A possible mechanism for the co-variability of the boreal spring Antarctic Oscillation and the Yangtze River valley summer rainfall

INTERNATIONAL JOURNAL OF CLIMATOLOGY Int. J. Climatol. 29: 1276 1284 (2009) Published online 3 December 2008 in Wiley InterScience (www.interscience.wiley.com).1773 A possible mechanism for the co-variability of the boreal spring Antarctic Oscillation and the Yangtze River valley summer rainfall Jianqi Sun, a * Huijun Wang a and Wei Yuan b a Nansen-Zhu International Research Centre (NZC), Institute of Atmospheric Physics, Chinese Academy of Sciences, PO Box 9804, Beijing, China b China Meteorological Administration Training Centre, Beijing, China ABSTRACT: A significant correlation between the boreal spring Antarctic Oscillation (AAO) and the Yangtze River valley summer rainfall (YRVSR) has been found in previous studies, although the mechanisms that might lead to such far-reaching teleconnection remain unresolved. In this study, one of possible mechanisms responsible for the co-variability of the boreal spring AAO and the YRVSR is proposed. It follows that the convection activity over the region of the Maritime Continent serves as a bridge linking the boreal spring AAO and the YRVSR. This physical process can be described schematically as follows: during the boreal spring, a positive-phase (negative-phase) AAO is concurrent with a strong (weak) convection activity over the region of the Maritime Continent via anomalous meridional circulations along the central South Pacific and two meridional teleconnection wave train patterns, with one over the southern Indian Ocean at the lower level and the other along the central South Pacific at the upper level. Thereafter, the anomalous convection propagates northward along the seasonal cycle, and then changes the western Pacific subtropical high in the following seasons, consequently impacting on the summer rainfall in the Yangtze River valley. Copyright 2008 Royal Meteorological Society KEY WORDS Antarctic Oscillation; precipitation; western Pacific subtropical high Received 14 July 2007; Revised 15 December 2007; Accepted 17 September 2008 1. Introduction The interannual variability of the East Asian summer monsoon (EASM), associated with frequent occurrences of extreme droughts/floods, has substantial influences on society. Many studies have therefore been carried out to learn about the EASM rainfall variability (e.g., Tao and Chen, 1987; Ding and Chan, 2005). It is found that the variability of the EASM rainfall is very complicated and attributed to many reasons. First, the large-scale circulations of the Northern Hemisphere have strong impacts on the EASM. Particularly, the western Pacific subtropical high (WPSH) (e.g., Chang et al., 2000), East Asian jet (Liang and Wang, 1998; Lau et al., 2000), Arctic Oscillation (Gong and Ho, 2003), and blocking over the Okhotsk (Zhang and Tao, 1998) to a large extent determine the EASM variability. In addition, the sea surface temperature (SST) anomalies over the tropical and northern midlatitude oceans are also documented having close linkages with the EASM. For example, Wang et al. (2000a) and Wang et al. (2000b) demonstrated the impact of the tropical eastern Pacific SST on the EASM. Huang and Li (1988) as well as Li and Mu (2001) addressed that the tropical western Pacific * Correspondence to: Jianqi Sun, NZC, Institute of Atmospheric Physics, Chinese Academy of Sciences, PO Box 9804, Beijing 100029, China. E-mail: sunjq@mail.iap.ac.cn and Indian Ocean play important roles in the variability of the EASM. Ueda and Kawamura (2004) and Li and Bates (2007) pointed out that the SST anomalies over the North Pacific and North Atlantic should also be emphasized in the study of the East Asian monsoon. More recent studies have examined the analogous relationship between the EASM rainfall and the Southern Hemispheric circulations, especially the relationship between the boreal spring Antarctic Oscillation (AAO) and the Yangtze River valley summer rainfall (YRVSR) (Gao et al., 2003; Xue et al., 2004; Wang and Fan, 2005; Fan, 2006). The AAO is the dominant annular mode in the extratropical Southern Hemisphere. It features primarily a large-scale seesaw oscillation of atmospheric circulations between the mid latitudes and high latitudes of the Southern Hemisphere (Rogers and van Loon, 1982; Kidson, 1988; Gong and Wang, 1999; Thompson and Wallace, 2000). The impacts of the AAO on the Southern Hemispheric climate have been extensively examined in numerous studies. It follows that the AAO-related variations may involve the Southern Hemispheric total column ozone, tropopause height over middle and high latitudes, strength of trade winds (Thompson and Wallace, 2000), sea ice and ocean conditions (Hall and Visbeck, 2002), Antarctic surface air temperature (Kwok and Comiso, 2002), precipitation over southeastern South America (Silvestri and Vera, 2003) and over western South Africa (Reason and Rouault, 2005). In addition, more recent Copyright 2008 Royal Meteorological Society

