Interannual Relationships Between the Tropical Sea Surface Temperature and Summertime Subtropical Anticyclone Over the Western N

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 116, D13111, doi:10.1029/2010JD015554, 2011

Interannual relationships between the tropical sea surface temperature and summertime subtropical anticyclone over the western North Pacific Pei‐Hsuan Chung,1 Chung‐Hsiung Sui,2 and Tim Li3 Received 27 December 2010; revised 8 March 2011; accepted 6 April 2011; published 13 July 2011.

[1] The interannual variability of the Western North Pacific Subtropical High (WNPSH) in boreal summer is investigated with the use of the NCEP/NCAR Reanalysis Data. The most significant change of the 500 hPa geopotential height field appears at the western edge of the WNPSH, with dominant 2–3 year and 3–5 year power spectrum peaks. The 2–3 year oscillation of the WNPSH and associated circulation and sea surface temperature (SST) patterns possess a coherent eastward propagating feature, with a warm SST anomaly (SSTA) and anomalous ascending motion migrating from the tropical Indian Ocean in the preceding autumn to the maritime continent in the concurrent summer of a strong WNPSH. A strong WNPSH is characterized by anomalous anticyclonic circulation and maximum subsidence in the western North Pacific (WNP). The anomalous WNPSH circulation has an equivalent barotropic vertical structure and resides in the sinking branch of the local Hadley circulation, triggered by enhanced convection over the maritime continent. A heat budget analysis reveals that the WNPSH is maintained by radiative cooling. The 3–5 year oscillation of the WNPSH exhibits a quasi‐stationary feature, with a warm SSTA (anomalous ascending motion) located in the equatorial central eastern Pacific and Indian Ocean and a cold SSTA (anomalous descending motion) located in the western Pacific. The anomaly pattern persists from the preceding winter to the concurrent summer of a high WNPSH. The greatest descent is located to the southeast of the anomalous anticyclone center, where a baroclinic vertical structure is identified. The zonal phase difference and the baroclinic vertical structure suggest that the anomalous anticyclone on this timescale is a Rossby wave response to a negative latent heating associated with the persistent local cold SSTA. ECHAM4 model experiments further confirm that the 2–3 year mode is driven by the SSTA forcing over the maritime continent, while the 3–5 year mode is driven by the local SSTA in the WNP. Citation: Chung, P.‐H., C.‐H. Sui, and T. Li (2011), Interannual relationships between the tropical sea surface temperature and summertime subtropical anticyclone over the western North Pacific, J. Geophys. Res., 116, D13111, doi:10.1029/2010JD015554.

1. Introduction mer monsoon. The zonal shift of the western part of the WNPSH dominates the position of large‐scale frontal zones [2] The subtropical high over the North Pacific is a domi- in East Asia [Tao and Chen, 1987]. A southwestward nant component of atmospheric general circulation that extension of the WNPSH would result in more water vapor experiences a distinct seasonal cycle with the maximum convergence and excessive rainfall along the Yangtze River strength and northernmost position occurring in the boreal valley [Chang et al., 2000a, 2000b; Zhou and Yu, 2005]. The summer. The evolution of the subtropical high over the precipitation characters in Japan and Korea were also greatly western North Pacific (hereafter WNPSH) in summer is influenced by the movement of the WNPSH through chang- associated with the onset and withdrawal of the Asian sum- ing the transport of water vapor into the Baiu or Changma frontal area [e.g., Tomita et al., 2004; Lee et al., 2005]. 1Department of History and Geography, Taipei Municipal University of [3] Traditionally, subtropical highs are regarded as the Education, Taipei, Taiwan. descending branch of the Hadley circulation maintained by 2Department of Atmospheric Sciences, National Taiwan University, radiative cooling. This idealized, zonally symmetric Hadley Taipei, Taiwan. circulation model, however, would lead to a weaker 3International Pacific Research Center, University of Hawai’iatMānoa, Honolulu, Hawaii, USA. (stronger) subtropical high in the summer (winter) hemisphere, which is contrary to observations [Hoskins, 1996]. Copyright 2011 by the American Geophysical Union. Moreover, the observed summertime subtropical highs are 0148‐0227/11/2010JD015554

