persist with time oscillations, however, no obvious breaks of the Indian monsoon exist. It is also seen that the upward motions over the SCS during the monsoon season are much larger than that over the IP, especially in the early developing stages of monsoon. The above analyses point out again that the SCS monsoon onset is about one month earlier than that of the Indian monsoon. This conclusion is coincident with that from Fig. 7. To further discuss the dynamic mechanism of the earlier onset of the SCS monsoon, the meridional distributions of the pentad mean vertical p-velocity along 115 E (through the SCS) and along 75 E (through the IP) are given in Fig. 9 and Fig. 10, respectively. It is seen that in the SCS longitude in the pentad of May 1 to 5 (Fig. 9a), the ascending motion locates over the equatorial region and the mainland between 20 Nand30 N. Then the ascending branch over the equator moves northward gradually and increases in intensity. Up to the pentad of May 21±25 (Fig. 9b), the ascending motion reaches the SCS region between 10 Nand 20 N and connects to that over the Asian mainland. At the same time, the ascending motion over the equatorial region weakens and the weak sinking motion appears in the lower levels. After the onset of the SCS monsoon, the ascending motion over the SCS weakens and moves southward slightly, while its center is still over the subtropical region of the Northern Hemisphere. Local sinking motion appears in the lower levels. On June 11±15 (Fig. 9c), the equatorial region returns to be an ascending area. The downdraft prevails at the lower levels over the SCS and the updraft at the middle and the upper levels. The Hadley cell in the Southern Hemisphere moves completely to the north side of the equator and forms the so-called vertical monsoon circulation in the SCS area. Fig. 9. Meridional circulation along 115 E (the topography shaded) averaged a before the monsoon onset (May 1±5), b during the onset of the SCS monsoon (May 21±25) and c during the onset of the Indian monsoon (June 11±15). Unit: v (m/s),! (10 2 Pa/s) The IP is situated around 75 E and the changes of the meridional circulation (Fig. 10) lag after that over the SCS. Before the monsoon onset (Fig. 10a), a Hadley circulation appears over the subtropical region in the Southern Hemisphere and its ascending branch mainly lies in the tropical region south to the equator. There is another Hadley circulation at the upper levels north to the equator, which causes the air in the lower levels to ow

A possible mechanism effecting the earlier onset of southwesterly monsoon 247 Fig. 10. Meridional circulation along 75 E (the topography shaded) averaged a before the monsoon onset (May 1±5), b during the onset of the SCS monsoon (May 21±25), and c during the onset of the Indian monsoon (June 11±15). Unit: v (m/s),! (10 2 Pa/s) from the Northern Hemisphere to the Southern Hemisphere. This shows that the meridional belt along 75 E is still characterized by the typical vertical circulation in winter at this stage. Up to May 21±25 (Fig. 10b), the meridional circulation changes. The Southern Hemisphere Hadley circulation dominates in the upper atmosphere between the equator and 20 S, with its northward expansion the ascending branch moves northward signi cantly and the wind directions in the lower and middle layers over the equator reverse, owing from the Southern to the Northern Hemisphere in turn. The north Hadley circulation shrinks to a narrow one but still maintains over the Northern Hemisphere. Its sinking branch south to the TP controls the most subtropical region of the Northern Hemisphere. Therefore, the circulation pattern over the IP is still in the transition period from winter to summer during the onset of the SCS monsoon. Not until June 11±15 (Fig. 10c) does the pattern change radically. During June 11 to 15, the Southern Hemisphere Hadley cell moves further north to the equator with sinking in the south and ascending in the north, opposite to the pattern before that stage. At the same time, an ascending ow dominates most regions south to the Plateau as well. The Indian monsoon bursts out. By examining Fig. 9 and Fig. 10 together, we can nd that the seasonal variation of the meridional circulation near the equator exhibits remarkable regional features. Along the longitude of 115 E, the SCS is roughly located from 7.5 N to 20 N and the east China is to the north. The topography there is comparatively smooth and favorable to the seasonal adjustment of meridional circulation. However, at 75 E, the TP is there in the north and blocks the air ow from south. Strong downdrafts, south to the Plateau, block the northward shift of the upward motions and therefore delay the seasonal adjustment of meridional circulation. This clearly presents the dynamic aspect of the in uence of the TP topography on the onset time of monsoon, too. 5. Conclusions and discussions Based on the above analyses, we can conclude this study as follows: During the onset of the SCS monsoon, the meteorological element elds, the horizontal and the vertical circulation change signi cantly. By taking the appearance of the southwesterly wind in the lower layers and the easterly wind in the upper layers as the indicator of the SCS monsoon onset, the SCS monsoon onsets roughly on May 23, 1998, which is a little later than normal. The Indian monsoon bursts out around June 16. The time variations of the daily mean divergence elds in the upper and the lower layers are consistent with the above monsoon process.