MECHANISM FOR CO-VARIABILITY OF AAO AND YRVSR 1277 study indicated that the Australian rainfall and surface temperature are also related to the AAO (Hendon et al., 2007). The AAO dominates the extratropical Southern Hemisphere; therefore it has a direct impact on the Southern Hemispheric climate. However, the influence of AAO on the Northern Hemispheric climate is complicated since the two are far apart. Fan and Wang (2004) suggested that the meridional teleconnection from the Antarctica to the Northern Hemisphere in the boreal spring might play a role bridging AAO and East Asian dust weather frequency. However, it is still unclear as to which physical process is responsible for the lagging correlation between the boreal spring AAO and the YRVSR. Hence, the aim of the current work is to explore the possible physical mechanism for such far-reaching teleconnection. This paper is outlined as follows. Datasets are described in Section 2. The results are shown in Section 3. The discussion and conclusions are given in Sections 4 and 5, respectively. 2. Datasets The datasets employed in this study are as follows: (1) The National Centers for Environmental Prediction/National Center for Atmospheric Research (NCEP/NCAR) global atmospheric reanalysis data (Kalnay et al., 1996; Kistler et al., 2001) are used to investigate the features of atmospheric circulations associated with AAO and YRVSR. (2) The outgoing long-wave radiation (OLR) data are used to infer tropical convection. The OLR data are available from June 1974 with a missing period between March and December in 1978. (3) Two monthly precipitation datasets are used in this study. One is China s 160-station precipitation, and the other is the Climate Prediction Center Merged Analysis of Precipitation (CMAP) (Xie and Arkin, 1997). The NCEP/NCAR reanalysis data, OLR data, and CMAP data are all obtained from the National Oceanic and Atmospheric Administration (NOAA) Climate Diagnostic Center (CDC) and they are all gridded on a 2.5 2.5 mesh. China s 160-station precipitation data are provided by the China Meteorological Administration. 3. Results Significant positive correlation between the boreal spring AAO and the YRVSR has been reported in several previous studies (e.g., Gao et al., 2003; Xue et al., 2004). Here we only show the time series of the normalized boreal spring AAO index and the YRVSR index to depict their connection. The AAO index is defined by the leading principal component of Southern Hemispheric sea level pressure (SLP) anomalies (south of 20 S). A positive-phase (negative-phase) AAO index is characterized by negative (positive) SLP anomalies over Antarctica and positive (negative) anomalies over the Southern Hemispheric mid latitudes (Gong and Wang, 1999; Thompson and Wallace, 2000). The YRVSR index is referred to as the mean of the summer rainfall of 18 stations. These 18 stations are located along the Yangtze River valley, and their rainfalls are all significantly correlated with the boreal spring AAO index at 95% confidence level in the period 1958 2004 (estimated by a local student s t-test). Thus, the averaged rainfall of these 18 stations is used to represent the mean value of the rainfall of the whole Yangtze River valley. Immediately evident in Figure 1 is a strong increasing trend exhibited in the AAO index, with negative-phase in the 1960s and positive-phase in the 1990s. Correspondingly, the YRVSR also shows an increasing trend with less rainfall in the 1960s and more rainfall in the 1990s, although its linear trend is weaker than that of the AAO index. These two time series are correlated with a value of 0.53 in 1958 2004, which is significant at 99% confidence level. Besides, considering the influence of the strong positive trend in these two indices, the correlation of the detrended indices is then calculated. The correlation is now decreased, but the coefficient with a value of 0.37 is still significant at 95% confidence level. Thus, it can be concluded that there is a close relationship between the boreal spring AAO and the YRVSR. While this close connection has been established, the cause-and-effect relationship between the AAO and YRVSR has not been determined yet. The purpose of the present study is to address this question. Because of concerns with data quality prior to the incorporation of global satellite soundings, particularly in the Southern Hemisphere (Kistler et al., 2001), we restrict our analysis to the 26 years following 1978 (1979 2004) to get a more reliable result. In this period (1979 2004), the AAO index is also significantly correlated with the YRVSR index with a coefficient of 0.48 (significant at 95% confidence level). Here, the correlation and composite statistical methods are used. Based on the criterion that the AAO index is larger than 1 (less than 1) standard deviation, the positive-phase AAO years YRVSR Index AAO and YRVSR Indices 300 250 200 150 100 50 0 1958 1964 1970 1976 1982 1988 1994 2000 Year Figure 1. Time series of the normalized boreal spring AAO index (solid curve) and the YRVSR index (dashed curve, in mm month 1 ) during the period 1958 2004. The linear trends of AAO and YRVSR are indicated by the solid and dashed lines respectively. 3 2 1 0-1 -2-3 Normalized AAO Index

1278 J. SUN ET AL. (1979, 1982, 1989, 1998, 1999, and 2000) and negativephase AAO years (1980, 1981, 1986, 1990, 1992, and 2002) are tagged. (a) 700 hpa in the boreal spring 3.1. How does the AAO affect the tropical climate systems in the boreal spring? AAO occurs in the Southern Hemispheric middle to high latitudes; how can it co-vary with the YRVSR? Naturally, it is speculated that some tropical climate systems may serve as a bridge. Thus we start with the analysis of the tropical variations associated with the AAO. Since the anomalies associated with positive-phase and negative-phase AAO tend to have opposite polarities, the composite difference will be shown hereafter. Figure 2(a) shows the composite differences of 700- hpa horizontal wind between positive-phase and negativephase AAO years as well as the correlation of 700 hpa geopotential height with AAO index in the boreal spring. It follows that the boreal spring AAO is closely related to the Southern Hemispheric atmospheric circulations. With