D13111 1of19 D13111 CHUNG ET AL.: TROPICAL SSTA AND SUBTROPICAL HIGH D13111 in the form of high‐pressure cells rather than a zonally [6] In addition to the warm pool heating, the El Nino‐ uniform high‐pressure belt. The dynamics of the summer- Southern Oscillation (ENSO) may also influence the East time subtropical high is better understood in the context of Asian climate through the modulation of low‐level winds planetary waves, as suggested by Ting [1994], who found over the WNP [e.g., Wang et al., 2000; Chang et al., 2000a, that summertime subtropical stationary wave patterns are 2000b; Lau and Nath,2000;Kawamura et al., 2001; Lau maintained by the global diabatic heating. and Wu, 2001; Chou, 2004]. An anomalous low‐level [4] Hoskins [1996] and Rodwell and Hoskins [2001] anticyclone dominates the WNP during the mature phase of applied the monsoon‐desert mechanism proposed by an El Nino, whereas an anomalous cyclone occurs during Rodwell and Hoskins [1996] to argue that the remote diabatic the La Nina mature winter. The anomalous anticyclone heating over the North American monsoon region could persists through the following spring and summer, leading induce a Rossby wave pattern to the west and contribute to to a weak WNP monsoon in the ENSO decaying summer the lower tropospheric subtropical anticyclone in the North [e.g., Wang et al., 2000, 2003; Chou et al., 2003]. The Pacific. Rodwell and Hoskins [2001] further examined the anticyclone in early summer favors the northward move- dynamics of the summertime subtropical high in a global ment of the Mei‐yu/Baiu front and thus enhances (weakens) nonlinear model and concluded that the combined effect of theEastAsia(WNP)summer monsoon rainfall [Chang the topography and the monsoonal heating is of primary et al., 2000a, 2000b]. importance for the generation of the surface subtropical [7] The wintertime anomalous low‐level anticyclone in highs and subsidence aloft over the eastern Pacific. This WNP may be regarded as a Rossby wave response to the differs from Chen et al. [2001], who argued that summertime suppressed convection over the western Pacific, which is subtropical highs can be maintained by downstream energy possibly caused both by the remote warm SSTA over the propagation of stationary Rossby waves forced by convec- eastern Pacific [Wang et al., 2000; Chou, 2004; Chen et al., tion heating over the Asian monsoon region in a linear quasi‐ 2007] and the local cold SSTA [Lau and Nath, 2000; Wang geostrophic model. Their experiments, however, could not et al., 2000]. The maintenance of the WNP anticyclone from produce the distinct meridional dependence of the observed the El Nino mature winter to its decaying early summer is vertical structures of the summertime planetary wave, pos- attributed to a season‐dependent thermodynamic air‐sea sibly due to the lack of other heating sources such as sensible feedback [Wang et al., 2000, 2003; Kawamura et al., 2001; heating over the western portion of the continent and radia- Chou, 2004; Li and Wang, 2005]. During the El Nino tion cooling off the west coast of the continent. Liu et al. decaying summer, the SSTA in the eastern Pacific has [2004] performed atmospheric GCM experiments to show diminished, yet the low‐level anticyclone still exists over the the importance of the global surface sensible heating and WNP. It is likely that both the cold SSTA in the WNP in late longwave radiative cooling in the formation of the subtrop- spring/early summer [Sui et al., 2007; Wu et al., 2010] and ical anticyclones. Miyasaka and Nakamura [2005] demon- the basin‐wide warming in the IO in summer [Yang et al., strated that the summertime subtropical high can be 2007; Xie et al., 2009; Wu et al., 2009a, 2009b] are accurately reproduced by the near‐surface thermal contrast responsible for maintaining the WNPSH during the summer. between a cooler ocean and a hotter land to the east. The role [8] The Asian summer monsoon shows a strong quasi‐ of local air‐sea interactions in generating the subtropical biennial variability [e.g., Shen and Lau, 1995; Barnett, anticyclone was investigated numerically by Seager et al. 1991; Chang and Li, 2000; Li et al., 2001; Li and Zhang, [2003]. All studies above suggest the important role of the 2002], which is termed as the tropospheric biennial oscil- land‐ocean‐atmosphere feedback in maintaining the sub- lation (TBO) in the Asian‐Australian monsoon region tropical high. [Meehl, 1994, 1997]. The tropical Pacific and southeast IO [5] In addition to the annual cycle, the subtropical high in (SEIO) SSTA associated with the Asian summer monsoon the boreal summer also exhibits remarkable interannual also exhibits a biennial characteristic [Chang et al., 2000a; variations. Convective heating over the western Pacific Wang et al., 2001]. A hybrid coupled GCM simulation warm pool is regarded as an important forcing that affects shows that the TBO may exist even in the absence of ENSO the extent of the subtropical high in the boreal summer. [Li et al., 2006]. On the other hand, ENSO may be linked to According to Nitta [1987], convective activity in the trop- the monsoon TBO through the following three branches of ical western Pacific during summer can influence the teleconnection: (1) WNP monsoon‐SEIO SSTA tele- Northern Hemispheric circulation; that is, Rossby waves connection or the Philippine‐Sumatra pattern [Li et al., generated by the anomalous tropical heat source can prop- 2002; Li et al., 2006; Wu et al., 2009b]; (2) IO convec- agate into higher latitudes. This leads to a north–south tion‐Central Eastern Pacific (CEP) SSTA teleconnection dipole between subtropics and middle latitude East Asia, [Meehl, 1987; Chang and Li, 2000; Yu et al., 2002; known as the Pacific‐Japan (PJ) Pattern. Lu [2001] found Watanabe and Jin, 2002; Li et al., 2003]; and (3) Pacific‐ that stronger convection over the warm pool is associated East Asia teleconnection [Wang et al., 2000, 2003]. with more eastward and northward migration of the sub- [9] Because the WNPSH is regarded as one of the tropical high. Lu and Dong [2001] performed numerical important components of the Asian summer monsoon sys- model experiments and showed that suppressed convection tem, the objective of this study is to conduct an analysis of caused by a cold SSTA in the warm pool results in anom- the interannual evolution of the WNPSH and investigate alous anticyclonic circulation in the lower troposphere over possible mechanisms that link to the variability. The rest the subtropical western North Pacific (WNP), providing a of the paper is organized as follows. Section 2 describes the favorable condition for the westward extent of the sub- data used in this study. In section 3 a WNPSH index is tropical high. defined and two distinctive spectrum peaks on quasi‐biennial

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(2–3 year) and lower‐frequency (3–5 year) periods are [13] According to Yanai et al. [1973], the large‐scale heat identified. A regression analysis is subsequently performed and moisture budget residuals, Q1 and Q2, respectively, are to obtain the large‐scale circulation, SST, outgoing long- defined and computed by wave radiation, and heat and moisture budgets associated k with the two oscillation periods. By synthesizing all features, p @ @ Q1 cp þ ~ r þ ! ð1Þ possible mechanisms responsible for the two oscillations of p0 @t @p the WNPSH are proposed. The idealized experiments of the ECHAM4 AGCM are also examined in section 4. In section and 5 the seasonal lead‐lag patterns of regressed SSTA and cir- – – @q @q culation fields associated with the 2 3 year and 3 5 year Q2 L þ ~ rq þ ! ð2Þ oscillations of the WNPSH are discussed. Section 6 com- @t @p prises a summary and discussions. where is the potential temperature, q is the water vapor mixing ratio, ~ is the horizontal velocity, w is the vertical p velocity. L is the latent heat of vaporization, k = R/c with R 2. Data and Model Description p the gas constant and cp the specific heat capacity at constant 2.1. Reanalysis Data pressure of dry air, and p0 = 1000 hPa. 14 ~ ‐ [10] The primary data set used in this study is the National [ ] In equations (1) and (2), , , and q denote large scale Centers for Environmental Prediction–National Center for fields that are mathematically defined as Reynolds averaged Atmospheric Research (NCEP/NCAR) Reanalysis [Kalnay quantities (normally with an overbar, being neglected here) et al., 1996] that includes wind, geopotential height, and are customarily assumed resolvable by reanalysis data. ‐ stream function, vertical p velocity, sea level pressure, So Q1 and Q2 are computed by 6 hourly data at each grid ‐ velocity potential, and outgoing longwave radiation flux on point from ERA products. The 6 hourly estimates of Q1 – and Q2 are averaged into monthly mean values and multi- a 2.5 × 2.5 grid for 48 years (1958 2005). This is the lon- −1 gest reanalysis product available. Another set of data used in plied by cp to be expressed in terms of equivalent warming/ the study is the NOAA Extend Reconstructed Sea Surface cooling rate (K/day). Temperature version 2 (ERSST v2) developed on a 2° × 2° [15] The Q1 and Q2 can be interpreted by: – grid for the 48 years (1958 2005). The extended recon- @s′!′ struction of historical SST was using improved statistical Q ¼ Q þ LðÞrc e s′v′ ð3Þ 1 R @p methods from 1854 to the present. A full description of the ERSST v2 is available in the work of Smith and Reynolds [2004]. and [11] Further, this study used the European Centre for @q′!′ ‐ Q ¼ LðÞþc e Lrq′v′ þ L ð4Þ Medium Range Weather Forecasts (ECMWF) Reanalysis 2 @p (ERA), which also contains the above variables at the same resolution from 1957 to 2002. A full description of the ERA where QR is radiative heating rate, c and e are condensa- is available in the work of Gibson et al. [1996, 1997]. The tion and evaporation per unit mass of air respectively, s is analysis results from this study, based on NCEP/NCAR the dry static energy, and the prime denotes deviation from data, were compared against those from ERA and the major the Reynolds average. The primed variables denote subgrid features were found to be essentially the same. Most of the scale fluctuations that are assumed to be well separated from results discussed in this study are based on NCEP/NCAR the resolvable large‐scale variable. Equation (3) represents reanalysis, except the following. the total effect of radiative heating, latent heat released by [12] In order to discuss the dynamic features of the cli- net condensation L(c − e), and the horizontal and vertical con- mate oscillations in this study, the heat and moisture bud- ‐ ‐ vergence of sensible heat fluxes due to subgrid scale eddies gets were computed, the so called apparent heat source (Q1) such as cumulus convection and turbulence. Equation (4), and apparent moisture sink (Q2) [e.g., Yanai et al., 1973], on the other hand, represents the total effect of net con- using ERA. This decision is made due to the fact that the densation and divergence of eddy moisture fluxes due to overall simulation of the diabatic heating (and therefore subgrid‐scale eddies. The contributions of the eddy hori- the moisture) field over the monsoon region in ERA is quite zontal transport terms r · s′v′ in equation (3) and r · q′v′ in close to reality [Annamalai et al., 1999]. In addition, a equation (4) are generally small and may be negligible com- comparison of the heat and moisture budgets based on ‐ ‐ pared to the horizontal transports by the large scale motion ERA, NCEP/NCAR R 1 Reanalysis, and NCEP/DOE [Yanai and Johnson, 1993]. AMIP‐II (R‐2) Reanalysis [Kanamitsu et al., 2002] was completed. The results show that budgets computed based on both ERA and R‐2 reanalysis data are quite similar to the 2.2. Atmospheric General Circulation Model observed rainfall variability in the tropical and East Asian [16] The Atmospheric General Circulation Model (AGCM) monsoon regions. However, the R‐2 reanalysis data only used in the study is ECHAM version 4.6 (hereafter ECHAM4), cover the satellite period from 1979 to present. Therefore which was developed by the Max Planck Institute for Mete- the Q1 and Q2 budgets were computed at each vertical level orology [Roeckner et al., 1996]. The model was run at a (23 pressure levels) using the 6‐hourly ERA data from 1958 horizontal resolution of spectral triangular 42 (T42), with 2.8° to 2001. grid and 19 vertical levels in a hybrid sigma pressure coor- dinate system. The climatological monthly SST constructed

