Abstract

Modulation of the Pacific decadal oscillation (PDO) on the behavior of the East Asian summer monsoon (EASM) in El Niño decaying years has been studied. When El Niño is in phase with the PDO (El Niño/high PDO), the low-level atmospheric anomalies are characterized by an anticyclone around the Philippines and a cyclone around Japan, inducing an anomalous tripolar rainfall pattern in China. In this case, the western Pacific subtropical high (WPSH) experiences a one-time slightly northward shift in July and then stays stationary from July to August. The corresponding anomalous tripolar rainfall pattern has weak subseasonal variations. When El Niño is out of phase with the PDO (El Niño/low PDO), however, the anomalous Philippines anticyclone has a much larger spatial domain, thereby causing an anomalous dipole rainfall pattern. Accordingly, WPSH experiences clearly two northward shifts. Therefore, the related dipole rainfall pattern has large subseasonal variations. One pronounced feature is that the positive rainfall anomalies shift northward from southern China in June to central China in July and finally to northern China in August.

The different El Niño–EASM relationships are caused by the influences of PDO on the decaying speed of El Niño. During the high PDO phase, El Niño decays slowly and has a strong anchor in the north Indian Ocean warming, which is responsible for the anomalous EASM. Comparatively, during the low PDO phase, El Niño decays rapidly and La Niña develops in summer, which induces different EASM anomalies from that during the high PDO phase. Additionally, PDO changes El Niño behaviors mainly via modifying the background tropical winds.

1. Introduction

The East Asian summer monsoon (EASM) manifests a large year-to-year variation (Chang et al. 2011; Ding and Chan 2005; Huang et al. 2012; Lau 1992). The unusual EASM can bring serious hazards such as floods and droughts and cause unexpected economic losses. As a dominant air–sea coupled phenomenon in the tropical Pacific, El Niño–Southern Oscillation (ENSO) is suggested to be an important predictor for the EASM (e.g., Chang et al. 2000a,b; Wu et al. 2003; Zhang et al. 1996; Zhou and Chan 2007). The ENSO–EASM relationship is proved to be complicated and exhibits different features during different phases of ENSO (Chen 2002; Huang and Wu 1989; Wu et al. 2003). Huang and Wu (1989) proposed that, in the developing summer of El Niño, below-normal rainfall anomalies are observed in northern and southern China, while above-normal rainfall anomalies are seen in central China. In contrast, rainfall anomalies with an opposite polarity are observed in the decaying summer of El Niño. This result is confirmed by the case study of Lau and Weng (2001), who compared the Chinese summer rainfall anomalies between the developing and decaying stages of the 1997/98 El Niño. Meanwhile, several studies further reported that the rainfall variability in the vicinity of Yangtze River valley presents quasi-biennial signals, which is associated with the eastern Pacific sea surface temperature (SST) anomalies in the previous winter (Chen et al. 2000; Shen and Lau 1995; Weng et al. 1999).

The distinct ENSO–EASM relationship can be attributed to different atmospheric responses to ENSO between the developing and decaying phases. In the El Niño developing summers, warm SST anomalies could enhance convection in the tropical central Pacific and trigger an anomalous cyclone in the lower troposphere over the western North Pacific via the Gill–Matsuno mechanism (Gill 1980; Matsuno 1966), which in turn causes the anomalous EASM (Wu et al. 2003). In the El Niño decaying summers, however, the positive SST anomalies in the tropical central and eastern Pacific are weakened or even disappear (Wang et al. 2003; Wang and Wu 2012). Therefore, the influences from the central and eastern Pacific are negligible, while those from the Indian Ocean (Xie et al. 2009) or the in situ air–sea interaction over the western North Pacific (Wang et al. 2000) play a dominant role. Hence, an anomalous anticyclone around the Philippines is formed through the Indian Ocean capacitor mechanism (Xie et al. 2009) and the local air–sea interaction (Wang et al. 2000). This anomalous Philippines anticyclone connects El Niño to the following EASM. On the other hand, the East Asia–Pacific (EAP; Huang and Li 1987; Huang and Yan 1999) or namely the Pacific–Japan (PJ; Nitta 1987; Kosaka and Nakamura 2010) teleconnection is formed with opposite polarity in the El Niño developing and decaying summers, leading to roughly opposite EASM rainfall anomalies in East Asia (Huang et al. 2012).

The impacts of ENSO on climate are not stationary and can be modulated by the Pacific decadal oscillation (PDO; Mantua et al. 1997; Mantua and Hare 2002) (e.g., Chen et al. 2013; Gershunov and Barnett 1998; Power et al. 1999). This modulation effect, on one hand, is possibly caused by the interference of the PDO on the ENSO’s high-latitude teleconnections (Barlow et al. 2001; Gershunov and Barnett 1998; Yoon and Yeh 2010; Yu and Zwiers 2007; among others). For example, Yu et al. (2007) reported that the Pacific–North American (PNA)-like tropospheric circulation and climate anomalies in North America are enhanced only when ENSO and PDO are out of phase. In this case, anomalous atmospheric energy in both the North Pacific and the tropical Pacific regions transports in the same direction toward the North America, which facilitates the occurrence of the PNA-like stationary waves. Yoon and Yeh (2010) proposed that the PDO could modulate the lag impacts of El Niño on the extratropics-related summer rainfall over northeast Asia through changing the Eurasian-like pattern. On the other hand, the modulation effect of PDO could be also attributed to its influence on ENSO’s low-latitude teleconnections. L. Wang et al. (2008) contrasted the relationship of ENSO and the East Asian winter monsoon (EAWM) between the low and high PDO phases. They suggested that the anomalous Philippines anticyclone in the lower troposphere, which acts as a linkage between ENSO and the EAWM, is strongly enhanced (suppressed) when ENSO and the PDO are out of phase (in phase). Thus, the significant (insignificant) ENSO–EAWM relationship is established in the low (high) phase of the PDO. Chan and Zhou (2005) found that the early summer (May–June) monsoon rainfall anomalies in South China are significant when ENSO and the PDO are in phase but not so when they are out of phase. Such relationships are likely to result from different effects of ENSO and the PDO on the intensity of the western Pacific subtropical high.

Regarding the EASM, its relationship with ENSO also experiences changes in the interdecadal time scales (e.g., B. Wang et al. 2008b; Wu and Wang 2002; Xie et al. 2010), which may be related to the PDO phase transition. Zhu and Yang (2003) presented a general comparison of the EASM rainfall anomalies in the decaying phase of El Niño between high and low PDO phases. A modulation effect of the PDO on the El Niño–EASM relationship was found, but the related physical process is not clear yet. Yoon and Yeh (2010) attributed the different relationship between the northeast Asian (30°–50°N, 110°–145°E) summer rainfall and El Niño to the modulation of the PDO on the high-latitude atmospheric teleconnections. For the EASM, in contrast, the teleconnection in the low latitudes (i.e., the anomalous lower tropospheric anticyclone over the Philippines) is important to bridge El Niño and the following EASM (e.g., Wang et al. 2000; Zhang et al. 1999). Therefore, it is necessary to investigate how the PDO modulates the El Niño’s low-latitude teleconnections during its decaying stage and further changes the EASM. Moreover, it should be noticed that all the studies mentioned above focused on the summer mean [June–August (JJA)] rainfall. The EASM, however, has clear subseasonal variations that cannot be fully reflected in the seasonal mean (e.g., Ding and Chan 2005). Hence, it is also necessary to investigate the possible influences of the PDO on the ENSO–EASM relationship on the subseasonal time scales. These two issues will be addressed in this study. Here, we focus on the decaying years of El Niño because its influence on the EASM is stronger and more significant than that in the developing years of El Niño (e.g., Wu and Wang 2002).

The structure of this paper is arranged as follows: Section 2 describes the datasets and methods employed in this study. In section 3, we contrast the characteristics of the anomalous EASM in the decaying summer of El Niño between the high and low PDO phases. Section 4 delineates the evolution of the tropical SST anomalies in the decaying stage of El Niño during the high and low PDO phases. The mechanism responsible for the different El Niño–EASM relationships between the high and low PDO phases is discussed in this section. Finally, conclusions and a discussion are given in section 5.

