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Apr 1, 2012 - from the NOAA website (http://www.esrl.noaa.gov/psd/ data/gridded/data.noaa.ersst.html). The rainfall data analyzed here are ..... rainfall, we use a bootstrap technique (e.g., Li and Smith. 2009) to assess the correlations ...
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Remote Influence of the Tropical Atlantic on the Variability and Trend in North West Australia Summer Rainfall ZHONGDA LIN State Key Laboratory of Numerical Modeling for Atmospheric Sciences and Geophysical Fluid Dynamics, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China

YUN LI CSIRO Mathematics, Informatics, and Statistics, CSIRO Climate Adaptation Flagship, Wembley, Western Australia, Australia (Manuscript received 5 January 2011, in final form 6 October 2011) ABSTRACT Rainfall in North West Australia (NWA) has been increasing over the past decades, occurring mainly in the austral summer season (December–March). A range of factors such as decreased land albedo in Australia and increasing anthropogenic aerosols in the Northern Hemisphere, identified using simulations from climate models, have been implicated in this wetting trend. However, the impact of land albedo and aerosols on Australian rainfall remains unclear. In addition, previous studies showed that dominant sea surface temperature (SST) signals in the Pacific–Indian Ocean including El Nin˜o–Southern Oscillation (ENSO), ENSO Modoki, and the Indian Ocean dipole mode have no significant impact on the NWA rainfall trend. The present study proposes another viewpoint on the remote influence of tropical Atlantic atmospheric vertical motion on the observed rainfall variability and trend in NWA. It is found that, with the atmospheric ascent instigated by the warming of SST over the tropical Atlantic, a Rossby wave train is emanating southeastward from off the west coast of subtropical South America to the midlatitudes of the South Atlantic Ocean. It then travels eastward embedded in the westerly jet waveguide over the South Atlantic and South Indian Oceans. The eastward-propagated Rossby wave induces an anticyclonic anomaly in the upper troposphere over Australia, which is at the exit of the westerly jet waveguide. This leads to an in situ upper-tropospheric divergence, ascending motion and a lower-tropospheric convergence, and the associated increase in rainfall in NWA. Thus, the increasing trend in atmospheric upward motion induced by the warming trend of SST in the tropical Atlantic may partially explain the observed rainfall trend in NWA.

1. Introduction North West Australia (NWA; 258S northward and 1358E westward) receives the bulk of its annual rainfall during the austral summer [December–March (DJFM)]. Rainfall has been increasing in NWA since the 1950s, mainly from December to February (Shi et al. 2008; Taschetto and England 2009). This positive trend in NWA rainfall is identified not only in long-term station rainfall data in Australia since the 1950s (Shi et al. 2008), but also in other global indicators, such as outgoing longwave Corresponding author address: Zhongda Lin, State Key Laboratory of Numerical Modeling for Atmospheric Sciences and Geophysical Fluid Dynamics, Institute of Atmospheric Physics, Chinese Academy of Sciences, P.O. Box 9804, Beijing 100029, China. E-mail: [email protected] DOI: 10.1175/JCLI-D-11-00020.1 Ó 2012 American Meteorological Society

radiation and upper-troposphere divergence since the 1970s (Taschetto and England 2009). Contrary to NWA, rainfall exhibits a decreasing trend in most of East Australia, especially along the east coast (Shi et al. 2008; Taschetto and England 2009). The variability of Australian rainfall has been linked to a suite of remote climate drivers (Risbey et al. 2009). Among all climate drivers of rainfall variability in Australia, the El Nin˜o–Southern Oscillation (ENSO) plays a key role. In the austral summer season, a significant negative correlation is detected between the Nin˜o-3.4 SST index and simultaneous rainfall in eastern Australia (McBride and Nicholls 1983; Ropelewski and Halpert 1987; Shi et al. 2008), though it exhibits decadal-scale variations (Suppiah 2004). However, the ENSO’s effect on the increasing trend of NWA rainfall in austral summer season is negligible (Shi et al. 2008).

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On the other hand, Ashok et al. (2007) noticed a central Pacific warming-type ENSO (ENSO Modoki) pattern, with a warm SST anomaly in the central Pacific and cool SST anomalies in the tropical eastern Pacific and tropical western Pacific. They discussed the influence of this ENSO Modoki on global climate, but did not find a significant response in Australian rainfall in austral summer season. This weak ENSO Modoki–Australian rainfall relationship in the whole austral summer is probably due to the opposite response of Australian rainfall in early and late summer with that in midsummer to El Nin˜o Modoki events (Taschetto et al. 2010). Namely, during El Nin˜o Modoki events, it tends to be wetter in NWA in January and February, which results from an anomalous cyclone due to a Gill-type (Gill 1980) response to the intensified diabatic heating in the central-western Pacific, while it tends to be drier in NWA in December and March induced by an anomalous Walker circulation. Accordingly, ENSO and ENSOModoki has no significant influence on the rainfall variability and trend in NWA in DJFM. Risbey et al. (2009) noted the significant impact of the Indian Ocean dipole (IOD) mode on South Australian rainfall for June– October. Unfortunately, the IOD mode has no signal in DJFM, so its effect on NWA rainfall in DJFM can be excluded. In addition to the IOD mode, they also showed some other drivers, such as atmospheric blocking at 1408E on rainfall in the Northern Territory and the southern annular mode on rainfall in the middle of Australia (Risbey et al. 2009). However, these drivers cannot explain the consistently increasing trend in NWA rainfall in the austral summer season. There are some efforts made to understand the dynamics of NWA rainfall variability and trends. Wardle and Smith (2004) suggested that the observed increase in the temperature gradient between Australia and neighboring oceans might drive a stronger monsoonal circulation, resulting in the increased rainfall. Their argument was based on a climate model simulation by artificially altering the temperature contrast through the reduction of land albedo by a factor of 4. They could simulate an increase in rainfall over the entire continent, with higher totals in the north, and a temperature response similar to that observed. However, the reduction in prescribed land albedo and the resulting surface temperature changes in Australia were much larger than could be justified based on current knowledge, so the cause of the land–ocean temperature contrast still remains as an open question (Shi et al. 2008; Wardle and Smith 2004). Rotstayn et al. (2007, 2009), on the other hand, highlighted the effect of anthropogenic aerosol change in the twentieth century on the increased trend in NWA