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Figure 1. The standard deviation of 2–7yearband‐pass filtered summer‐mean geopotential height at 500 hPa (Z500) normalized by the corresponding zonal mean Z500 (shaded) and 5870‐m contour line for a strong (red), normal (green), and weak (blue) WNPSH composite. by phase II of the Atmospheric Model Intercomparison WNPSH. For comparison, two additional indices are defined Project (AMIP II) was used as the model low boundary based on 850 hPa stream function (y850) and geopotential forcing. The ECHAM4 model is one of the Top‐3AMIPII height anomalies (Z850) in JJA averaged over the region of models in simulating the interannual variability modes of the (120–150°E, 15–25°N), where the strongest interannual Asian‐Australian monsoon [Zhou et al., 2009a]. It shows variation of y850 and Z850 occurs (hereafter termed as y850 excellent performance in simulating the El Nino‐related tro- and Z850 indices). The y850 and Z850 fields were used by Lee pospheric temperature anomalies [Zhou and Zhang,2011]. et al. [2005] and Lu [2001] to study the subtropical high Thus the model is used in the study for investigating the variability over the WNP. interannual variations of the summertime circulation in WNP. [18] The power spectrums of the Z500, y850, and Z850 indices are examined. The Z500 index shows two dominant periods at 2–3 years and 3–5 years, both of which pass the 3. Characteristics Associated With y the Anomalous Subtropical High 95% confidence level (Figure 2a). The 850 and Z850 index also presents two similar peaks although with the 3–5 year 3.1. WNPSH Index band (2–3 year band) being above (slightly below or at) the [17] In this study the interannual variation of the summer 95% confidence level (figures not shown). Additional power ‐ subtropical high over the WNP is measured using the 500 hPa spectrum analyses with 44 year ERA40 Z500 data (figure not shown) show a similar result. Thus in the following, the Z geopotential height field (Z500). Although the summer 500 Pacific subtropical high can be detected in the lower tropo- index will be used to represent the WNPSH quasi‐biennial sphere, the WNPSH is most pronounced in the middle tro- (2–3 year) and lower‐frequency (3–5 year) variability. posphere. The center of the subtropical high in the lower [19] Through a Fourier series decomposition of the troposphere is located in the eastern Pacific, while the center WNPSH index series, two band‐pass filtered time series of the subtropical high in the middle troposphere resides over were constructed by summing Fourier components within the western North Pacific [Liu et al., 2001; Liu and Wu, the 2–3 year and 3–5 year periods (Figure 2b). Hereafter, the 2004]. For this reason, the variations of the WNPSH in two time series are referred to as the 2–3 year and 3–5 year WNPSH indices. Positive indices represent the strong and this study are measured by the Z500 variability, following many previous studies [e.g., Tao and Chen, 1987; Zhou and westward extending phase of the subtropical high, while negative indices represent the opposite. In the following Li, 2002]. Figure 1 shows the standard deviation (~Z) of the analysis, the circulation and SST fields are regressed against summer (JJA) mean Z500 during the period of 1958–2005. It is calculated from 2 to 7 year band‐pass filtered data and has the two WNPSH indices to investigate spatial and temporal been normalized by its zonal mean value to illustrate the structures associated with the 2–3 year and 3–5 year oscil- zonally asymmetric character. Figure 1 shows that the largest lations of the WNPSH. variability occurs in the northwest Pacific region, southwest – of the climatological mean subtropical high as depicted by 3.2. Circulation Patterns Associated With the 2 3 Year and 3–5 Year Oscillations the contour line of Z500 = 5870 m. A WNPSH index is defined as the area mean value of Z500 anomalies in JJA in [20] The structures related to 2–3 year and 3–5 year the region of (120°–140°E, 10°–30°N). This index is used to oscillations of the WNPSH were investigated by examining measure the strength of the interannual variability of the the horizontal patterns of regressed fields of summer mean