2. Data and methods

The monthly mean data used in this study include the atmospheric variables, SST, outgoing longwave radiation (OLR), and Chinese station rainfall. The atmospheric variables, including wind, geopotential height, and sea level pressure (SLP), are from the National Centers of Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis dataset with a horizontal resolution of 2.5° × 2.5°, covering the period from January 1948 to the present (Kalnay et al. 1996). Interpolated OLR data from the National Oceanic and Atmospheric Administration (NOAA) with a 2.5° × 2.5° horizontal resolution are used as a proxy for the tropical convection (Liebmann and Smith 1996). This dataset starts from June 1974, but from March 1978 to December 1978 it is not available because all grid points are covered by missing values. Therefore, considering the continuity of the OLR data, the period from 1979 to 2010 is employed in this study. The Chinese rainfall data comprising 160 observational stations are derived from the Chinese Meteorological Data Center. This dataset starts from 1951 and updates every month to the present. The SST data are from the Hadley Centre Global Sea Ice and Sea Surface Temperature (HadISST) dataset, compiled by the Met Office Hadley Center (Rayner et al. 2003). It has 1° latitude by 1° longitude resolution and covers a long period since 1870. Given the consistency and reliability, the data period spanning from 1957 to 2010 is considered in this study. In addition, all data are detrended before analyses in order to remove the possible influence of the long-term trends.

The selection of El Niño cases is based on the winter [December–February (DJF)] mean Niño-3.4 index (Fig. 1), which is derived from the SST anomalies averaged over the central and eastern Pacific (5°S–5°N, 170°–120°W). An El Niño year is selected when the normalized Niño-3.4 index is greater than 0.5 (Table 1), following the method in Yoon and Yeh (2010). The PDO is identified with the empirical orthogonal function (EOF) analysis on the cool-season (November–April) mean SST anomalies in the North Pacific (north of 20°N). The first mode of the EOF is the PDO pattern and the corresponding principal component is defined as the PDO index (Fig. 1). This definition is consistent with that of Mantua et al. (1997). In this study, high (low) PDO phase denotes the years that the PDO index is greater (less) than zero. Following this definition, 24 (29) years are identified as the high (low) phase of the PDO. Thus, El Niño events are divided into two groups according to the different PDO phases (Table 1). El Niño during the high (low) PDO phase is denoted as El-hPDO (El-lPDO).

Fig. 1.

The normalized time series of winter (DJF) mean Niño-3.4 index (dashed line) and the PDO index (solid line).

Fig. 1.

The normalized time series of winter (DJF) mean Niño-3.4 index (dashed line) and the PDO index (solid line).

Table 1.

El Niño events classified based on the different PDO phases.

El Niño events classified based on the different PDO phases.
El Niño events classified based on the different PDO phases.

A popular EASM index proposed by Wang and Fan (1999) is employed in this study. It is based on the definition of difference in 850-hPa zonal winds between the regions (22.5°–32.5°N, 110°–140°E) and (5°–15°N, 90°–130°E) (indicated with boxes in Fig. 3). This index reflects the low-latitude shear vorticity in the tropical western Pacific. B. Wang et al. (2008a) contrasted 25 EASM indices and found that this index could best capture the dominant mode of the EASM that is defined by the multivariate EOF analysis based on a set of meteorological fields including precipitation and atmospheric circulation fields. In addition, this EASM index has a better correlation with the Niño-3.4 index than other indices (B. Wang et al. 2008a). Figure 2 depicts sliding correlations between winter mean Niño-3.4 index and the following EASM index with a 13-yr window. The correlations between Niño-3.4 index and the EASM index manifest clear interdecadal variations. Interestingly, this interdecadal change roughly synchronizes the transition of the PDO phases, especially around the mid-1970s. Significant strong positive correlations are observed from the late 1970s to late 1980s and from the early 1990s to early 2000s, corresponding to the high phase of the PDO (Fig. 2). This result implies that the PDO is likely to affect the ENSO–EASM relationship. Note that the strong positive correlation after the 1980s has an obvious interdecadal change around the 1990s, with more significant correlation before the 1990s and less significant correlation after the 1990s. This might be associated with the change in the spatial structure of El Niño. After the 1990s, El Niño Modoki with warming in the tropical central Pacific has a more frequent occurrence than the conventional El Niño with the warming in the tropical eastern Pacific (Ashok et al. 2007; Kug et al. 2009; among others). The atmospheric circulation responses to these two flavors of El Niño display distinct features, thereby possibly causing a different ENSO–EASM relationship (Feng et al. 2011; Weng et al. 2007).

Fig. 2.

Sliding correlations between winter (DJF) mean Niño-3.4 index and the EASM index of the following summer with a 13-yr window (dashed line) and the PDO index (solid line). The dotted reference lines indicate the 90% and 95% confidence levels, respectively.

Fig. 2.

Sliding correlations between winter (DJF) mean Niño-3.4 index and the EASM index of the following summer with a 13-yr window (dashed line) and the PDO index (solid line). The dotted reference lines indicate the 90% and 95% confidence levels, respectively.

3. Anomalous EASM in the decaying phase of El Niño during different PDO phases

a. Atmospheric circulation anomalies

The influence of El Niño on the following EASM is mainly through an anomalous Philippines anticyclone in the lower troposphere that persists from winter to the following summer (Wang et al. 2000; Zhang et al. 1999). This anomalous anticyclone has an important impact on the moisture transport from ocean to the East Asian continent, the movement of the western Pacific subtropical high (WPSH), and the accompanying Chinese rainfall pattern (Zhang et al. 1999). Hence, the behavior of this anomalous anticyclone could determine the different EASM anomalies in the decaying summer of El Niño during different PDO phases.

Figure 3 shows the composite summer [JJA (+1)] mean anomalous 850-hPa winds and related streamfunction in the decaying summer of El Niño during the low and high PDO phases, respectively. Here, +1 denotes the decaying year of El Niño. For the El-hPDO cases, the anomalous Philippines anticyclone is zonally elongated and located at about 20°N. The westerly and easterly wind anomalies on the southern and northern sides of this anticyclone indicate a strong horizontal shear vorticity in the tropical western Pacific. It facilitates the establishment of the Pacific–East Asian teleconnection (Wang et al. 2000) and a strong positive Niño-3.4–EASM index correlation (Fig. 2). On the other hand, there are ordinal cyclonic and anticyclonic anomalies on the north side of this Philippines anticyclone, resembling the EAP (Huang and Li 1987) or namely the PJ pattern (Nitta 1987). Compared with the El-hPDO years, this Philippines anticyclone exhibits completely different features in the El-lPDO years. It has a much larger spatial scale and covers the entire western North Pacific (Fig. 3b). Its northern edge extends 15° more northward (at about 45°N) compared to that (at about 30°N) in the El-hPDO years. Therefore, the shear vorticity in the subtropical western Pacific is weak because of the northward extension of the anomalous Philippines anticyclone, resulting in a weak Niño-3.4–EASM index relationship (Fig. 2) in the El-lPDO years. The strong anomalous Philippines anticyclone as shown in Fig. 3 suggests that this weak Niño-3.4–EASM index correlation as shown in Fig. 2 does not actually mean the weak impact of El Niño on the EASM. This result further reflects the complexity of the EASM as documented in other studies (Ding and Chan 2005; B. Wang et al. 2008a; among others). So far, there is no index that can fully describe the EASM features, although the EASM index we chose is broadly used and considered as a better one. Comparing Fig. 3a with Fig. 3b, one can find that the impacts of El Niño on the western North Pacific circulation are quite different in the following summer when the phase of PDO is considered.

Fig. 3.

Composite anomalies of the summer [JJA (+1)] mean wind (vector; m s−1) and streamfunction (shading; contour interval = 0.3 × 106 m2 s−1) at 850 hPa for (a) El-hPDO and (b) El-lPDO years. Wind vectors above 90% confidence level according to two-tailed Student’s t test are plotted. Here, +1 denotes the decaying years of El Niño. The two boxes indicate the regions for defining the EASM index.

Fig. 3.

Composite anomalies of the summer [JJA (+1)] mean wind (vector; m s−1) and streamfunction (shading; contour interval = 0.3 × 106 m2 s−1) at 850 hPa for (a) El-hPDO and (b) El-lPDO years. Wind vectors above 90% confidence level according to two-tailed Student’s t test are plotted. Here, +1 denotes the decaying years of El Niño. The two boxes indicate the regions for defining the EASM index.

The WPSH is another salient factor that could impact the EASM rainfall. The strength and location of the WPSH can affect the location where the droughts and floods tend to occur in East Asia (Chang et al. 2000a; Tao and Chen 1987). Figure 4 presents the composite maps of WPSH in the decaying summer of El Niño from June (+1) to August (+1) during different PDO phases. For the climatological mean state, the WPSH experiences progressive northward shifts twice during summer with the first shift in July and the second in August (Ding 2004; also see Fig. 4a). In the case of El-hPDO, however, the WPSH experiences slightly northward shift from the South China Sea to the East China Sea in July (+1) and then stays stationary from July (+1) to August (+1) (Fig. 4b). In the case of El-lPDO, in contrast, the WPSH experiences pronounced northward shift twice (Fig. 4c) just as the climatological mean conditions (Fig. 4a). One slight difference is that the WPSH shifts more northwestward in El-lPDO years than in the climatology.