248 Y.-F. Qian et al. From the heating elds we nd that in the SCS area, the surface sensible heating exists almost all the time at both the tropics and the subtropics, and the subtropical heating becomes dominant from March. The north-south gradients of the SH uxes at 30 N change their sign completely from March, reach maximum values in mid May and result in the earlier burst of the SCS summer monsoon. While in the IP area, the tropical heating is the dominant one and the subtropical heating is small even negative at the early period before May. The north-south gradients of the SH uxes do not change their sign completely until May. Therefore, the heating effects of the SH uxes at the middle latitudes and their north-south gradients at 30 N are critically important to the onsets of monsoon. It is further found that the west-east gradients of the SH uxes at middle latitudes between the northwestern Paci c and the east Asian continent and between the east plains and the TP are also favorable to the earlier onset of the SCS summer monsoon. The SCS monsoon onsets after the west-east gradients between the northwestern Paci c and the East Asian continent change sign. And the Indian monsoon onsets after the west-east gradients between the east plains and the TP change sign. The time evolutions of circulation indicate the dynamic mechanism and are closely connected to the time changes of heating eld. The change of a strong cyclone, which forms between the equator and 10 N at the south tip of the IP, can be taken as an indicator of the summer monsoon developing process. Due to the favorite location of the SCS in the front of the cyclone the southwesterly coming from the tropics easily reaches the SCS area. The IP locates in the northwest rear part of the cyclone. The combined impact of the cyclone and the blocking effect of the TP makes the northwesterly stronger and lasting a long time over the IP area. When and only when the cyclone moves towards the TP, becomes weaker and weaker and forms a shallow trough due to the increasing heating effect of the TP, the southwesterly monsoon ow comes to the IP area and the blocking effect of the TP is weakened. It is at that time that the Indian summer monsoon bursts out and is hence later than the SCS summer monsoon onset by about one month. The local conditions of topography can also in uence the onset time of monsoon. In the SCS area, there is no high north-south contrast of topography, and the Hadley cell from the Southern Hemisphere can faster reach to the SCS and induce upward motions over that area causing large amount of latent heat to release and heat the atmosphere. While in the IP, the large and high TP to the north blocks the Hadley cell to come. Moreover, the strong downward motions on the south slope of the TP dominate which again block the upward motions from the south to enter the area. Such condition does not change until time when the heating over the TP increases to some extent and attracts the monsoon low to locate near it. Therefore, the circulation change and the blocking effect of the topography are also the cause of the earlier onset of the SCS monsoon though, the fundamental mechanism should be found in the heating effect of the underlying surface. References Chen LX, Zhu QG, Luo HB et al. (1991) The East Asian Monsoon. Beijing China: China Meteorological Press, 362 pp (in Chinese) Ding YH, Murakami M (1994) The Asian Monsoon. Beijing China: China Meteorological Press, 263 pp He H, McGinnis JW, Song Z, Yanai M (1987) The onset of the Asian monsoon in 1979 and the effect of the Tibetan Plateau. Mon Wea Rev 115: 1966±1995 Krishnamurti TN (1985) The summer monsoon experiment, a review. Mon Wea Rev 113: 1590±1626 Kuo HL, Qian YF (1981) In uence of the Tibetan Plateau on cumulative and diurnal changes of weather and climate in summer. Mon Wea Rev 109 (11): 2337±2356 Kuo HL, Qian YF (1982) Numerical simulation of the development of mean monsoon circulation in July. Mon Wea Rev 110(12): 1879±1897 Lau K-M, Yang S (1996) The Asian monsoon and predictability of the tropical ocean-atmosphere system. Q J Roy Meteor Soc 122: 945±957 Lau K-M, Yang S (1997) Climatology and interannual variability of the Southeast Asian monsoon. Adv Atmos Sci 14: 141±162 Ma HN, Ding YH (1997) The present status and future of research of the East Asian monsoon. Adv Atmos Sci 14: 125±140 Shen S, Lau K-M (1995) Biennial oscillation associated with the East Asian summer monsoon and tropical sea surface temperatures. J Meteor Soc Japan 73: 105±124 Ramage CS (1971) Monsoon Meteorology. New York and London: Academic Press, 296 pp Tao SY, Chen LX (1987) A review of recent research of the East Asian summer monsoon in China. Monsoon Meteorology. Chang CP, Krishnamurti TN (eds) Oxford University Press: 60±92

A possible mechanism effecting the earlier onset of southwesterly monsoon 249 Wang B, Wu RG (1997) Peculiar temporal structure of the South China Sea summer monsoon. Adv Atmos Sci 14: 177±194 Wang SY, Qian YF (2000) Diagnostic study of apparent heat sources and moisture sinks in the South China Sea and its adjacent areas during the onset of 1998 SCS monsoon. Adv Atmos Sci, 17: 285±298 Webster PJ, Yang S (1992) Monsoon and ENSO: Selectively interactive systems. Q J Roy Meteor Soc 118: 877± 926 Ye T, Gao Y (1979) The Tibetan Plateau Meteorology. Beijing China: China Science Press, 278 pp (in Chinese) Zhu Q, He J, Wang P (1986) A study of the circulation differences between East Asian and Indian monsoons with their interaction. Adv Atmos Sci 3: 466±477 Authors' address: Y.-F. Qian, S.-Y. Wang and H. Shao, Nanjing University, Department of Atmospheric Sciences, 11 Hankou Road, Nanjing P.R. China, 210093 (E-mail: qianzh2@netra.nju.edu.cn)