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tion and discussion of physical mechanisms afterward will be merely based on the features of the positive phase. [21] The regressed Z500, V500, y850, and y200 are shown in Figure 3. The WNPSH high anomaly of the 2–3 year oscillation is associated with a meridional standing wave pattern with a zonally elongated positive Z500 anomaly and anticyclonic circulation in the WNP, a negative Z500 anomaly over the northern East Asia, and a positive anomaly over Eurasia (Figure 3b). The horizontal distribution of y850 in Figure 3c shows a pair of anticyclonic circulations symmetric about 10 N. One is located in the WNP (overlapping with a positive Z500 anomaly center) and the other is located south of the Equator. The circulation pattern of y200 (Figure 3a) shows that an anticyclonic circulation extends from southern to southeastern China and the WNP. [22] In the 3–5 year oscillation, an anticyclone and a positive Z500 anomaly center (Figure 3e) are located near Taiwan when the WNPSH extends westward. The positive Z500 anomaly extends further to the tropical central eastern Pacific than that in 2–3 year oscillation. The positive anomaly in Z500 and y850 fields (Figures 3e and 3f) acts as the source of a Rossby wave train propagating from the WNP to North America. The distribution of y200 (Figure 3d) exhibits a cyclonic circulation pattern over China. In the WNP, y200 is in general out of phase with y850. [23] Figure 4 shows the regressed SST patterns for the 2– 3 year and 3–5 year modes. The SST field has been normalized at each grid point based on the climatological standard deviation to capture the obvious SSTA variations. The SST distribution associated with the westward extending WNPSH in the 2–3 year oscillation (Figure 4a) features a warm SSTA extending from the northern IO southeastward to South China Sea, maritime continent and northeastern Australia, and a cold SSTA in the central eastern equatorial and northeastern and southeastern Pacific. For the 3–5 year oscillation (Figure 4b), however, a warm SSTA appears in the tropical IO and the central eastern equatorial Pacific. The SSTA in the central eastern Pacific is surrounded by cold SST anomalies to its north and south and in the maritime continent. [24] In the 2–3 year oscillation, the large‐scale divergent circulation in upper and lower levels are shown by regressed c200 and c850 fields in Figures 5a and 5b. A broad overturning circulation occurs in the summer of a strong WNPSH over the western Pacific, with the low‐level con- Figure 2. (a) Power spectrums of the summer‐mean geo- vergent (divergent) flow toward Java and the eastern IO (out potential height at 500 hPa (WNPSH index) and (b) decom- of the tropical central eastern Pacific region) (Figure 5b). posed band‐pass filtered time series of the (top) 2–3years The regressed OLR field in Figure 5c shows enhanced and (bottom) 3–5 years WNPSH index. The pink curve in convection over the maritime continent and suppressed Figure 2a denotes the 95% confidence level. convection over the Philippine Sea and the equatorial eastern Pacific, consistent with the divergent wind in Figures 5a and 5b. A zonal band of enhanced convection is also observed to the north of the WNPSH. The anomalous OLR ERSST, outgoing longwave radiation (OLR), Z500, 500 hPa pattern in the western Pacific resembles a negative phase of vertical p velocity (w500), 200 hPa and 850 hPa stream y y the PJ pattern [Nitta, 1987]. function ( 200 and 850), 200 hPa and 850 hPa velocity – c c [25] For the 3 5 year oscillation, the divergent circulation potential ( 200 and 850), and wind fields at 850 hPa and pattern differs from that of the 2–3 year oscillation. The 500 hPa (V850 and V500). Since the derived results of regres- regressed c200 and c850 fields (shown in Figures 5d and 5e) sion can be applied to both the positive phase (represented by reveal two anomalous Walker cells in the tropics, with two a strengthened and westward WNPSH) and the negative upper‐tropospheric divergence centers located in the tropical phase (represented by a weakened WNPSH), the interpreta- central‐eastern Pacific and IO, respectively, and a convergence center located in the region extending from southern

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Figure 3. Regressed summer‐mean fields of stream function at (a, d) 200 hPa and (c, f) 850 hPa and (b, e) geopotential height with significant wind fields at 500 hPa for the 2–3 year (Figures 3a, 3c, and 3e) and 3–5 year (Figures 3b, 3d, and 3f) oscillations of WNPSH. Shaded (vectors) areas pass the 90% confidence level.

Philippines to the maritime continent and eastern Australia low‐level southward divergent flow to the south of the (Figure 5d). The regressed OLR field (Figure 5f) shows the WNPSH may bring the moisture into the equatorial region regions of suppressed convection in the maritime continent and enhance convection in the maritime continent. The and subtropical North Pacific and the regions of enhanced enhanced convection further reinforces the WNPSH subsi- convection in the equatorial eastern Pacific and western dence via a positive feedback process. tropical IO, consistent with the features of the divergent [27] For the 3–5 year oscillation, the regressed vertical flows shown in the c200 and c850 fields (Figures 5d and 5e). overturning circulation along 110°E–150°E (Figure 6b) [26] The meridional overturning circulation of the 2– shows two descending branches at 0°–10°S and 15°–25°N, 3 year oscillation over the western Pacific is further revealed coinciding with the two suppressed convective centers in by the regressed zonally averaged vertical p velocity (w)at the OLR field. There is no direct linkage between the WNP each pressure level and latitude‐vertical divergent wind and equatorial regions (Figure 5f). It can be seen from c profile (v, w) along 110°E–140°E (Figure 6a). Consistent analysis in Figures 5d and 5e that the equatorial branch of with the regressed OLR and c fields, the flow is upward the descending motion within the western Pacific is a result over the maritime continent (0°–10°S) and East Asia near of the anomalous large‐scale Walker circulation. In the c Japan (30°–40°N), with the former being stronger than the field, however, the convergent flow over the WNP is not as latter. The southern branch of upward motion causes upper‐ strong as that in the maritime continent/eastern Australia tropospheric divergent flows toward the Philippine Sea (at (Figure 5d). The result suggests that unlike the 2–3 year 10°–20°N) where the downward motion is pronounced. The oscillation, the subsidence in the WNP is not directly

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Figure 4. Regressed summer‐mean SSTA fields in (a) the 2–3yearand(b)3–5 year oscillations of WNPSH, normalized by climatological standard deviations at each grid point. Shaded areas pass the 90% confidence level. attributed to the equatorial forcing in the 3–5 year oscilla- phase with the anticyclone center, the maximum subsidence tion. Rather, it may be forced through a Rossby wave (suppressed convection in Figure 5f) in the 3–5 year oscil- response to anomalous heating associated with the local lation is located to the southeast of the anomalous anticy- cold SSTA in Figure 4b [Sui et al., 2007; Wu et al., 2010]. clone center. Such a phase difference suggests that the anomalous WNPSH in the former is possibly a direct 3.3. Vertical Structures of the WNPSH Anomalies response to the local Hadley circulation anomaly, whereas in the latter it is a baroclinic Rossby wave response to a neg- [28] The analysis above indicates that the large‐scale circulations associated with an enhanced summertime ative heating anomaly induced by the local cold SSTA (Figure 4b). WNPSH in the 2–3 year and 3–5 year oscillations are dis- [30] To reveal the vertical structure of the anomalous tinctly different. The 2–3 year oscillation is accompanied by WNPSH, the domains of the strongest downward motion an anomalous local Hadley circulation that connects the and anomalous subtropical high center on the 2–3 year and maritime continent and the WNP, while the 3–5 year 3–5 year timescales respectively were chosen. The vertical oscillation is characterized by the anomalous Walker cir- profiles of the domain‐averaged regressed fields of vertical culation and a cold SSTA to the southeast of the anomalous p velocity, relative vorticity, Q and Q are computed at WNPSH. To reveal the specific processes that affect 1 2 each pressure level. WNPSH, the horizontal and vertical structures of various [31] Figure 8a shows the regressed vertical p velocity dynamic and thermodynamic variables associated with the two oscillations are further compared. profiles over the area of strongest downward motion in the 2–3 year and 3–5 year oscillations. The strongest subsidence [29] First, the relative zonal position between the anom- is located in the upper troposphere (300 hPa) in the 2–3 year alous anticyclone and descending motion centers during a – strong WNPSH phase are examined by depicting the re- oscillation but in the middle troposphere (400 500 hPa) in the 3–5 year oscillation. The magnitude of the subsidence in gressed y (contours) and w (blue shades) fields in 850 500 the 3–5 year oscillation is weaker than that in the 2–3 year Figure 7. Only the stronger (at the 95% significant level) oscillation. Figure 8b shows the vertical profiles of the portion of the regressed positive w field is shown in 500 regressed relative vorticity fields over the anomalous anti- Figure 7. While the subsidence center (suppressed convec- cyclone center associated with the 2–3 year and 3–5 year tion in Figure 5c) in the 2–3 year oscillation is zonally in