Fig. 4.

The monthly location of the western Pacific subtropical high (indicated by 5870-gpm contour) for (a) climatology (averaged for the period 1957–2010), (b) El-hPDO, and (c) El-lPDO. The dotted, dashed, and solid lines indicate the conditions in June (+1), July (+1), and August (+1), respectively.

Fig. 4.

The monthly location of the western Pacific subtropical high (indicated by 5870-gpm contour) for (a) climatology (averaged for the period 1957–2010), (b) El-hPDO, and (c) El-lPDO. The dotted, dashed, and solid lines indicate the conditions in June (+1), July (+1), and August (+1), respectively.

Besides the differences in the meridional direction, the motion of the WPSH also experiences different features in the zonal direction between the El-hPDO and El-lPDO cases. Generally, the western boundary of the WPSH is located on the coast of East Asia at about 122°E for the climatological condition (Fig. 4a). In the El Niño decaying summer, the WPSH can extend farther westward and reach the East Asian inland regardless of the PDO phase (Figs. 4b,c). However, during the high PDO phases, the western boundary of the WPSH stretches westward to about 112°E but only to about 117°E during the low PDO phases. This difference is consistent with the difference in the anomalous Philippines anticyclone that stretches more westward for the El-hPDO case (Fig. 3).

b. Rainfall anomalies

Anomalous rainfall in East Asia may directly lead to destructive disasters to the inhabitants in this region. Hence, it is the key variable in the EASM studies. In this section, we will evaluate the Chinese rainfall anomalies in the decaying summer of El Niño when the phase of PDO is taken into account. Figure 5 depicts the summer mean and monthly rainfall anomalies in China for the El-hPDO and El-lPDO cases. In the summer mean sense, the rainfall anomalies in the El-hPDO years are featured with tripolar pattern: that is, robust less rainfall in the southern and northern parts of China and slightly more rainfall in the central part of China (Fig. 5a). In the El-lPDO years, in contrast, more rainfall is observed in almost the whole China especially in the northern part (Fig. 5e). The most prominent change in Figs. 5a,e is observed in northern China where the rainfall anomalies shift from negative to positive when the PDO phase changes from high to low. This result is consistent with Wu and Wang (2002) and Zhu and Yang (2003).

Fig. 5.

Composite rainfall anomalies (mm month−1) for (a) summer [JJA (+1)] mean, (b) June (+1), (c) July (+1), and (d) August (+1) in El-hPDO years. (e)–(h) As in (a)–(d), but for El-lPDO years. The dark red and blue circles indicate 90% confidence level based on two-tailed Student’s t test, respectively.

Fig. 5.

Composite rainfall anomalies (mm month−1) for (a) summer [JJA (+1)] mean, (b) June (+1), (c) July (+1), and (d) August (+1) in El-hPDO years. (e)–(h) As in (a)–(d), but for El-lPDO years. The dark red and blue circles indicate 90% confidence level based on two-tailed Student’s t test, respectively.

These different rainfall anomalies in El-hPDO and El-lPDO summers can be attributed to different water vapor transport (Fig. 7). In El-hPDO summers, the troposphere-integrated water vapor flux is featured with an anomalous anticyclone over the Philippines and an anomalous cyclone around Japan (Fig. 7a). Abundant moisture convergence is then induced along the climatological mean mei-yu–baiu band (at about 30°N) by both the southwesterly and northeasterly wind anomalies. Meanwhile, anomalous moisture divergence is formed in the subtropical (at about 20°N) and midlatitude (at about 40°N) East Asia. Such anomalous moisture distribution gives rise to the related rainfall anomalies in Fig. 5a. In the El-lPDO years, in contrast, the water vapor flux is characterized by one large anticyclonic anomaly over the western North Pacific (Fig. 7e). It leads to abundant water vapor convergence in the midlatitudes (at about 40°N) and favors above-normal rainfall anomalies in northern China (Fig. 5e).

The EASM rainfall has distinct seasonal march and subseasonal characteristics (e.g., Ding 2004). Therefore, it is meaningful to further investigate the monthly rainfall anomalies in addition to the above summer mean features. Hence, Figs. 5 and 6 show the spatial and temporal evolutions of the monthly rainfall anomalies in El Niño decaying years. In the case of El-hPDO, the rainfall anomalies in June (+1) show a tripolar pattern in the eastern China that resembles the JJA (+1) mean result, with above-normal rainfall in central China and below-normal rainfall straddling on the northern and southern sides (Figs. 5b, 6a). In July (+1), both the positive and negative rainfall anomaly belts move slightly northward (Figs. 5c, 6a). In August (+1), however, the rainfall anomaly belts stay stationary and show a similar distribution to that in July (+1). One difference between July (+1) and August (+1) is that the rainfall anomalies are enhanced in August (+1), especially in southern China (Figs. 5d, 6a). This evolution of the rainfall anomalies from June (+1) to August (+1) can be attributed to the subseasonal movement of the WPSH that advances slightly northward from June (+1) to July (+1) and stays stationary from July (+1) to August (+1) (Fig. 4b).

Fig. 6.

Temporal evolutions of composite zonal mean (east of 100°E in Chinese mainland) rainfall anomalies in El Niño decaying months during the (a) high PDO phase and (b) low PDO phase. The contour interval is 8 mm month−1. Shading denotes the positive anomalies.

Fig. 6.

Temporal evolutions of composite zonal mean (east of 100°E in Chinese mainland) rainfall anomalies in El Niño decaying months during the (a) high PDO phase and (b) low PDO phase. The contour interval is 8 mm month−1. Shading denotes the positive anomalies.

The monthly rainfall anomalies (Figs. 5b–d) can be accounted for by the corresponding water vapor transport (Figs. 7b–d). In June (+1), when the WPSH is located at about 20°N in El-hPDO cases, the anomalous moisture flux is featured with an anomalous cyclone to the south of Japan (Fig. 7b). It leads to abundant moisture convergence in central China and moisture divergence in northern China (Fig. 7b). This anomalous moisture distribution favors the anomalous rainfall pattern in Fig. 5b. Accompanied to the northward progressive of the WPSH in July (+1), an anticyclonic anomaly around the Philippines dominates the moisture flux whereas the cyclonic anomaly around Japan retreats northeastward (Fig. 7c). Hence, the related rainfall anomaly belts shift northward accordingly. In August (+1), the anomalous moisture flux pattern is roughly similar to that in July (+1) (Fig. 7d), consistent with the relative standing stage of the WPSH from July (+1) to August (+1) (Fig. 4d). Therefore, the anomalous rainfall pattern in August (+1) is roughly in agreement with that in July (+1).

Fig. 7.

As in Fig. 5, but for the vertically integrated water vapor flux (vectors; kg m−1 s−1) and moisture content (shading; kg m−2). Vectors above the 90% confidence level according to two-tailed Student’s t test are plotted.

Fig. 7.

As in Fig. 5, but for the vertically integrated water vapor flux (vectors; kg m−1 s−1) and moisture content (shading; kg m−2). Vectors above the 90% confidence level according to two-tailed Student’s t test are plotted.

In the case of El-lPDO, the evolution of rainfall anomalies shows distinct features among individual months (Figs. 5f–h, 6b). The rainfall anomalies in June (+1) take on a dipole pattern with negative signals in central China and positive signals in southern China (Fig. 5f). In July (+1), in contrast, the anomalous rainfall pattern is generally opposite to that in June (+1), displaying wet conditions in central China and dry conditions in southern China (Figs. 5g, 6b). This anomalous dipole rainfall pattern moves northward in August (+1), with above-normal rainfall anomalies in northern China and below-normal rainfall anomalies in central China (Fig. 5h). In this monthly rainfall anomaly evolving process, one pronounced feature is that the anomalous rainfall pattern displays a clear dipole pattern (Figs. 5f–h, 6b) that is different from the tripolar pattern in the El-hPDO years (Figs. 5b–d, 6a). Another pronounced feature in El-lPDO years is that the positive rainfall anomaly belt displays clear northward shifts from southern China in June (+1) to central China in July (+1) and finally to northern China in August (+1) (Figs. 5f–h, 6b). This is in synchronized with the climatological advances of the EASM rainfall band (Ding 2004). Hence, in the case of El-lPDO, the major rainfall band of the EASM is strongly enhanced in each summer months. The northward advance of the positive rainfall anomalies can be attributed to the northward shift of the WPSH. Accompanied with the two northward shifts of the WPSH in July (+1) and August (+1) in the El-lPDO years, the anomalous moisture convergence zone extends to higher latitudes from June (+1) to August (+1) (Figs. 7f–h) in the El-lPDO. Therefore, the northward shifts of the positive rainfall anomalies are formed.