a positive-phase AAO, the zonal wind over the Southern Hemispheric high latitudes is enhanced and accompanied by a deepening of Antarctic circumpolar low. On the other hand, the zonal wind over the Southern Hemispheric Subtropics is depressed, co-varying with strengthened subtropical highs over the southern Indian Ocean and the southern Pacific to southeastern Australia (indicated by AI 700 and AP 700 in Figure 2(a), respectively). Such a reverse variation of atmospheric circulations between the middle latitudes and high latitudes of the Southern Hemisphere is a major feature of AAO (Gong and Wang, 1999; Thompson and Wallace, 2000). In addition, there is a significant meridional teleconnection wave train pattern, emanating from AI 700 northeastward to the Tropics with another centre at the northwestern seacoast of Australia (CI 700 ) (Figure 2(a)). This teleconnection pattern links AAO to the tropical climate at the lower level, resulting in a strong anomalous westerly over the tropical eastern Indian Ocean. Such an anomalous westerly converges with an anomalous easterly over the tropical western Pacific, leading to a strong convergence over the Maritime Continent in the boreal spring. Figure 2(b) shows the composite differences of 150 hpa horizontal wind between positive-phase and negative-phase AAO years as well as the correlation of 150 hpa geopotential height with AAO index in the boreal spring. Evidently, a significant meridional teleconnection wave train pattern is present along the central South Pacific. It includes three centres: AP 150, CP1 150, and CP2 150. This teleconnection wave train pattern bridges the AAO and the tropical circulations at the upper level, resulting in an anomalous westerly over the tropical western Pacific. Since the lower level and upper level zonal winds are inversely related in the Tropics (Webster and Yang, 1992), this upper level anomalous westerly matches well with the pronounced anomalous easterly at the lower level over the tropical western Pacific (Figure 2(a)). (b) CI 700 AI 700 AP 700 150 hpa in the boreal spring CP2 150 CP1 150 AI 150 AP 150 Figure 2. Correlations of (a) 700 hpa and (b) 150 hpa geopotential heights with AAO index (contours), superimposed with composite differences of horizontal wind (vectors, in m s 1 ) between positive-phase and negative-phase AAO years in the boreal spring. Areas with significant correlations at 95% (90%) confidence level are shaded dark (light) (estimated by a local student s t-test). A ( C ) indicates anticyclone (cyclone). I and P denote the regions over the Indian Ocean and the Pacific. 700 ( 150 ) indicates the 700 hpa (150 hpa). On another aspect, regional meridional circulation change in the atmosphere is also an important factor contributing towards the polar-tropical connection (e.g., Liu et al., 2002). Bals-Elsholz et al. (2001) documented that there is a high correlation between the Southern Hemispheric split jet and AAO. In the scenario of a positive-phase (negative-phase) AAO, the anomalous westerly (easterly) zonal wind near 60 S over the central South Pacific enhances (depresses) the polar front jet (PFJ), and anomalous easterly (westerly) zonal wind near 30 S depresses (enhances) the subtropical jet (STJ). This feature is well exhibited in Figure 3. The PFJ (STJ) in positive-phase AAO years tends to be stronger (weaker) than in negative-phase AAO years. As a consequence, the changes of the jets result in an anomalous local Ferrel Cell along the central South Pacific (Liu et al., 2002; Yuan, 2004). As shown in Figure 4(a), both the anomalous upward motion centred at about 35 S and anomalous downward motion centred at about 55 S tend to weaken the equatorward segment of the local Ferrel Cell during positive-phase AAO years. At the same time, an anomalous poleward air near the surface south of about

MECHANISM FOR CO-VARIABILITY OF AAO AND YRVSR 1279 (a) Positive-phase AAO years in the boreal spring (a) 180-190 E in the boreal spring (b) Negative-phase AAO years in the boreal spring Pressure Levels (hpa) (b) 80-90 E in the boreal spring Figure 3. Composites of 150 hpa zonal wind (in m s 1 ) in (a) positivephase and (b) negative-phase AAO years in the boreal spring. The lines with values less than 23 m s 1 are omitted. 30 S in positive-phase AAO years tends to enhance the poleward segment of the local Ferrel Cell. This anomalous local Ferrel Cell then changes the local Hadley Cell to its north in two ways: (1) The ascending branch of the anomalous local Ferrel Cell over the Southern Hemispheric Subtropics and its aloft northward divergent flow suppress the local Hadley Cell to its north; (2) The anomalous southward air advection south of about 30 S from the surface to the upper troposphere causes anomalous poleward heat transportation from low latitudes to high latitudes, consequently raising the tropospheric temperature in the Southern Hemispheric middle latitudes. The warming in the middle latitudes can decrease the Tropics-to-Extratropics meridional temperature gradient, consequently resulting in a weakened local Hadley Cell along the central South Pacific (Rind et al., 2001; Liu et al., 2002). It is well known that over the tropical Pacific, the dominant zonal circulation is the Walker circulation. On climatology, the air rises over the western side of the basin and moves eastward at the upper levels and sinks over the eastern side of the basin. The trade winds along the equatorial Pacific then complete the Walker circulation cell (Garcia and Kayano, 2008). The above pronounced relaxation of the local Hadley Cell over the central South Pacific produces a strong anomalous descending