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Figure 5. Regressed summer‐mean velocity potential and divergent wind at (a, d) 200 hPa and (b, e) 850hPaand(c,f)upwardlongwaveradiationflux(similar to outgoing longwave radiation) for the 2–3 year (Figures 5a, 5c, and 5e) and 3–5 year (Figures 5b, 5d, and 5f) oscillations of WNPSH. The red (blue) contours denote positive (negative) velocity potential anomalies or enhanced (suppressed) convection. Shaded regions pass the 90% confidence level. oscillations. While it exhibits a clear equivalent barotropic in the establishment of the anomalous WNPSH on the 2– structure in the 2–3 year oscillation, the vorticity has a 3 year timescale. On the other hand, a peak of anomalous baroclinic structure in the 3–5 year oscillation, changing Q2 at 850 hPa in the 3–5 year oscillation implies strong from negative anomalies in the lower troposphere to a negative condensation heating at the top of the boundary positive anomaly in the upper troposphere (above 250 hPa). layer. The negative condensation heating is induced by This result is consistent with circulation patterns over the the local cold SSTA over the region of (150°–170°E, 15°– WNP area shown in Figures 3a, 3c, 3d, and 3f. 20°N) that is to the southeast of the anomalous anticyclone [32] The vertical distributions of regressed Q1 and Q2 and is responsible for a baroclinic vertical structure. anomalies averaged over the strongest descent region are [33] In both the oscillations the warm SSTA under the shown in Figures 8c and 8d. Note that the amplitude of the anomalous anticyclone may be regarded as a passive negative Q1 anomaly in the middle‐upper troposphere is response to the atmospheric forcing [Wang et al., 2005; Wu greater in the 2–3 year oscillation than in the 3–5 year and Kirtman, 2005; Zhou et al., 2008]. The anomalous oscillation (Figure 8c), while the amplitude of the negative Q2 anticyclone leads to the suppression of convection and thus anomaly in the lower troposphere is greater in the 3–5 year the increase of downward solar radiation into the ocean oscillation than in the 2–3 year oscillation (Figure 8d). This surface. Meanwhile, less evaporation due to the weakening implies that enhanced longwave radiative cooling due to wind speed may also increase SST. the reduction of the vertically integrated moisture caused by [34] The distinctive horizontal and vertical structures the descent‐induced dry advection plays an important role above suggest that the physical mechanisms responsible

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Figure 6. Regressed vertical velocity (contours) and divergent wind fields (a) averaged over 110°E–140°E for 2–3 year and (b) averaged over 110°–150°E for 3–5 year oscillations of WNPSH. Shaded regions pass the 90% confidence level. for the anomalous WNPSH may differ in the 2–3 year and emitted into space. The change of local vertical motion and 3–5 year oscillations. According to the Gill [1980] model, a circulation in the WNP is induced by the remote forcing via Rossby wave response is located to the west of a heating anomalous local Hadley circulation associated with enhanced source. This corresponds well to the 3–5 year oscillation convective heating in the maritime continent. case, in which there is a phase shift between the anomalous [35] The regressed results above are consistent with the anticyclone and the anomalous heating (or vertical motion). summertime subtropical LOSECOD diabatic heating pattern Further, the Rossby wave response to the anomalous heating identified by Wu and Liu [2003] and Wu et al. [2009], who is baroclinic. These features are accurately observed in the showed that the dominant diabatic heating sources over the 3–5 year oscillation. Therefore the diagnosis result in this northwestern Pacific region are convective heating and study suggests that the variation of the WNPSH on the 3– longwave radiation cooling. The corresponding diagnoses of 5 year timescale is primarily caused by an atmospheric Figures 5, 7, and 8 indeed demonstrate that in the 2–3 year Rossby wave response to a negative heating associated with oscillation, the longwave radiation cooling associated with a cold SSTA to the southeast of the anticyclone center. In the anomalous descent by the local Hadley circulation pre- contrast, in the 2–3 year oscillation, the anomalous WNPSH vails over the negative convective heating; while in the 3– center is zonally colocated with the vertical motion, and its 5 year mode, the negative convective heating due to the cold vertical structure is approximately barotropic. This implies SSTA prevails over the longwave radiation cooling. The that the Rossby wave response is unlikely and the down- negative convective heating will induce an atmospheric ward motion is a major driving mechanism. That is, the Rossby wave response to the west as the anomalous strengthening of the WNPSH on the 2–3 year timescale WNPSH. As a consequence, the subtropical anticyclone is primarily attributed to the radiative cooling due to the anomaly of the 2–3 year oscillation possesses a barotropic descent‐induced dry advection; the less clouds and water structure, whereas in the 3–5 year oscillation possesses a vapor in the atmosphere, the more longwave radiation is baroclinic structure as presented in Figure 8b. This is

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Figure 7. Regressed summer‐mean stream function at 850 hPa (contour) and positive p velocity (only most significant descending region is indicated here by shading) at 500 hPa for (a) the 2–3 year and (b) 3– 5 year oscillations of WNPSH. because a convective heating in the subtropics can generate the same simulation strategy was used as in the WNP run a Sverdrup balance stream function pattern with anticyclone except with a +1°C SSTA specified in the maritime conti- (cyclone) to the east and cyclone (anticyclone) to the west nent region (90°–150°E, 20°S–5°N) (red box in Figure 9b). of the heating (negative heating) in the lower troposphere The amplitude of this specified SSTA in the WNP and MC and the opposite in the upper troposphere, as discussed by regions is 2–3 times larger than the composite SSTA field Wu and Liu [2003], Liu et al. [2004], and Wu et al. [2009]. associated with the observed high WNPSH indices, in order Therefore the distinct difference between the two oscillation to obtain a robust response. The WNP and MC runs each modes should be attributed to the different dominant dia- consists of seven ensemble members which have initial batic heating over the northwest Pacific. conditions starting from 27 April 0000 UTC to 30 April 0000 UTC, separating by 12 h apart. All experiments were integrated from May to August. The atmospheric response 4. Model Experiment Results to the local and remote SSTA forcing is estimated based on the difference between the ensemble mean of the WNP/MC [36] To test the hypothesis that the WNPSH variability on ‐ ‐ the 3–5 year timescale is primarily forced by local SSTA in run and the CTRL run (i.e., WNP CTRL and MC CTRL). – [37] The anomalous summertime (JJA mean) precipitation the WNP while the WNPSH variability on the 2 3 year y timescale is attributed to the forcing over the maritime and 850 field responses to the WNP SSTA are shown in continent, the following control and sensitivity experiments Figure 9a. The imposed local cold SSTA causes the decrease have been developed. In the control run (CTRL), ECHAM4 of local precipitation to the southwest and an increase of precipitation in the tropical central western Pacific along was integrated for 20 years, forced by monthly climato- ‐ logical SST (based on the period of 1955–2000). In the 5°N. An anomalous low level anticyclone forms to the east of WNP run, the model was forced by the climatological SST Taiwan in response to the cold SSTA forcing. The simulated plus a uniform SSTA of −1°C over a WNP domain (150°E– patterns are qualitatively consistent with the observations 160°W, 10°–25°N) (blue box in Figure 9a). In the MC run, from the reanalysis data sets (Figures 3f, 5f, and 7b), but