Since the negative rainfall anomalies persisting in southern China in the El-hPDO years and the positive rainfall anomalies propagating from southern China to northern China in the El-lPDO years are strong and significant (Fig. 6), Fig. 8 evaluates the magnitude of these rainfall anomalies for the individual El Niño years during the different PDO phases. Almost all the El-lPDO cases are accompanied by positive rainfall anomalies in each summer month (Fig. 8b). However, in the El-hPDO cases (Fig. 8a), the consistency of negative rainfall anomalies is not as good as that in the El-lPDO cases. Nevertheless, 7 out of 8 El-hPDO years are accompanied with negative rainfall anomalies in August (Fig. 8a), indicating robust and significant rainfall anomalies.

Fig. 8.

The monthly rainfall anomalies averaging the stations in the boxes as shown in Fig. 5 for the individual El Niño years during the (a) high PDO phase and (b) low PDO phase.

Fig. 8.

The monthly rainfall anomalies averaging the stations in the boxes as shown in Fig. 5 for the individual El Niño years during the (a) high PDO phase and (b) low PDO phase.

4. Possible mechanism

a. Causes for the different seasonal mean EASM conditions

The different atmospheric circulations and rainfall anomalies in East Asia between El-hPDO and El-lPDO years could be attributed to the distinct evolutions of the SST anomalies in the tropical Pacific and Indian Ocean. Figure 9 displays the composite evolutions of the SST anomalies in the El-hPDO and El-lPDO years from winter to the following summer. In the El-hPDO years, warm SST anomalies over the tropical central and eastern Pacific peak in winter and decay in the following seasons, while warm SST anomalies in the Indian Ocean sustain strong intensity from winter to the following summer (Figs. 9a–c). In addition, there exist negative SST anomalies over the North Pacific and western Pacific during the decaying seasons, especially during the winter and following spring, consistent with the high PDO phase. In this case, the north Indian Ocean warming is believed to play a dominant role in the formation of the anomalous Philippines anticyclone in the decaying summer of El Niño by the Kelvin wave–Ekman divergence mechanism proposed by Xie et al. (2009) (Fig. 3a). Summertime north Indian Ocean warming could lead to an increase in the tropospheric temperature that triggers a Kelvin wave to propagate into the western Pacific. This Kelvin wave could in turn induce Ekman divergence in the western Pacific, thereby giving rise to the zonally elongated anomalous anticyclone around the Philippines (Xie et al. 2009). Meanwhile, the local air–sea interactions may also contribute to the formation of this anomalous anticyclone via the wind–evaporation mechanism (Wang et al. 2000).

Fig. 9.

Composite SST anomalies for the El Niño decaying years in (a) winter [D(0)JF(+1)], (b) spring [MAM(+1)], and (c) summer [JJA(+1)] during the high PDO phases. (d)–(f) As in (a)–(c), but for the low phase of the PDO. Contour interval is 0.2°C. Zero lines are omitted. The light, middle, and dark shading indicates the 90%, 95%, and 99% confidence levels based on two-tailed Student’s t test, respectively.

Fig. 9.

Composite SST anomalies for the El Niño decaying years in (a) winter [D(0)JF(+1)], (b) spring [MAM(+1)], and (c) summer [JJA(+1)] during the high PDO phases. (d)–(f) As in (a)–(c), but for the low phase of the PDO. Contour interval is 0.2°C. Zero lines are omitted. The light, middle, and dark shading indicates the 90%, 95%, and 99% confidence levels based on two-tailed Student’s t test, respectively.

In the El-lPDO winters, the positive SST anomalies in the tropical central and eastern Pacific are weaker and shift slightly more westward than those in the El-hPDO winters (Fig. 9d). Over the North Pacific and western Pacific, there are no negative SST anomalies, which is clearly different from that during the high PDO phase. Conversely, the weak positive SST anomalies are observed over the North Pacific, corresponding to the low phase of PDO. Importantly, the positive SST anomalies in the central and eastern Pacific decay in the subsequent spring (Fig. 9e) and rapidly change into negative in the following summer (Fig. 9f). These negative SST anomalies are intensified in the following seasons and reach a maximum in the following winter (figure not shown), indicating the formation of a La Niña event. Meanwhile, weak positive SST anomalies are observed in the Indian Ocean in winter and the following spring, but they disappear in the following summer. Therefore, the anomalous large-scale anticyclone around the Philippines (Fig. 3b) in the El-lPDO decaying summer is possibly induced by a developing La Niña pattern. The cooling forcing in the central Pacific (Fig. 9f) could excite a Rossby wave on the northwestern side of this cooling region (Gill 1980; Matsuno 1966), forming an anomalous large-scale anticyclone around the Philippines (Figs. 3b). This result is consistent with that obtained by Chen et al. (2012), who suggested that a robust large-scale anomalous Philippines anticyclone tends to be seen when a La Niña develops after a quickly decaying El Niño with 1000-yr outputs from a coupled general circulation model.

To further confirm the different atmospheric responses over East Asia arising from the different decaying features of El Niño, Fig. 10 gives composite seasonal evolutions of SLP anomalies in the decaying periods of El Niño during different PDO phases. An anomalous anticyclone in the western Pacific persists from winter to the following summer for the El-hPDO cases (Figs. 10a–c), which can be explained by the strong El Niño–type SST forcing (Wang et al. 2000) in winter and the following spring (Figs. 9a,b). In the following summer when El Niño decays, north Indian Ocean warming is responsible for the persistence of this anticyclone (Xie et al. 2009). Comparatively, for the El-lPDO cases, this anomalous anticyclone is observed in winter (Fig. 10d). However, it is much weaker than that for the El-hPDO cases. This result is caused by the weaker El Niño–type SST forcing (Fig. 9d). This anticyclone is quickly weakened in the following spring (Fig. 10e). Interestingly, it is reinvigorated in the subsequent summer (Fig. 10f). This behavior of the anomalous anticyclone can be explained by the rapidly decaying step of El Niño in MAM (+1) and the fast development of La Niña in JJA (+1) (Figs. 9e,f). This result further confirms that the large-scale summertime Philippines anticyclone during the low PDO phase is induced by the La Niña SST pattern.

Fig. 10.

As in Fig. 9, but for the sea level pressure. Contour interval is 0.3 hPa. Zero lines are omitted. The shading indicates 90% confidence level based on two-tailed Student’s t test.

Fig. 10.

As in Fig. 9, but for the sea level pressure. Contour interval is 0.3 hPa. Zero lines are omitted. The shading indicates 90% confidence level based on two-tailed Student’s t test.

b. Causes for the different subseasonal EASM conditions

1) Monthly evolutions of SST anomalies

To elaborate the evolutions of the SST anomalies in more details, we examine the monthly composite of zonal wind anomalies averaged in the tropical Pacific (10°S–10°N, 140°E–120°W) and SST anomalies averaged in the tropical central and eastern Pacific (5°S–5°N, 180°–90°W) and north Indian Ocean (0°–20°N, 40°–100°E) for the El-hPDO and El-lPDO events. It is known that the zonal wind anomalies in the tropical Pacific could affect the evolving SST anomalies in the tropical central and eastern Pacific. In the case of El-hPDO, the anomalous westerly winds in the tropical Pacific sustain a long period until early summer in July (+1) (Fig. 11a). With these anomalous westerly winds, the trade winds at the equator are weakened, facilitating the persistence of El Niño in the decaying stage through weakening the oceanic upwelling by Ekman transport. Therefore, warm SST anomalies in the tropical central and eastern Pacific can persist until August (+1) (Fig. 11b). On the contrary, in the case of El-lPDO, the anomalous westerly winds are weakened quickly and changed into anomalous easterly winds since April (+1) (Fig. 11a). In this case, the trade winds at the equator are strengthened after April (+1); thus, El Niño decays quickly due to the enhanced oceanic upwelling by Ekman transport (Fig. 11b). The positive SST anomalies become negative from April (+1) to May (+1) in the central and eastern Pacific. Moreover, the negative SST anomalies are intensified in the subsequent months (Fig. 11b) and mature in the following winter (figure not shown), indicating the formation of a La Niña event.