motion over the central-eastern tropical Pacific as depicted in Figures 4(a) and 5. This anomalous descending air, combined with the upper level anomalous westerly (Figure 2(b)) and lower level anomalous easterly (Figure 2(a)) over the tropical western Pacific, Latitudes Figure 4. Composite differences of latitude pressure cross-section of air temperature (contours, in C) and meridional circulations (vectors) averaged along (a) 180 190 E and(b) 80 90 E between positive-phase and negative-phase AAO years in the boreal spring. The unit is m s 1 for meridional wind and hpa s 1 for vertical motion. The values of vertical velocity are multiplied by 200. Areas with significant composite differences of vertical velocity at 95% (90%) confidence level are shaded dark (light) (estimated by a local student s t-test). strengthens the Walker circulation and in turn induces a significant ascending anomaly over the Maritime Continent (Figure 5). Thus, through the aforementioned two anomalous meridional wave train patterns and meridional circulations, the signal of the AAO is transported to the tropical circulations and in turn stimulates a strong upward motion over the region of the Maritime Continent. As a consequence, the convection activity is apt to be stronger over there. This hypothesis is well proved by the OLR analysis. As shown in Figure 6, the Maritime Continent is distinguished by significant negative correlations, indicating the convection is stronger than normal. One issue, regarding the impacts of the AAO on the tropical circulations and convection, needs further consideration. This issue is the impact of the El Niño- Southern Oscillation (ENSO) phenomenon on the composite anomalies developed above. The ENSO, acting as the major source of interannual variability in the Tropics, has a widely established impact on quasi-global climate. Moreover, L Heureux and Thompson (2006) and

1280 J. SUN ET AL. Pressure Levels (hpa) 15 S-15 N in the boreal spring Longitudes Figure 5. Composite differences of longitude pressure cross-section of vertical velocity (in hpa s 1 ) averaged along 15 S 15 N between positive-phase and negative-phase AAO years in the boreal spring. Areas with significant composite differences at 95% (90%) confidence level are shaded dark (light) (estimated by a local student s t-test). OLR in the boreal spring Figure 6. Correlation of OLR with AAO index in the boreal spring. Areas with significant correlations at 95% (90%) confidence level are shaded dark (light) (estimated by a local student s t-test). Carvalho et al. (2005) recently discussed that the fluctuations of the ENSO are associated with the variability of AAO in the austral summer. Although the period of analysis is during the boreal spring in this study, it still needs to ensure whether the atmospheric anomalies pronounced in this paper are closely related to the AAO or just a residual of the ENSO signal. To answer this question, we recompute the AAO-related composite differences by excluding the springs during El Niño [1983, 1987, 1992, 1993, 1995, 1998, and 2003) and La Niña [1989, 1999, 2000, and 2001; the dates are identified by Smith and Sardeshmukh (2000) and are updated online at http://www.cdc.noaa.gov/people/cathy.smith/ best/#years]. Considering the number of extreme years, the positive-phase (negative-phase) AAO years are now defined as the AAO index greater than 0.5 (less than 0.5) standard deviation. Based on the above two criteria, there are four positive-phase (1979, 1982, 1996, and 2004) and five negative-phase (1980, 1981, 1986, 1990 and 2002) AAO years. The results show that the AAO-related atmospheric and convective composite differences in the non-enso years (figures not shown) are quite similar to those depicted in Figures 2 6 when the ENSO years are included. This suggests that the results in this study are dominated by variations in the AAO and are not heavily contaminated by the impact of ENSO. 3.2. Roles of the convection over the region of the Maritime Continent The EASM is a portion of the greater Asian Australian monsoon. The evolution of the Asian Australian monsoon shows a strong seasonal cycle (Wang et al., 2002). As shown in Figure 7, during the Australian summer monsoon time (December February), the major rain belt is located south of the equator. From March, after the ending of the Australian summer monsoon, the rain belt starts propagating northward gradually. By the end of May to the beginning of June, the major rain belt reaches about 10 N. And the north portion of the rain belt covers the Yangtze River valley, implying the beginning of the Yangtze River valley rainy season. This rainfall will persist until the ending of the boreal summer (June August). Then the rain belt retreats southward from September. It returns back to the Southern Hemisphere in the austral summer (December February) and completes the seasonal cycle of Asian Australian monsoon. Accompanying the Asian Australian monsoon seasonal cycle, the pronounced anomalous convection associated with the boreal spring AAO also propagates northward from the boreal spring to summer as shown in Figure 8. The only difference is that the anomalous convection is divided into two branches in its northward propagation. This attributes to the interaction between the monsoon-related convection heating and WPSH. Some previous studies have documented that the Asian monsoon-related deep convection can strengthen Month Precipitation along 105-140 E Latitudes Figure 7. Latitude month cross-section of the climatological (1979 2004) CMAP data (in mm day 1 ) averaged along 105 140 E. The values less than 5 (mm day 1 ) are not plotted.