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Figure 8. Vertical profiles of regressed (a) vertical p velocity, (c) Q1,and(d)Q2 averaged over the region of strongest descending motion and (b) relative vorticity averaged over the region of anomalous WNPSH center for the 2–3 year (red) and 3–5 year (blue) oscillations. the amplitude of the simulated precipitation and anticyclonic the observed anticyclone‐vertical motion phase relation. In circulation is too strong, due to the greater SSTA specified. the tropical western Pacific, anomalous easterly flows in the In the observation, the SSTA in the WNP decreases grad- southern part of the anticyclone extend to the north IO and ually from June to August [Wu et al., 2010]. converge toward the MC region. The simulated patterns in [38] The strong positive SSTA forcing in the MC induces the MC run are quite similar to those observed shown in a positive precipitation anomaly over the northern part of the Figure 3c, 5c, and 7a but with less zonal extent and a SSTA domain and a negative precipitation anomaly in the slightly northward shift due to the neglect of tropical central northern Philippines along 15°N (Figure 9b). An anomalous Pacific cold SSTA forcing in Figure 4a. anticyclone is simulated near Taiwan and is located to the [39] It is seen from Figures 9a and 9b that the spatial north of the negative precipitation anomaly center, similar to phase patterns between the anomalous precipitation center

11 of 19 D13111 CHUNG ET AL.: TROPICAL SSTA AND SUBTROPICAL HIGH D13111 and the anticyclone center in both the WNP and MC runs anomalous WNPSH through an anomalous local Hadley are similar to those observed. In addition, the vertical pro- circulation as discussed in section 3. files of the vertical velocity and relative vorticity fields averaged over the maximum precipitation and anticyclonic vorticity center in both the WNP and MC runs were eval- 5. Temporal Evolution Features uated. In Figure 9c the anomalous descending motion in the [41] To better understand the interannual variability of the MC run is larger than that in the WNP run. The strongest WNPSH, the seasonal evolution of anomalous SST and descending motion is located in the middle troposphere circulation fields of w500, y850, and V850 associated with the around 400–600 hPa in the MC run but at 500–700 hPa in WNPSH 2–3 year and 3–5 year oscillations are examined. the WNP run. In the lower troposphere below 850 hPa, More than one full cycle (from MAM(0) to JJA(2)) and the amplitude of the descent in the WNP and MC run is nearly one full cycle (from MAM(0) to JJA(3)) of regressed slightly reversed as in the observation of NCEP/NCAR R‐1 fields are presented for 2–3 year and 3–5 year oscillations. (Figure 8a). In addition to the p velocity, the vertical dis- Seasons of MAM(1), JJA(1), SON(1) denote the concurrent tributions of relative vorticity over the anticyclone center in spring, summer and autumn of the westward WNPSH, the WNP and MC run also shows a distinct difference, whereas the numbers 0, 2, and 3 in parentheses denote the similar to the observation. The anomalous anticyclone in the lead and lag years relative to the current year (1). MC run (Figure 9d) is associated with the negative vorticity throughout the troposphere that resembles the observed 5.1. SSTA and Vertical Velocity barotropic structure in the 2–3 year oscillation (Figure 8b). [42] The evolution of regressed SSTA for the 2–3 year In the WNP run (Figure 9d), the negative vorticity in the oscillation (Figure 10a) shows the following features. lower troposphere changes sign to a positive one at 400 hPa, Starting from the preceding spring MAM(0), cold SST consistent with the observed baroclinic structure in the 3– anomalies dominated the northern IO, MC, and northern 5 year oscillation (Figure 8b), although the baroclinicity in Australia, while a weak warm SSTA appeared in the equa- the model is a little stronger due to neglecting other minor torial central Pacific. The warm SSTA in the central Pacific remote SSTA forcing in the simulation. The vertical distri- moved eastward and intensified from JJA(0) to SON(0), bution of the specific humidity (figure not shown) also weakened afterward through MAM(1), and became nega- resembles similar result in observed Q2 (Figure 8d) that the tive by the current JJA(1). After JJA(1), the cold SSTA amplitude of the negative specific humidity anomaly in the strengthened until D(1)JF(2) and then decayed from MAM(2) lower troposphere is greater in the WNP run than in the MC to JJA(2) when the cold SSTA turned into a weak warm run. SSTA. Accompanying the above evolution, a weak warm [40] The horizontal and vertical structures of the anoma- SSTA in the central IO in the previous summer JJA(0) grew lous circulation in the idealized numerical experiments through SON(0). By D(0)JF(1), the warm SSTA had spread accurately capture the observed characteristics in the 2– widely to the northern IO, MC, and northern Australia. The 3 year and 3–5 year oscillations. Numerical experiments in broad‐scale warm SSTA further extended southeastward the study demonstrate the role of SSTA forcing in the WNP from D(0)JF(1) to MAM(1) and eastward from MAM(1) to and MC regions in generating the anomalous anticyclone SON(1). By the following summer JJA(2), the SSTA had associated with the westward extension of the WNPSH evolved to a pattern similar to that in MAM(0). The above in the two oscillations. The baroclinic vertical structure in evolution feature resembles the quasi‐biennial oscillation of the WNP run confirms that a Rossby wave response to the the tropical SSTA signal found in previous studies [e.g., local cold SSTA is a key process involved in the 3–5 year Barnett, 1991; Li et al., 2006]. oscillation. Since the cold SSTA in the WNP weakens from [43] Contrary to the 2–3 year oscillation, the evolution early to late summer when El Nino is decaying in the 3– of the 3–5 year oscillation shows a standing oscillation, 5 year oscillation, the SSTA forcing in the tropical IO may with the SSTA persisting from MAM(0) to SON(1) and also contribute to the westward movement of the subtropical reversing its sign from D(1)JF(2) to JJA(3) (Figure 10b). A high [Xie et al., 2009; Wu et al., 2010]. The numerical warm SSTA developed in the central eastern Pacific from experiment of Zhou et al. [2009b] has also shown that a MAM(0), reached its warmest phase in D(0)JF(1), and then warmer tropical IO may contribute to the westward exten- weakened in the following three seasons to D(1)JF(2) when sion WNPSH in the interdecadal time scale. In an additional a weak cold SSTA grew. The cold SSTA became stronger idealized experiment with a +1°C SSTA specified in the from MAM(2) to D(2)JF(3), weakened in MAM(3), and tropical IO basin (20°S–20°N, figure not shown), an continued to last for two seasons until SON(3) (figure not anomalous low‐level anticyclonic circulation is simulated in shown). Accompanying the warm SSTA in the equatorial the WNP, with a maximum center located slightly south- Pacific, the surrounding cold SSTA to the northwest and westward over the South China Sea. Since the simulated southwest also persisted from JJA(0) to SON(1). In D(1)JF(2), anticyclone differs from the observed WNPSH in its loca- the areas surrounding the equatorial Pacific were replaced tion (a southwestward shift from the observed center) and by a warm SSTA that can persist to JJA(3). In addition, warm vertical structure (the simulated baroclinic structure is much SST anomalies developed in the IO from D(0)JF(1) strength- weaker), the warm SSTA in the tropical IO may not be the ened and extended to the WNP and marginal seas along the major forcing in causing the westward shift of the WNPSH East Asia coast through MAM(1) to SON(1). In the following in the 3–5 year oscillation. In the MC run, the simulated seasons, weak cold SST anomalies appeared in the IO from horizontal pattern and the vertical structure resemble the MAM(2), extended to the north and south from JJA(2) to D(2) observed in the 2–3 year oscillation. This confirms the role JF(3), and reached the strongest intensity over the whole of the remote warm SSTA forcing in the MC in inducing the ocean basin in JJA(3). The evolution of the regressed SSTA