Fig. 11.

The evolutions of (a) 850-hPa zonal wind anomalies averaged over the tropical Pacific region (10°S–10°N, 140°E–120°W), (b) SST anomalies averaged over the central and eastern Pacific (5°S–5°N, 180°–90°W), and (c) SST anomalies averaged over the north Indian Ocean (0°–20°N, 40°–100°E) in El Niño decaying years during the high (dashed lines) and the low (solid lines) PDO phases.

Fig. 11.

The evolutions of (a) 850-hPa zonal wind anomalies averaged over the tropical Pacific region (10°S–10°N, 140°E–120°W), (b) SST anomalies averaged over the central and eastern Pacific (5°S–5°N, 180°–90°W), and (c) SST anomalies averaged over the north Indian Ocean (0°–20°N, 40°–100°E) in El Niño decaying years during the high (dashed lines) and the low (solid lines) PDO phases.

Figure 11c shows the monthly evolution of the SST anomalies in the north Indian Ocean. It reveals that the positive SST anomalies in the north Indian Ocean display a similar intensity from January (+1) to April (+1) during both the high and the low PDO phases. This situation is changed in May (+1) when the decaying steps of El Niño diverge. When El Niño decays in the early spring during the low PDO phases, the warming SST anomalies in the north Indian Ocean weaken fast (Fig. 11c). In contrast, when the El Niño decaying step is postponed into the early summer during the high PDO phases, the north Indian Ocean warming tends to reach a peak in June (+1) and persist into the whole summer (Fig. 11c). Hence, the different decaying steps of El Niño affect evolutions of the north Indian Ocean warming. This result is supported by Du et al. (2009), Li et al. (2012), and Xie et al. (2010). They suggested that a slow (fast)-decaying El Niño has a strong (weak) anchor in the north Indian Ocean warming in JJA (+1) through the internal air–sea interaction over the tropical Indo-Pacific Ocean.

2) Western Pacific convective activities

In section 3, the circulation and rainfall anomalies over East Asia were shown to exhibit clear subseasonal characteristics besides the seasonal mean features. The different monthly evolutions of the SST anomalies over the Indo-Pacific Ocean in Fig. 11 can induce different convective activities over the tropical western Pacific, which is suggested to account for the subseasonal characteristics of the EASM (Huang and Sun 1992). When the convective activities are enhanced over the tropical western Pacific, the WPSH usually experiences pronounced northward shifts from June to August. When the summer convective activities are suppressed over the tropical western Pacific, in contrast, the progressive northward shift of the WPSH is not obvious (Huang and Sun 1992).

Figure 12 shows the composite OLR anomalies in the El-hPDO and El-lPDO years, which is a good measure of the convection in the tropical region. Due to the availability of the OLR data, the composite is based on El Niño events after 1979. For the El-hPDO events, the positive OLR anomalies are observed from January (+1) to August (+1) (Fig. 12a), indicating suppressed convection over the western Pacific. According to the studies of Huang and Sun (1992), this suppressed convective activity in summer facilitates the slowly northward movement of the WPSH and is in favor of anchoring the WPSH south of 30°N (Fig. 4b). This western Pacific convective evolution is closely related to the evolutions of SST anomaly in the tropical Pacific and Indian Ocean. In the El-hPDO years, the tropical central and eastern Pacific has strong warm SST anomalies (Fig. 11b) from January (+1) to May (+1) that could induce anomalous Walker circulation with ascending anomalies over the central Pacific and descending anomalies over the western Pacific. Therefore, the convection is suppressed over the western Pacific from January (+1) to May (+1). From June (+1) to August (+1), the warm SST anomalies in the central and eastern Pacific are weakened (Fig. 11b), so the anomalous Walker circulation in the Pacific sector is weakened simultaneously. At this time, the north Indian Ocean has robust warm SST anomalies (Fig. 11c) and is responsible for the suppressed convection over the western Pacific through the Kelvin wave–Ekman divergence mechanism (Xie et al. 2009).

Fig. 12.

Composite OLR anomalies averaged over the western Pacific (110°–140°E) in the El Niño decaying months during (a) the high PDO phase and (b) the low PDO phase. Counter interval is 4.0 W m−2. Zero lines are omitted. The shading denotes 90% confidence level based on a two-tailed Student’s t test.

Fig. 12.

Composite OLR anomalies averaged over the western Pacific (110°–140°E) in the El Niño decaying months during (a) the high PDO phase and (b) the low PDO phase. Counter interval is 4.0 W m−2. Zero lines are omitted. The shading denotes 90% confidence level based on a two-tailed Student’s t test.

For the El-lPDO events, in contrast, the anomalous convective activities over the western Pacific are suppressed from January (+1) to May (+1) but enhanced from June (+1) to August (+1) (Fig. 12b). The suppressed convection from January (+1) to May (+1) is induced by the descending branch of the anomalous Walker circulation just as that in the El-hPDO years. On the contrary, from June (+1) to August (+1), the cold SST anomalies develop rapidly in the tropical central and eastern Pacific. It could cause opposite anomalous Walker circulation with the ascending branch in the tropical western Pacific and descending branch in the tropical eastern Pacific. Therefore, convection is enhanced in the western Pacific. This enhanced western Pacific convection then favors the pronounced northward shift of the WPSH (Fig. 4c) and the rainband (Figs. 5f–h). Note that the enhanced convection shifts northward from June (+1) to August (+1). It reaches about 10°N in July (+1) and 25°N in August (+1). This tends to cause two northward shifts of the WPSH (Fig. 4b).

c. Influences of PDO on the behavior of El Niño

On the basis of the aforementioned studies, an important question is raised that how the PDO phase impacts the evolution and intensity of El Niño. To illustrate this question, Fig. 13 gives the differences in the surface winds between the high and low PDO years. The notable differences are characterized by significant cyclonic anomalies in the extratropical Pacific and associated strong westerly wind anomalies in the tropical central Pacific. These anomalous equatorial westerlies result in weaker trade winds during the high PDO phase than low PDO phase. In addition, slightly weak easterly wind anomalies are observed in the tropical eastern Pacific. Thus, a convergence is formed in the tropical eastern Pacific. Hence, the PDO conveys its influence to the tropics through the atmospheric bridge as revealed by Barnett et al. (1999), Pierce et al. (2000), and Wang and An (2002).

Fig. 13.

Differences of the winter (November–March) mean surface wind between the high PDO years and low PDO years. The shading indicates wind above 90% confidence level.

Fig. 13.

Differences of the winter (November–March) mean surface wind between the high PDO years and low PDO years. The shading indicates wind above 90% confidence level.

The differences in the tropical Pacific background surface winds between the high and low PDO phases tend to induce the distinct equatorial westerly anomalies in El Niño years as shown in Fig. 14. The different background tropical winds between the high and low PDO phases lead to a displacement of the El Niño–related westerly anomalies (Fig. 14). During the high PDO phase, the westerly anomalies are displaced eastward about 15° relative to those during the low PDO phase. Wang and An (2002) attributed this eastward displacement of the westerly anomalies to the enhanced mean trade wind convergence in the tropical eastern Pacific, which tends to cause eastward shift of the anomalous atmospheric heating and then favors the El Niño–related westerly anomalies moving eastward. In this study, the enhanced wind convergence in the tropical eastern Pacific is also observed in Fig. 13, thereby favoring the eastward shift of the westerly anomalies during the high PDO phase, as shown in Fig. 14. Furthermore, the displacement of the equatorial westerly anomalies is a key factor that can change the intensity and period of El Niño through the delayed oscillator theory (Battisti and Hirst 1989; Suarez and Schopf 1988; Wang and An 2002). The eastward displacement of the equatorial westerly anomalies increases the distance in which the upwelling oceanic Rossby wave propagates to the western boundary. Thus, the reflected upwelling oceanic kelvin wave over the western boundary, which has a negative feedback to the warm SST anomalies, would be decayed. This decay favors the positive downwelling Kelvin wave–SST feedback and increases the warm SST anomalies during the high PDO phase. Importantly, this decay lengthens the turnabout period from El Niño to La Niña. Hence, El Niño is strong and decays slowly during the high PDO phase but is weak and decays fast during the low PDO phase (Fig. 9).

Fig. 14.

Zonal distributions of the composite winter (November–March) mean tropical surface zonal wind anomalies in El Niño years during the high PDO (dashed line) and low PDO (solid line) phases.

Fig. 14.

Zonal distributions of the composite winter (November–March) mean tropical surface zonal wind anomalies in El Niño years during the high PDO (dashed line) and low PDO (solid line) phases.