MECHANISM FOR CO-VARIABILITY OF AAO AND YRVSR 1281 OLR along 105-140 E 105-140 E in the boreal summer Month Pressure Levels (hpa) Latitudes Latitudes Figure 8. Composite differences of latitude month cross-section of OLR (in W m 2 ) averaged along 105 140 E between positive-phase and negative-phase AAO years. Areas with OLR values less than 3 are shaded. Figure 9. Composite differences of latitude pressure cross-section of meridional circulations (vectors) averaged along 105 140 E between positive-phase and negative-phase AAO years in the boreal summer. The unit is m s 1 for meridional wind and hpa s 1 for vertical motion. The values of vertical velocity are multiplied by 200. Areas with significant composite differences of vertical velocity at 95% (90%) confidence level are shaded dark (light) (estimated by a local student s t-test). the WPSH to its east (e.g., Webster, 1972; Rodwell and Hoskins, 2001; Liu et al., 2004). The enhanced WPSH then extends southwestward and dominating the tropical western North Pacific, thus depresses the convection in the region and in turn divides the western Pacific convection into two branches as shown in Figure 8. The northern branch is located at about 30 N to influence the YRVSR, and the southern one is blocked near the equator (Figure 8). This interactive process between the monsoon-related deep convection and WPSH will be discussed in Section 4.2. Figure 9 shows the composite differences of the local meridional circulations averaged along 105 140 E in summer. The descending motion is around 10 N, indicating an enhanced WPSH. This descending airflow combined with the ascending airflow around 30 N forms an enclosed meridional cell. That is the so-called EASM circulation cell (Ding and Chan, 2005) which can strengthen the summer rainfall in the Yangtze River valley. On the horizontal direction, an anomalous anticyclone circulation dominates the tropical western North Pacific corresponding to an enhanced WPSH (Figure 10). Such anomalous anticyclone also favours more rainfall in the Yangtze River valley through three ways: (1) It blocks the Mei-yu fronts from moving southward and thereby extends the time that the fronts produce stationary rainfall, (2) enhances the pressure gradient to its northwest resulting in a more intense front, and (3) induces anomalous warming of the South China Sea surface through increased downwelling, which leads to a higher moisture supply to the rain area (Chang et al., 2000). Therefore, the Yangtze River valley is covered by a zonally oriented belt of significant negative correlations (Figure 10). OLR and 850 hpa wind in the boreal summer Figure 10. Correlation of OLR with AAO index (contours), superimposed with composite differences of 850 hpa horizontal wind (vectors, in m s 1 ) between positive-phase and negative-phase AAO years in the boreal summer. Areas with significant correlations at 95% (90%) confidence level are shaded dark (light) (estimated by a local student s t-test). 4. Discussion 4.1. AAO-related teleconnection wave train patterns and the Southern Hemispheric climatological circulations The analysis in Section 3.1 indicates that AAO connects the tropical circulations through two meridional teleconnection wave trains, with one over the southern Indian Ocean at the lower level and the other along the central South Pacific at the upper level. The existence and location of these two meridional teleconnection wave trains attribute mainly to the Southern Hemispheric climatological circulations.