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Figure 9. Anomalous precipitation (shaded, only areas passing the 90% confidence level plotted) and 850 hPa stream function in (a) the WNP run and (b) the MC run, and vertical profiles of (c) anomalous vertical p velocity over the strongest descending region and (d) anomalous relative vorticity over the anticyclone center in the WNP (blue line) and the MC (red line) runs. The thick blue and red box in Figures 9a and 9b shows the region where uniform (either −1°C or +1°C) SSTA is specified.

13 of 19 D13111 CHUNG ET AL.: TROPICAL SSTA AND SUBTROPICAL HIGH D13111 patterns described above resembles a typical El Nino evo- on the 2–3 year oscillation. This differs from the 3–5 year lution with a long life cycle [e.g., Wang et al., 2000, 2003]. oscillation, in which the large‐scale upper‐layer convergent The 3–5 year oscillation of the WNPSH may be modulated flow over the tropical western Pacific persisted from the by the remote low frequency El Nino SSTA forcing [Wu and preceding summer JJA(0) to the current summer JJA(1). Zhou, 2008] via local air‐sea coupling in WNP [Wang et al., [48] The evolutions of the regressed y850 and V850 fields 2000; Sui et al., 2007]. (Figure 12) are, in general, consistent with the c200 field. In [44] The time‐longitude section of the regressed SST and the 2–3 year oscillation (Figure 12a), a strong anomalous w500 fields along the Equator from MAM(0) to JJA(2) (to anticyclone existed over southern Asia in SON(0) and JJA(3) in 3–5 year oscillation) is shown in Figure 11. For moved to the Philippine Sea in D(0)JF(1). The anomalous the 2–3 year oscillations (Figures 11a and 11c), there is an anticyclone weakened and disappeared in MAM(1) but obvious eastward propagation of the warm SSTA and reemerged over the WNP in JJA(1), again indicating a spring upward motion from the IO to the western Pacific during the barrier for the Philippine Sea anticyclone to pass through the period from JJA(0) to JJA(2). The warm SSTA propagated boreal spring on this timescale. The local Hadley circulation to the tropical western Pacific in D(0)JF(1) where the in JJA(1), therefore, is essential for the development of the upward motion developed two seasons later in JJA(1). In the summer anticyclonic circulation over the WNP on this central eastern Pacific, warm SSTA (upward motion) existed timescale. The evolution is also consistent with the vertical from MAM(0) to MAM(1) and was followed by a cold velocity fields that descending motion within the tropical SSTA (downward motion) from JJA(1) to JJA(2) with a western Pacific weakens in MAM(1), where the ascending slight eastward propagation. The center of the upward motion becomes stronger and statistically significant in JJA motion within (170°W–150°W) is located to the west of the (1) within the green box in Figures 11a and 11c. warm SSTA in the region (150°W–120°W) over the tropical [49] In the 3–5 year oscillation (Figure 12b), however, central eastern Pacific. However, over the tropical western the anomalous anticyclone over the Philippine Sea and Pacific, the two fields do not show a significant phase low‐level northeasterly wind over WNP persisted from difference. D(0)JF(1) to JJA(1), consistent with the c200,SSTAand [45] The regressed fields for the 3–5 year oscillation vertical velocity pattern in Figure 10b and Figures 11b and (Figures 11b and 11d) exhibit a standing sandwich pattern 11d. A positive thermodynamic air‐sea feedback can characterized by a warm (cold) SSTA and upward (down- explain how the anomalous anticyclone is maintained from ward) motion in the IO, a cold (warm) SSTA and downward the preceding winter to the current summer [Wang et al., (upward) motion in the western Pacific, and a warm (cold) 2000, 2003; Sui et al., 2007]. Here the anomalous anticy- SSTA with upward (downward) motion in the central east- clone, initiated by a cold SSTA as a Rossby wave response ern Pacific from MAM(0) to D(1)JF(2) [from MAM(2) to to a negative heat source in the boreal winter, may further SON(3)]. Although warm (cold) SST anomalies are gener- reinforce the cold SSTA through the wind‐evaporation‐ ally accompanied by upward (downward) motion in the SST feedback under prevailing climatological northeasterly ocean basins, this feature is not so clear in the IO where a trade winds to the spring. Although the seasonal change significant SSTA is only associated with weak vertical of prevailing winds leads to a negative air‐sea feedback motion. In the tropical western Pacific, the strongest des- and thus a weakening of the local cold SSTA in the boreal cending (ascending) motion lagged the strongest cold summer, the negative SSTA may still have a significant (warm) SSTA by 1 month, developing in JJA(0) (MAM(2)), impact on the WNP monsoon, particularly in early summer. reaching the strongest intensity in D(0)JF(1) [D(2)JF(3)], Thus this local cold SSTA may have a large impact on the and then persisting to JJA(1) (JJA(3)). summertime WNPSH variability. [46] The spatial and temporal evolutions of the SSTA in the 2–3 year and 3–5 year oscillations are consistent with the study by Barnett [1991], who found that the near‐equatorial 6. Conclusion and Discussion characteristics of the quasi‐biennial SST/SLP mode are those of a quasi‐progressive wave while the 3–7 year mode [50] The western edge of the summertime Pacific sub- is closer to a standing wave. tropical high in the Northern Hemisphere (WNPSH) exhibits a significant zonal shift on the interannual timescale. The 5.2. Stream Function and Wind largest variability is found in the western North Pacific. This [47] The examination of temporal evolution of the variability is represented in this study by an index defined by geopotential height anomalies at 500 hPa averaged in the regressed c200 (figures not shown) and w500 (Figure 11c) from JJA(0) to SON(1) indicates that in the 2–3 year oscil- region where the largest WNPSH variability occurs. A power lation, a large‐scale upper‐level divergence (ascending) was spectrum analysis of the WNPSH index indicates two – – located in the eastern Pacific and northern IO from JJA(0) dominant peaks at 2 3 year and 3 5 year periods. The to MAM(1), while a large scale upper‐level convergence oscillation characteristics of the WNPSH at the two periods (descending) appeared over the western Pacific. The pattern are investigated through a regression analysis of atmospheric was nearly out of phase in JJA(1), with a dominant upper‐ and oceanic variables against the two WNPSH indices at the – – level divergence near Java and a large‐scale convergence 2 3 year and 3 5 year periods. The idealized numerical over the tropical central eastern Pacific (Figure 5a). It should experiments with the ECHAM4 model are also carried out to examine the dominated SSTA and associated mechanisms be noted that the c200 pattern has no significant signal in the transition season MAM(1). This implies that the signal from for the WNPSH variability on the two timescales. Since the the preceding boreal summer JJA(0) to winter D(0)JF(1) may location and strength of the subtropical anticyclone greatly have difficulty to pass through the spring season MAM(1) change from June to August, the conclusion obtained from

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Figure 10. (a) Evolutions of the SSTA field from the preceding spring (MAM(0)) to the summer of the following year (JJA(2)) regressed against the WNPSH index for the 2–3 year oscillation. (b) Evolutions of the SSTA field from the preceding spring (MAM(0)) to the summer of the year after next year (JJA(3)) regressed against the WNPSH index for the 3–5 year oscillation.