In addition, the equatorial westerly anomalies in the central–eastern Pacific associated with El Niño have different intensities during the different PDO phases (Fig. 14). They are strong during the high PDO phase possibly arising from the suppressed background trade winds. With these robust westerly anomalies, the equatorial upwelling is reduced and the amplitude of El Niño is then strengthened. Conversely, during the low PDO phase, the enhanced background trade winds tend to induce the reversed conditions with weaker tropical westerly anomalies and stronger equatorial upwelling, thereby weakening the amplitude of El Niño.

5. Conclusions and discussion

Based on NCEP–NCAR reanalysis dataset, NOAA OLR data, HadISST data, and observed rainfall data from 160 China stations, this study investigated the behavior of the EASM in El Niño decaying years during different PDO phases. It reveals that, when El Niño occurs during the high PDO phase, a zonally elongated anticyclonic anomaly in the lower troposphere is observed around the Philippines in the following summer. To the northern side of this anticyclone, there is an anomalous cyclone around Japan. The anomalous anticyclone and cyclone dominate the East Asian summer moisture transport. Accordingly, moisture convergence over central China is induced by the anomalous southwesterly and northeasterly winds. Meanwhile, moisture divergence over southern and northern China is formed. Hence, this moisture distribution gives rise to an anomalous tripolar rainfall pattern in China from June (+1) to August (+1), with more rainfall in central China and less rainfall in southern and northern China. Moreover, The WPSH has a weak northward movement. It only experiences a one-time slightly northward shift in July (+1) and then stays stationary from July (+1) to August (+1). Corresponding to this motion of the WPSH, the monthly rainfall anomalies display a similar tripolar pattern from month to month.

In contrast, when El Niño occurs during the low PDO phase, the anomalous Philippines anticyclone has a much larger domain and covers the entire western North Pacific. In this case, the strong anticyclone controls the anomalous water vapor transportation in East Asia, thereby causing enhanced and reduced moisture in its northern and southern sides. Hence, a dipole anomalous rainfall pattern is formed in China from June (+1) to August (+1), which is different from the tripolar rainfall pattern for the El-hPDO cases. In this case, the WPSH has a pronounced northward movement. It experiences two northward shifts that happen in July (+1) and August (+1). Accordingly, the rainfall anomalies have robust subseasonal variations from June to August. One pronounced variation is that the positive rainfall anomalies experience two northward advances from southern China in June (+1) to central China in July (+1) and finally to northern China in August (+1). Hence, the major rainfall band of the EASM is strongly enhanced in each summer months. This evolution of the positive rainfall anomalies is closely related to the northward shift of the enhanced moisture convergence that is transported by the anomalous Philippines anticyclone.

The above distinct EASM behaviors can be attributed to the different El Niño decaying features when PDO is in its different phases. During the high PDO phase, El Niño is strong and decays slowly. The associated SST anomalies in JJA (+1) are featured with weak warming in the tropical central and eastern Pacific and significant warming in the Indian Ocean. This Indian Ocean warming is responsible for the formation of the anomalous Philippines anticyclone in the low latitudes based on the Indian Ocean capacitor mechanism. In this case, the western Pacific convection is suppressed from preceding winter to the following summer. The suppressed convection in summer then facilitates the weak movement of the WPSH. Unlike the situations during the high PDO phases, El Niño is weak and decays rapidly during the low PDO phases. The tropical Pacific SST anomalies in JJA (+1) evolve into a developing La Niña pattern with cooling in the central and eastern Pacific and warming in the western Pacific. Hence, the developing La Niña induces a stronger and larger anomalous anticyclone around the Philippines via the Gill–Matsuno mechanism (Gill 1980; Matsuno 1966). In this case, the western Pacific convection is suppressed from preceding winter to the following spring, but is enhanced in the following summer. The enhanced convection then facilitates two shifts of the WPSH in July (+1) and August (+1), leading to different subseasonal variations of the EASM from that during the low PDO phase. In addition, the spatial pattern of El Niño during the different PDO phases exhibits slightly different features. The warm SST anomalies in the tropical Pacific are located more eastward during the high PDO phase than the low PDO phase. Therefore, the anomalous warm SST pattern during the high PDO phase is similar to conventional El Niño with the warming in the eastern Pacific, whereas during the low PDO phase it is similar to the El Niño Modoki with warming in the central Pacific (Ashok et al. 2007; Kug et al. 2009; among others). These two types of El Niño have different impacts on the EASM, as revealed by many documents (Feng et al. 2011; Weng et al. 2007; among others). These different spatial patterns of El Niño possibly cause the distinct EASM conditions.

The evolving feature of ENSO is substantially influenced by the extratropical ocean conditions (Kleeman et al. 1999; Wang and An 2001; Wang and An 2002). One way is possibly through the ocean process. Kleeman et al. (1999) reported that the midlatitude decadal changes of the heat transport in the oceanic subtropical cell could convey its impact into the tropical ocean. Fedorov and Philander (2000) proposed the midlatitude SST anomalies could intrude into the equatorial region, influence the thermocline behavior, and thus induce tropical SST anomalies. The other way is suggested to be the atmosphere process (Barnett et al. 1999; Pierce et al. 2000; Wang and An 2001, 2002). Vimont et al. (2001, 2003a,b) proposed that the tripolar anomalous SST pattern in the extratropical region persisting from winter to the following summer could induce tropical wind anomalies and further lead to tropical SST anomalies. Wang and An (2001, 2002) identified that the North Pacific SST changes can give rise to the changes in El Niño properties via modifying the tropical wind behavior. In this study, the results show that the PDO conveys its impact to the tropics through changing the equatorial background surface winds. During the high (low) PDO phase, the mean trade winds in the tropical Pacific are enhanced (suppressed) and the wind convergence moves to the east (west). The changes in the tropical background winds could induce strong (weak) and eastward (westward) displacement of the El Niño–related equatorial westerly anomalies, which increases (decreases) the SST anomalies and prolongs (shortens) the El Niño lifetime via the delayed oscillator theory.

Acknowledgments

The valuable comments and suggestions from three reviewers and editor have led to a significant improvement of this paper. This study is supported jointly by the National Natural Science Foundation of China (41025017, 41230527, and 41205047) and the Chinese Academy of Sciences (KZCX2-EW-QN204). This work is also supported by the Jiangsu Collaborative Innovation Center for Climate Change.