1282 J. SUN ET AL. In the boreal spring, the Mascarene High (MH) is the main subtropical system in the eastern Southern Hemispheric middle latitudes at the lower level (Figure 11(a)). A stronger AAO can strengthen the MH (Xue et al., 2004). The strengthened MH will then enhance the southeasterly at its northeast. On the way northward, the southeasterly will be defected to the right due to Coriolis force and will become the southwesterly, thus resulting in a cyclonic circulation over the region to northwestern Australia. This cyclonic circulation, combined with the MH, forms the pronounced lower level meridional teleconnection pattern over the South Indian Ocean. Thus, the existence of this teleconnection pattern is strongly related to the MH and its associated northward flows. The AAO can influence this teleconnection pattern by strengthening or weakening the MH. At the upper level, a distinct characteristic of the climatological circulations in the Southern Hemisphere is the split jet at the longitudes of Australia and New Zealand, with a strong STJ centred between 25 and 35 S, a relatively weaker PFJ centred between 55 and 65 S, and an obvious westerly minimum between the two jets (Figure 11(b)). The positive-phase (negative-phase) AAO can strengthen (weaken) the PFJ and weaken (strengthen) the STJ (Figure 3). The opposite variations of the PFJ and STJ then result in an anomalous upper level meridional teleconnection and anomalous meridional circulations (Ferrel Cell and Hadly Cell) along the central South Pacific (Liu et al., 2002; Yuan, 2004). However, over (a) (b) MH SLP in the boreal spring 150 hpa zonal wind in the boreal spring Figure 11. Climatological (1958 2004) (a) SLP (in hpa) and (b) 150 hpa zonal wind (in m s 1 ) in the boreal spring. The arrow in (a) denotes the flow path emanating from the Mascarene High. MH indicates the Mascarene High. The bold line in (b) denotes the jet core. the Southern Hemispheric region to the east of 120 E, there is no jet split and the jet core rapidly retreats to nearly 50 S. The strong changes of jet related to AAO are only over the high latitude southern Indian Ocean. Over the midlatitude and tropical southern Indian Ocean, the changes of zonal wind are weaker than those over the split jet region (Figure 2(b)). As a consequence, the anomalous meridional circulations over the southern Indian Ocean are also weaker than those along the central South Pacific (Figure 4(b)) and there is no obvious meridional teleconnection pattern over the upper level South Indian Ocean (Figure 2(b)). Thus, the existence of this upper level teleconnection pattern along the central South Pacific is strongly related to the split jet over there. The AAO can impact on this teleconnection pattern by changing this split jet. 4.2. The convection over the region of the Maritime Continent and the WPSH Some previous observational and modelling studies have shown that the deep-convective condensation heating near the equator and Asian monsoon region has important impacts on the position and intensity of the WPSH in the warm season (Webster, 1972; Rodwell and Hoskins, 2001; Liu et al., 2004). Rodwell and Hoskins (2001) and Liu et al. (2004) gave a clear physical explanation for how the Asian monsoon forces the anticyclone to its east. They pointed out that, over the Asian monsoon region, maximum deep-convective heating occurs at the height about 300 400 hpa where the heating rate can be several degrees per day. This, following the Sverdrup vorticity balance, results in a poleward flow below the maximum heating level. This forced poleward meridional wind enhances the western flank of the subtropical high to the east of the deep convection region in the middleto-lower troposphere. In addition, the strong convective heating can stimulate a Kelvin wave to the east, which then enhances the southern flank of the subtropical high. Through the above two processes, the Asian monsoon rainfall heating can strengthen the WPSH to its east. In this study, the convective activity over the region of the Maritime Continent has an important effect on the WPSH. Figure 12 shows the composite monthly WPSH in positive-phase and negative-phase AAO years and composite difference of OLR between positive-phase and negative-phase AAO years from March to June. Compared to negative-phase AAO years, the WPSH in positive-phase AAO years becomes stronger and more southwestward with time. The more obvious enhancement of the WPSH begins in April. At that time, the main convective centre propagates across the equator and is located over the tropical western North Pacific. In May, the WPSH keeps on being strengthened. On another aspect, the enhanced WPSH begins to depress the convection over the middle tropical western North Pacific, and then splits the tropical western North Pacific convection into two parts: one to its north and the other to its south. In June, the enhanced WPSH dominates the tropical western North Pacific and pushes the northern part