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Figure 11. (a, b) Longitude‐time section of regressed SSTA and (c, d) 500 hPa vertical p velocity anomaly fields averaged between 15°S and 15°N from MAM(0) to JJA(2) for 2–3 year oscillation (Figures 11a and 11c) and from MAM(0) to JJA(3) for 3–5 year oscillation (Figures 11b and 11d). Contours in warm (cold) colors indicate warm (cold) SSTA or downward (upward) motion. Shaded regions pass the 90% confidence level. Green boxes denote the current summer of a strengthened WNPSH. this study may not be applicable to individual summer sponding vertical structure is equivalently barotropic. The months. anticyclonic circulation in the WNP is accompanied by an [51] The 2–3 year oscillation of the WNPSH and the anomalous cyclonic circulation to its north. As a result, associated SST and circulation possess a coherent eastward the overall structure appears like a meridionally oriented PJ propagating feature. The primary signals include the type pattern. The role of the remote SSTA forcing in the southeastward spread of a warm SSTA and ascending maritime continent in affecting the WNPSH variability is motion from the Indian Ocean in the preceding autumn, verified by the idealized ECHAM4 model experiment. SON(0), to the maritime continent in the current summer [52] The 3–5 year oscillation of the WNPSH exhibits a JJA(1) of the high WNPSH index. In the central and eastern standing wave feature, that the warm SSTA in the equatorial Pacific, the SSTA changes its sign from positive in the central eastern Pacific, its surrounding cold SSTA to the preceding summer through winter to negative in JJA(1), northwest and southwest, and a warm SSTA in the Indian analogous to the La Nina development condition. In the Ocean and along the East Asian marginal seas, persisting current summer, the WNPSH is characterized by an anom- from the preceding winter to the current summer of a high alous anticyclonic circulation and descending motion north WNPSH index. Along with the evolution above, the of the Philippines. The descent is a part of the anomalous ascending motion over the tropical eastern central Pacific local Hadley circulation, with enhanced convection and a and Indian Ocean and the descending motion over the warm SSTA appearing in the maritime continent. A heat western Pacific also persist from the preceding autumn to budget analysis reveals that the descent and anticyclone are the current summer. These features resemble the tripole maintained primarily by radiative cooling, and the corre- SSTA pattern and associated overturning Walker circulation

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Figure 12. Regressed 850 hPa stream function and wind fields from the preceding summer (JJA(0)) to the current autumn (SON(1)) for (a) the 2–3 year and (b) 3–5 year oscillations of WNPSH. Shadings and vectors denote regions passing the 90% confidence level. pattern frequently observed during the decaying phase of are quasi‐biennial and low‐frequency variabilities in the El Nino episodes [e.g., Wang et al., 2000, 2003]. For the 3– tropical SST/SLP fields and in Asian monsoon circulation 5 year oscillation, the maximum descent of the WNP is fields. Our findings are also consistent with the observa- slightly east of the anomalous anticyclone center. The tional fact that ENSO has both the quasi‐biennial and lower‐ greatest descent appears in the lower troposphere, accom- frequency spectral peaks. The consistency above indicates panied by a negative latent heating and a cold SSTA to the that the variability of the WNPSH connected with the Asian southeast. The corresponding vertical structure of relative Monsoon system might be closely linked to the tropical vorticity in the descent region bears a baroclinic nature, SSTA forcing in the interannual time scale. The possible and the zonal phase difference between the anomalous mechanisms that connect the subtropical high and tropical anticyclone center and the descent implies a Rossby wave SST include the monsoon‐warm ocean interaction associ- response to a negative heating anomaly in association with ated with the TBO [Li et al., 2006], ENSO‐East Asia tele- the local cold SSTA. The local SSTA forcing mechanism is connection [Wang et al., 2000; Sui et al., 2007], and a confirmed by the idealized ECHAM4 model experiment. season‐dependent Indian Ocean impact on the WNP mon- [53] Our findings of the 2–3 year and 3–5 year oscillations soon [Wu et al., 2009b]. of the WNPSH are consistent with previous studies of [54] In addition to the tropical forcing, midlatitude pro- Barnett [1991] and Wang et al. [2001] who found that there cesses may also play a role in the interannual variability of

17 of 19 D13111 CHUNG ET AL.: TROPICAL SSTA AND SUBTROPICAL HIGH D13111 the WNPSH [e.g., Enomoto, 2004]. Since the present study Hoskins, B. J. (1996), On the existence and strength of the summer sub- focus on the relationship between the WNPSH and tropical tropical anticyclones, Bull. Am. Meteorol. Soc., 77, 1287–1292. Kalnay, E., et al. (1996), The NCEP/NCAR 40‐Year Reanalysis Project, forcing, the influence of the midlatitude processes have not Bull. Am. Meteorol. Soc., 77, 437–471, doi:10.1175/1520-0477(1996) been discussed here. Our ECHAM4 model simulation 077<0437:TNYRP>2.0.CO;2. confirms that the 2–3 year mode is driven by the SSTA Kanamitsu, M., W. Ebisuzaki, J. Woollen, S.‐K. Yang, J. J. Hnilo, M. Fiorino, ‐ ‐ forcing over the maritime continent, while the 3–5 year and G. L. Potter (2002), NCEP/DOE AMIP II Reanalysis (R 2), Bull. Am. Meteorol. Soc., 83, 1631–1643, doi:10.1175/BAMS-83-11-1631. mode is driven by the local SSTA in the WNP. Despite the Kawamura, R., T. 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Wu (2001), Principal modes of rainfall‐SST variabil- needed to investigate the relative importance of these factors ity of the Asian summer monsoon: A reassessment of the monsoon‐ and their causality in causing the interannual variability of ENSO relationship, J. Clim., 14, 2880–2895, doi:10.1175/1520-0442 the WNPSH from dynamic and thermodynamic discussions. (2001)014<2880:PMORSV>2.0.CO;2. Lee, E. J., J. G. Jhun, and C. K. Park (2005), Remote connection of the Northeast Asian summer rainfall variation revealed by a newly defined monsoon index, J. Clim., 18, 4381–4393, doi:10.1175/JCLI3545.1. [55] Acknowledgments. We appreciate very much for the con- Li, T., and B. Wang (2005), A review on the western North Pacific mon- structive comments from three anonymous reviewers. This research soon: Synoptic‐to‐interannual variabilities, Terr. Atmos. Ocean. Sci., 16, was supported by NSC 98‐2111‐M‐133‐002 and NSC 99‐2111‐M‐ 285–314. 133‐001. 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