REFERENCES

REFERENCES
Ashok
,
K.
,
S. K.
Behera
,
S. A.
Rao
,
H.
Weng
, and
T.
Yamagata
,
2007
:
El Niño Modoki and its possible teleconnection
.
J. Geophys. Res.
,
112
,
C11007
,
doi:10.1029/2006JC003798
.
Barlow
,
M.
,
S.
Nigam
, and
E.
Berbery
,
2001
:
ENSO, Pacific decadal variability, and U.S. summertime precipitation, drought, and stream flow
.
J. Climate
,
14
,
2105
2128
,
doi:10.1175/1520-0442(2001)014<2105:EPDVAU>2.0.CO;2
.
Barnett
,
T. P.
,
D. W.
Pierce
,
M.
Latif
, and
D.
Dommenget
,
1999
:
Interdecadal interactions between the tropics and midlatitudes in the Pacific basin
.
Geophys. Res. Lett.
,
26
,
615
618
,
doi:10.1029/1999GL900042
.
Battisti
,
S. D.
, and
A. C.
Hirst
,
1989
:
Interannual variability in the tropical atmosphere–ocean system: Influence of the basic state and ocean geometry
.
J. Atmos. Sci.
,
46
,
1678
1712
,
doi:10.1175/1520-0469(1989)046<1687:IVIATA>2.0.CO;2
.
Chan
,
J. C. L.
, and
W.
Zhou
,
2005
:
PDO, ENSO and the early summer monsoon rainfall over south China
.
Geophys. Res. Lett.
,
32
, L08810, doi:10.1029/2004GL022015.
Chang
,
C.
,
Y.
Zhang
, and
T.
Li
,
2000a
:
Interannual and interdecadal variations of the East Asian summer monsoon and tropical Pacific SSTs. Part I: Roles of the subtropical ridge
.
J. Climate
,
13
,
4310
4325
,
doi:10.1175/1520-0442(2000)013<4310:IAIVOT>2.0.CO;2
.
Chang
,
C.
,
Y.
Zhang
, and
T.
Li
,
2000b
:
Interannual and interdecadal variations of the East Asian summer monsoon and tropical Pacific SSTs. Part II: The meridional structure of the monsoon
.
J. Climate
,
13
,
4326
4340
,
doi:10.1175/1520-0442(2000)013<4326:IAIVOT>2.0.CO;2
.
Chang
,
C.
,
Y.
Ding
,
N.-C.
Lau
,
R. H.
Johnson
,
B.
Wang
, and
T.
Yasunari
,
2011
: The Global Monsoon System: Research and Forecast. 2nd ed. World Scientific Series on Asia-Pacific Weather and Climate, Vol. 5, World Scientific, 590 pp.
Chen
,
W.
,
2002
:
Impacts of El Niño and La Niña on the cycle of the East Asian winter and summer monsoon
.
Chin. J. Atmos. Sci.
,
26
,
595
610
.
Chen
,
W.
,
H. F.
Graf
, and
R. H.
Huang
,
2000
:
The interannual variability of East Asian winter monsoon and its relation to the summer monsoon
.
Adv. Atmos. Sci.
,
17
,
48
60
.
Chen
,
W.
,
J. K.
Park
,
B.
Dong
,
R.
Lu
, and
W. S.
Jung
,
2012
:
The relationship between El Niño and the western North Pacific summer climate in a coupled GCM: Role of the transition of El Niño decaying phases
.
J. Geophys. Res.
,
117
, D12111, doi:10.1029/2011JD017385.
Chen
,
W.
,
J.
Feng
, and
R. G.
Wu
,
2013
:
Roles of ENSO and PDO in the link of the East Asian winter monsoon to the following summer monsoon
.
J. Climate
,
26
,
622
635
,
doi:10.1175/JCLI-D-12-00021.1
.
Ding
,
Y.
,
2004
: Seasonal march of the East-Asian summer monsoon. East Asian Monsoon, C.-P. Chang, Ed., World Scientific, 3–53.
Ding
,
Y.
, and
J. C. L.
Chan
,
2005
:
The East Asian summer monsoon: An overview
.
Meteor. Atmos. Phys.
,
89
,
117
142
,
doi:10.1007/s00703-005-0125-z
.
Du
,
Y.
,
S.-P.
Xie
,
G.
Huang
, and
K. M.
Hu
,
2009
:
Role of air–sea interaction in the long persistence of El Niño–induced north Indian Ocean warming
.
J. Climate
,
22
,
2023
2038
,
doi:10.1175/2008JCLI2590.1
.
Fedorov
,
A. V.
, and
S. G.
Philander
,
2000
:
Is El Niño changing?
Science
,
288
,
1997
2002
,
doi:10.1126/science.288.5473.1997
.
Feng
,
J.
,
W.
Chen
,
C.-Y.
Tam
, and
W.
Zhou
,
2011
:
Different impacts of El Niño and El Niño Modoki on China rainfall in the decaying phases
.
Int. J. Climatol.
,
31
,
2091
2101
,
doi:10.1002/joc.2217
.
Gershunov
,
A.
, and
T. P.
Barnett
,
1998
:
Interdecadal modulation of ENSO teleconnections
.
Bull. Amer. Meteor. Soc.
,
79
,
2715
2725
,
doi:10.1175/1520-0477(1998)079<2715:IMOET>2.0.CO;2
.
Gill
,
A. E.
,
1980
:
Some simple solutions for heat-induced tropical circulation
.
Quart. J. Roy. Meteor. Soc.
,
106
,
447
462
,
doi:10.1002/qj.49710644905
.
Huang
,
G.
, and
Z.
Yan
,
1999
:
The East Asian summer monsoon circulation anomaly index and its interannual variations
.
Chin. Sci. Bull.
,
44
,
1325
1329
,
doi:10.1007/BF02885855
.
Huang
,
R.
, and
W. J.
Li
,
1987
: Influence of the heat source anomaly over the western tropical Pacific on the subtropical high over East Asia. Proc. Int. Conf. on the General Circulation of East Asia, Chengdu, China, Chinese Academy of Sciences Institute of Atmospheric Physics,
40
51
.
Huang
,
R.
, and
Y. F.
Wu
,
1989
:
The influence of ENSO on the summer climate change in China and its mechanism
.
Adv. Atmos. Sci.
,
6
,
21
32
.
Huang
,
R.
, and
F.
Sun
,
1992
:
Impacts of the tropical western Pacific on the East Asian summer monsoon
.
J. Meteor. Soc. Japan
,
70
,
243
256
.
Huang
,
R.
,
J.
Chen
,
L.
Wang
, and
Z.
Lin
,
2012
:
Characteristics, processes, and causes of the spatio-temporal variabilities of the East Asian monsoon system
.
Adv. Atmos. Sci.
,
29
,
910
942
,
doi:10.1007/s00376-012-2015-x
.
Kalnay
,
E.
, and
Coauthors
,
1996
:
The NCEP/NCAR 40-Year Reanalysis Project
.
Bull. Amer. Meteor. Soc.
,
77
,
437
471
,
doi:10.1175/1520-0477(1996)077<0437:TNYRP>2.0.CO;2
.
Kleeman
,
R.
,
J. P.
McCreary
Jr.
, and
B. A.
Klinger
,
1999
:
A mechanism for generating ENSO decadal variability
.
Geophys. Res. Lett.
,
26
,
1743
1746
,
doi:10.1029/1999GL900352
.
Kosaka
,
Y.
, and
H.
Nakamura
,
2010
:
Mechanisms of meridional teleconnection observed between a summer monsoon system and a subtropical anticyclone. Part I: The Pacific–Japan pattern
.
J. Climate
,
23
,
5085
5108
,
doi:10.1175/2010JCLI3413.1
.
Kug
,
J.-S.
,
F.-F.
Jin
, and
S.-I.
An
,
2009
:
Two types of El Niño Events: Cold tongue El Niño and warm pool El Niño
.
J. Climate
,
22
,
1499
1515
,
doi:10.1175/2008JCLI2624.1
.
Lau
,
K.
,
1992
:
East Asian summer monsoon rainfall variability and climate teleconnection
.
J. Meteor. Soc. Japan
,
70
,
211
242
.
Lau
,
K.
, and
H.
Weng
,
2001
:
Coherent modes of global SST and summer rainfall over China: An assessment of the regional impacts of the 1997–98 El Niño
.
J. Climate
,
14
,
1294
1308
,
doi:10.1175/1520-0442(2001)014<1294:CMOGSA>2.0.CO;2
.
Li
,
Q.
,
R.-C.
Ren
,
M.
Cai
, and
G. X.
Wu
,
2012
:
Attribution of the summer warming since 1970s in Indian Ocean basin to the inter-decadal change in the seasonal timing of El Niño decay phase
.
Geophys. Res. Lett.
,
39
, L12702, doi:10.1029/2012GL052150.
Liebmann
,
B.
, and
C. A.
Smith
,
1996
:
Description of a complete (interpolated) outgoing longwave radiation dataset
.
Bull. Amer. Meteor. Soc.
,
77
,
1275
1277
.
Mantua
,
N. J.
, and
S. R.
Hare
,
2002
:
The Pacific decadal oscillation
.
J. Oceanogr.
,
58
,
35
44
,
doi:10.1023/A:1015820616384
.
Mantua
,
N. J.
,
S. R.
Hare
,
Y.
Zhang
,
J. M.
Wallace
, and
R. C.
Francis
,
1997
:
A Pacific interdecadal climate oscillation with impacts on salmon production