MECHANISM FOR CO-VARIABILITY OF AAO AND YRVSR 1283 March April May June Figure 12. Composite monthly WPSH at 500 hpa (indicated by 5875 gpm contour) in positive-phase (solid line) and negative-phase (dashed line) AAO years, superimposed with composite differences of OLR (in W m 2 ) between positive-phase and negative-phase AAO years, from March to June. Areas with OLR anomalies greater than 10 (less than 10) are shaded dark (light). of the convection to the Yangtze River valley, implying the beginning of the Yangtze River valley rainy season. Through such process, the signal of the boreal spring AAO is transported to East Asia, and in turn has an influence on the WPSH and YRVSR. This, to some extent, accounts for why the boreal spring AAO and WPSH covary in phase as shown in Section 3.2 and the previous study of Xue et al. (2004). 5. Conclusions One possible mechanism for the influence of the boreal spring AAO on the YRVSR is explored in this study. It is found that in an anomalous AAO boreal spring, there are anomalous meridional circulations along the central South Pacific and two concurrent meridional wave train patterns, with one over the southern Indian Ocean at the lower level and the other along the central South Pacific at the upper level. These two teleconnection wave train patterns and the anomalous meridional circulations connect AAO with the convective activity over the region of the Maritime Continent in the boreal spring. A positive-phase (negative-phase) AAO is associated with a strong (weak) convection over the region of the Maritime Continent. Then, the AAO-related anomalous convection propagates northward along the seasonal cycle, and in turn changes the WPSH. When the convection activity is stronger, the WPSH will be stronger and more southeastward, dominating the tropical western North Pacific. Finally, the changes of WPSH result in anomalous YRVSR. However, this mechanism may not be the whole story for the coupling between the boreal spring AAO and the YRVSR. Further explorations should be needed and may result in more valuable findings. Acknowledgements The authors are grateful to Prof. Helge Drange, Prof. Yongqi Gao, Prof. Jiping Liu, and two anonymous reviewers for their valuable comments and helpful advices. This research was jointly supported by the Knowledge Innovation Project of Chinese Academy of Sciences under grants KZCX2-YW-217 and the National Basic Research Program of China under Grant 2006CB403600. References Bals-Elsholz TM, Atallah EH, Bosart LF, Wasula TA, Cempa MJ, Lupo AR. 2001. The wintertime Southern Hemisphere split jet: Structure, variability, and evolution. Journal of Climate 14: 4191 4215. Carvalho LMV, Jones C, Ambrizzi T. 2005. Opposite phases of the Antarctic Oscillation and Relationships with Intraseasonal to Interannual activity in the tropics during the Austral Summer. Journal of Climate 18: 702 718. Chang CP, Zhang YS, Li T. 2000. Interannual and interdecadal variations of the East Asian summer monsoon and tropical Pacific SSTs. Part I: Roles of the subtropical ridges. Journal of Climate 13: 4310 4325. Ding YH, Chan JCL. 2005. The East Asian summer monsoon: an overview. Meteorology and Atmospheric Physics 89: 117 142. Fan K. 2006. Atmospheric circulation in southern Hemisphere and summer rainfall over Yangtze River valley. Chinese Journal of Geophysics 49(3): 672 679. Fan K, Wang HJ. 2004. Antarctic oscillation and the dust weather frequency in North China. Geophysical Research Letters 31: L10201, DOI:10.1029/2004GL019465. Gao H, Xue F, Wang HJ. 2003. Influence of interannual variability of Antarctic oscillation on Mei-yu along the Yangtze and Huaihe River

1284 J. SUN ET AL. valley and its importance to prediction. Chinese Science Bulletin 48: 61 67. Garcia SR, Kayano MT. 2008. Climatological aspects of Hadley, Walker and monsoon circulations in two phases of the Pacific Decadal Oscillation. Theoretical and Applied Climatology 91: 117 127, DOI:10.1007/s00704-007-0301-9. Gong DY, Ho CH. 2003. Arctic Oscillation signals in East Asian summer monsoon. Journal of Geophysical Research 108(D2): 4066, DOI:10.1029/2002JD002193. Gong DY, Wang SW. 1999. Definition of antarctic oscillation index. Geophysical Research Letters 26: 459 462. Hall A, Visbeck M. 2002. Synchronous variability in the Southern Hemisphere atmosphere, sea ice, and ocean resulting from the annular mode. Journal of Climate 15: 3043 3057. Hendon HH, Thompson DWJ, Wheeler MC. 2007. Australian rainfall and surface temperature variations associated with the Southern Hemisphere annular mode. Journal of Climate 20(11): 2452 2467. Huang RH, Li WJ. 1988. Influence of heat source anomaly over the western tropical Pacific on the subtropical high over East Asia and its physical mechanism. Acta Meteorologica Sinica: 107 116 (in Chinese). Kalnay E, Kanamitsu M, Kistler R, Collins W, Deaven D, Gandin L, Iredell M, Saha S, White G, Woollen J, Zhu Y, Chelliah M, Ebisuzaki W, Higgins W, Janowiak J, Mo KC, Ropelewski C, Wang J, Leetmaa A, Reynolds R, Jenne R, Joseph D. 1996. The NCEP/NCAR 40-year reanalysis project. Bulletin of the American Meteorological