.
Bull. Amer. Meteor. Soc.
,
78
,
1069
1079
,
doi:10.1175/1520-0477(1997)078<1069:APICOW>2.0.CO;2
.
Matsuno
,
T.
,
1966
:
Quasi-geostrophic motions in the equatorial area
.
J. Meteor. Soc. Japan
,
44
,
25
43
.
Nitta
,
T.
,
1987
:
Convective activities in the tropical western Pacific and their impact on the Northern Hemisphere summer circulation
.
J. Meteor. Soc. Japan
,
65
,
373
390
.
Pierce
,
D. W.
,
T. P.
Barnett
, and
M.
Latif
,
2000
:
Connections between the Pacific Ocean tropics and midlatitudes on decadal time scales
.
J. Climate
,
13
,
1173
1194
,
doi:10.1175/1520-0442(2000)013<1173:CBTPOT>2.0.CO;2
.
Power
,
S.
,
T.
Casey
,
C.
Folland
,
A.
Colman
, and
V.
Mehta
,
1999
:
Inter-decadal modulation of the impact of ENSO on Australia
.
Climate Dyn.
,
15
,
319
324
,
doi:10.1007/s003820050284
.
Rayner
,
N. A
.,
D. E.
Parker
,
E. B.
Horton
,
C. K.
Folland
,
L. V.
Alexander
,
D. P.
Rowell
,
E. C.
Kent
, and
A.
Kaplan
,
2003
:
Global analyses of sea surface temperature, sea ice, and night marine air temperature since the late nineteenth century
.
J. Geophys. Res.
,
108
,
4407
,
doi:10.1029/2002JD002670
.
Shen
,
S.
, and
K. M.
Lau
,
1995
:
Biennial oscillation associated with the East Asian summer monsoon and tropical sea surface temperatures
.
J. Meteor. Soc. Japan
,
73
,
105
124
.
Suarez
,
M. J.
, and
P. S.
Schopf
,
1988
:
A delayed oscillator for ENSO
.
J. Atmos. Sci.
,
45
,
3283
3287
,
doi:10.1175/1520-0469(1988)045<3283:ADAOFE>2.0.CO;2
.
Tao
,
S. Y.
, and
L. X.
Chen
,
1987
: A review of recent research of the East Asian summer monsoon in China. Monsoon Meteorology, C.-P. Chang and T. N. Krishnamurti, Eds., Oxford University Press, 60–92.
Vimont
,
D. J.
,
D. S.
Battisti
, and
A. C.
Hirst
,
2001
:
Footprinting: A seasonal connection between the tropics and mid-latitudes
.
Geophys. Res. Lett.
,
28
,
3923
3926
,
doi:10.1029/2001GL013435
.
Vimont
,
D. J.
,
D. S.
Battisti
, and
A. C.
Hirst
,
2003a
:
The seasonal footprinting mechanism in the CSIRO general circulation models
.
J. Climate
,
16
,
2653
2667
,
doi:10.1175/1520-0442(2003)016<2653:TSFMIT>2.0.CO;2
.
Vimont
,
D. J.
,
J. M.
Wallace
, and
D. S.
Battisti
,
2003b
:
The seasonal footprinting mechanism in the Pacific: Implications for ENSO
.
J. Climate
,
16
,
2668
2675
,
doi:10.1175/1520-0442(2003)016<2668:TSFMIT>2.0.CO;2
.
Wang
,
B.
, and
Z.
Fan
,
1999
:
Choice of South Asian summer monsoon indices
.
Bull. Amer. Meteor. Soc.
,
80
,
629
638
,
doi:10.1175/1520-0477(1999)080<0629:COSASM>2.0.CO;2
.
Wang
,
B.
, and
S. I.
An
,
2001
:
Why the properties of El Niño changed during the late 1970s
.
Geophys. Res. Lett.
,
28
,
3709
3712
,
doi:10.1029/2001GL012862
.
Wang
,
B.
, and
S. I.
An
,
2002
:
A mechanism for decadal changes of ENSO behavior: Roles of background wind changes
.
Climate Dyn.
,
18
,
475
486
,
doi:10.1007/s00382-001-0189-5
.
Wang
,
B.
,
R.
Wu
, and
X.
Fu
,
2000
:
Pacific–East Asian teleconnection: How does ENSO affect East Asian climate?
J. Climate
,
13
,
1517
1536
,
doi:10.1175/1520-0442(2000)013<1517:PEATHD>2.0.CO;2
.
Wang
,
B.
,
R.
Wu
, and
T.
Li
,
2003
:
Atmosphere–warm ocean interaction and its impacts on Asian–Australian monsoon variation
.
J. Climate
,
16
,
1195
1211
,
doi:10.1175/1520-0442(2003)16<1195:AOIAII>2.0.CO;2
.
Wang
,
B
.,
Z.
Wu
,
J.
Li
,
J.
Liu
,
C.-P.
Chang
,
Y.
Ding
, and
G.
Wu
,
2008a
:
How to measure the strength of the East Asian summer monsoon
.
J. Climate
,
21
,
4449
4463
,
doi:10.1175/2008JCLI2183.1
.
Wang
,
B
,
J.
Yang
,
T.
Zhou
, and
B.
Wang
,
2008b
:
Interdecadal changes in the major modes of Asian–Australian monsoon variability: Strengthening relationship with ENSO since the late 1970s
.
J. Climate
,
21
,
1771
1789
,
doi:10.1175/2007JCLI1981.1
.
Wang
,
L.
, and
R.
Wu
,
2012
:
In-phase transition from the winter monsoon to the summer monsoon over East Asia: Role of the Indian Ocean
.
J. Geophys. Res.
,
117
,
D11112
,
doi:10.1029/2012JD017509
.
Wang
,
L.
,
W.
Chen
, and
R.
Huang
,
2008
:
Interdecadal modulation of PDO on the impact of ENSO on the East Asian winter monsoon
.
Geophys. Res. Lett.
,
35
,
L20702
,
doi:10.1029/2008GL035287
.
Weng
,
H.
,
K. M.
Lau
, and
Y.
Xue
,
1999
:
Multi-scale summer rainfall variability over China and its long-term link to global sea surface temperature variability
.
J. Meteor. Soc. Japan
,
77
,
845
857
.
Weng
,
H.
,
K.
Ashok
,
S. K.
Behera
,
S. A.
Rao
, and
T.
Yamagata
,
2007
:
Impacts of recent El Niño Modoki on dry/wet conditions in the Pacific Rim during boreal summer
.
Climate Dyn.
,
29
,
113
129
,
doi:10.1007/s00382-007-0234-0
.
Wu
,
R.
, and
B.
Wang
,
2002
:
A Contrast of the East Asian summer monsoon–ENSO relationship between 1962–77 and 1978–93
.
J. Climate
,
15
,
3266
3279
,
doi:10.1175/1520-0442(2002)015<3266:ACOTEA>2.0.CO;2
.
Wu
,
R.
,
Z. Z.
Hu
, and
B. P.
Kirtman
,
2003
:
Evolution of ENSO-related rainfall anomalies in East Asia
.
J. Climate
,
16
,
3742
3758
,
doi:10.1175/1520-0442(2003)016<3742:EOERAI>2.0.CO;2
.
Xie
,
S. P.
,
K.
Hu
,
J.
Hafner
,
H.
Tokinaga
,
Y.
Du
,
G.
Huang
, and
T.
Sampe
,
2009
:
Indian Ocean capacitor effect on Indo–western Pacific climate during the summer following El Niño
.
J. Climate
,
22
,
730
747
,
doi:10.1175/2008JCLI2544.1
.
Xie
,
S. P.
,
Y.
Du
,
G.
Huang
,
X. T.
Zheng
,
H.
Tokinaga
,
K.
Hu
, and
Q.
Liu
,
2010
:
Decadal shift in El Niño influences on Indo–western Pacific and East Asian climate in the 1970s
.
J. Climate
,
23
,
3352
3368
,
doi:10.1175/2010JCLI3429.1
.
Yoon
,
J.
, and
S. W.
Yeh
,
2010
:
Influence of the Pacific decadal oscillation on the relationship between El Niño and the northeast Asian summer monsoon
.
J. Climate
,
23
,
4525
4537
,
doi:10.1175/2010JCLI3352.1
.
Yu
,
B.
, and
F.
Zwiers
,
2007
:
The impact of combined ENSO and PDO on the PNA climate: A 1,000-year climate modeling study
.
Climate Dyn.
,
29
,
837
851
,
doi:10.1007/s00382-007-0267-4
.
Yu
,
B.
,
A.
Shabbar
, and
F.
Zwiers
,
2007
:
The enhanced PNA-like climate response to Pacific interannual and decadal variability
.
J. Climate
,
20
,
5285
5300
,
doi:10.1175/2007JCLI1480.1
.
Zhang
,
R.
,
A.
Sumi
, and
M.
Kimoto
,
1996
:
Impact of El Niño on the East Asian monsoon: A diagnostic study of the ’86/87 and ’91/92 events
.
J. Meteor. Soc. Japan
,
74
,
49
62
.
Zhang
,
R.
,
A.
Sumi
, and
M.
Kimoto
,
1999
:
A diagnostic study of the impact of El Niño on the precipitation in China
.
Adv. Atmos. Sci.
,
16
,
229
241
,
doi:10.1007/BF02973084
.
Zhou
,
W.
, and
J. C. L.
Chan
,
2007
:
ENSO and the South China Sea summer monsoon onset
.
Int. J. Climatol.
,
27
,
157
167
,
doi:10.1002/joc.1380
.
Zhu
,
Y. M.
, and
X. Q.
Yang
,
2003
:
Relationship between Pacific decadal oscillation and climate variabilities in China
.
Acta Meteor. Sin.
,
61
,
641
654
.