Society 77: 437 471. Kidson JW. 1988. Interannual variations in the Southern Hemisphere circulation. Journal of Climate 1: 1177 1198. Kistler R, Kalnay E, Collins W, Saha S, White G, Woollen J, Chelliah M, Ebisuzaki W, Kanamitsu M, Kousky V, Dool H, Jenne R, Fiorino M. 2001. The NCEP-NCAR 50-year reanalysis: Monthly means CD-ROM and documentation. Bulletin of the American Meteorological Society 82(2): 247 268. Kwok R, Comiso JC. 2002. Spatial patterns of variability in Antarctic surface temperature: Connections to the Southern Hemisphere Annular Mode and the Southern Oscillation. Geophysical Research Letters 29(14): 29, 1705, DOI:10.1029/2002GL015415. Lau KM, Kim KM, Yang S. 2000. Dynamical and boundary forcing characteristics of regional components of the Asian summer monsoon. Journal of Climate 13: 2461 2482. L Heureux ML, Thompson DWJ. 2006. Observed relationships between the El Niño-Southern Oscillation and the extratropical zonal-mean circulation. Journal of Climate 19: 276 287. Li SL, Bates G. 2007. Influence of the Atlantic Multidecadal Oscillation on the winter climate of East China. Advance in Atmospheric Science 24(1): 126 135. Li CY, Mu MQ. 2001. The dipole in the Equatorial Indian Ocean and its impacts on climate. Chinese Journal of Atmospheric Sciences 24(4): 433 443, (in Chinese). Liang XZ, Wang WC. 1998. Association between China monsoon rainfall and tropospheric jets. Quarterly Journal of the Royal Meteorological Society 124: 2597 2623. Liu JP, Yuan X, Rind D, Martinson DG. 2002. Mechanism study of the ENSO and southern high latitude climate teleconnections. Geophysical Research Letters 29: 1679, DOI:10.1029/2002GL015143. Liu YM, Wu GX, Ren RC. 2004. Relationship between the subtropical anticyclone and Diabatic heating. Journal of Climate 17: 682 698. Reason CJC, Rouault M. 2005. Links between the Antarctic Oscillation and winter rainfall over western South Africa. Geophysical Research Letters 32: L07705, DOI:10.1029/2005GL022419. Rind D, Chandler M, Lerner J, Martinson DG, Yuan X. 2001. The climate response to basin-specific changes in latitudinal temperature gradients and the implications for sea ice variability. Journal of Geophysical Research 106: 20161 20173. Rodwell MR, Hoskins BJ. 2001. Subtropical anticyclones and summer monsoons. Journal of Climate 14: 3192 3211. Rogers JC, van Loon H. 1982. Spatial variability of sea level pressure and 500 mb height anomalies over the Southern Hemisphere. Monthly Weather Review 110: 1375 1392. Silvestri GE, Vera CS. 2003. Antarctic Oscillation signal on precipitation anomalies over southeastern South America. Geophysical Research Letters 30(21): 2115, DOI:10.1029/2003GL018277. Smith CA, Sardeshmukh P. 2000. The effect of ENSO on the intraseasonal variance of surface temperature in winter. International Journal of Climatology 20: 1543 1557. Tao SY, Chen LX. 1987. A review of recent research on the East Asian summer monsoon. China, Monsoon Meteorology. Oxford University Press: London. Thompson DWJ, Wallace JM. 2000. Annular modes in the extratropical circulation, Part I: Month-to-month variability. Journal of Climate 13: 1000 1016. Ueda H, Kawamura R. 2004. Summertime anomalous warming over the midlatitude Western North Pacific and its relationship to the modulation of the Asian monsoon. International Journal of Climatology 24: 1109 1120. Wang HJ, Fan K. 2005. Central-north China precipitation as reconstructed from the Qing dynasty: Signal of the Antarctic atmospheric oscillation. Geophysical Research Letters 32: L24705, DOI:10.1029/2005GL024562. Wang HJ, Matsuno T, Kurihara Y. 2000a. Ensemble hindcast experiments for the flood period over China in 1998 by use of the CCSR/NIES atmospheric general circulation model. Journal of the Meteorological Society of Japan 78(4): 357 365. Wang B, Wu R, Fu X. 2000b. Pacific-East Asian teleconnection: How does ENSO affect East Asian Climate? Journal of Climate 13: 1517 1536. Wang HJ, Xue F, Zhou GQ. 2002. The spring monsoon in South China and its relationship to large-scale circulation features. Advance in Atmospheric Science 19: 651 664. Webster PJ. 1972. Response of the tropical atmosphere to local, steady forcing. Monthly Weather Review 100: 518 541. Webster PJ, Yang S. 1992. Monsoon and ENSO: selectively interactive systems. Quarterly Journal of the Royal Meteorological Society 118: 877 926. Xie P, Arkin PA. 1997. Global precipitation: A 17-year monthly analysis based on gauge observations, satellite estimates, and numerical model outputs. Bulletin of the American Meteorological Society 78: 2539 2558. Xue F, Wang HJ, He JH. 2004. Interannual variability of Mascarene high and Australian high and their influences on East Asian summer monsoon. Journal of the Meteorological Society of Japan 82(4): 1173 1186. Yuan XJ. 2004. ENSO-related impacts on. Antarctic sea ice: a synthesis of phenomenon and mechanisms. Antarctic Science 16: 415 425. Zhang QY, Tao SY. 1998. Influence of Asian mid-high latitude circulation on East Asian summer rainfall. Acta Meteorologica Sinica 56: 199 211, (in Chinese).