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rainfall. They reproduced an increased trend in NWA rainfall by including anthropogenic aerosol changes in a low-resolution coupled GCM. Their results suggested that the strong impact of aerosols is mainly from the massive Asian aerosol haze, which changes the meridional temperature and pressure gradients over the tropical Indian Ocean in the model, thereby increasing the tendency of monsoonal winds toward Australia. However, as revealed by Shi et al. (2008), the modeled rainfall trend in this coupled GCM is due to an unrealistic relationship between Australian rainfall and eastern Indian Ocean sea surface temperature (SST), which suffers from a bias in the equatorial tropical Pacific cold tongue, being too far into the eastern Indian Ocean. That is, in the presence of increasing aerosols, a significant SST increase occurs in the tropical eastern Indian Ocean and attributes to the modeled rainfall increase in NWA (Shi et al. 2008; Cai et al. 2011). Therefore, the impact of aerosols on Australian rainfall is still unclear. To examine the dynamics of the observed rainfall variability in NWA, Shi et al. (2008) found that the observed positive trend in NWA summer (DJF) rainfall can be projected onto two modes of variability. The first mode is characterized by coherent change of rainfall in North Australia and associated with an anomalous low mean sea level pressure (MSLP) off the coast of NWA. The second mode is characterized by an out-of-phase rainfall variation between east and west of 1308E, which is associated with an anomalous high MSLP over most of the Australian continent. Their results also showed that the two modes are closely related to the lowertropospheric circulation and SST changes in the South Indian Ocean. However, it is not clear what triggers the atmospheric circulation and SST changes in the South Indian Ocean and around Australia. This motivates us to investigate the influence of remote climate drives on rainfall variability and trend in NWA. The above review provides a background that the ultimate cause and the influence of remote climate drivers on NWA rainfall trend and variability remain unclear. In this study, we reveal a remote forcing from the tropical Atlantic on variability and trend in NWA rainfall in the austral summer season (DJFM). The linkage between the tropical Atlantic and NWA rainfall is possibly through a midlatitude teleconnection pattern associated with the upper-troposphere westerly jet waveguide. The atmospheric circulation changes found by Shi et al. (2008) may be attributed to the responses of the teleconnection pattern over the South Indian Ocean and Australia. The rest of this paper is arranged as follows. In section 2, data used in this study are presented. We depict variability and trend of NWA summer rainfall in the austral summer season and associated local circulation in section 3, and

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then the midlatitude teleconnection pattern linking the tropical Atlantic atmospheric vertical motion to the NWA rainfall variation in section 4. The SST anomaly associated with the tropical Atlantic atmospheric vertical motion is discussed in section 5. Finally, the conclusions and discussion are presetented in section 6.

2. Data The atmospheric circulation data are extracted from 40-yr European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-40) data (Uppala et al. 2005) during the period from September 1957 to August 2002. The data analyzed here comprise the DJFM average values for each year including zonal and meridional winds; pressure vertical velocity; geopotential height at the pressure levels of 200, 500, and 850 hPa; and MSLP. In addition, the DJFM SST over the same ERA-40 period from 1957 to 2002 are created using monthly National Oceanic and Atmospheric Administration (NOAA) extended reconstructed SST (ERSST; Smith et al. 2008), which can be downloaded from the NOAA website (http://www.esrl.noaa.gov/psd/ data/gridded/data.noaa.ersst.html). The rainfall data analyzed here are provided by the National Climate Centre (NCC) of the Bureau of Meteorology (BoM) and consist of gridded data on a 0.258 3 0.258 grid. These were described by Lo et al. (2007) who indicated their confidence in the validity of the data after 1948. In addition, the target region (NWA) was selected as the box area bounded by 108–258S, 1108–1358E, which is same as that in Shi et al. (2008). The averaged summer (DJFM) rainfall totals over the ERA-40 period from 1957 to 2002 are generated (from the BoM website http:// www.bom.gov.au/cgi-bin/silo/cli_var/area_timeseries.pl) and covers areas with remarkably increasing rainfall trends in DJFM (Fig. 1). Note that the summer season used here is DJFM rather than DJF as in Shi et al. (2008) because NWA rainfall shows an increasing trend with a slope of 1.14 mm yr21, significant at the 0.01 level in March, which is comparable to that in DJF. We also use rainfall data gauged over 11 stations in NWA region, of which 9 stations are along the NWA coast and 2 stations in the middle of Australia (Fig. 1). The Nin˜o-3.4 index taken as the averaged SST (48S– 48N, 1708–1208W) based on the ERSST is used as the index of ENSO. After removing the ENSO effect (i.e., linearly removing the DJFM Nin˜o-3.4 index–related anomalies), the averaged summer rainfall totals in NWA (108–258S, 1108–1358E) is referred as the NWA rainfall index (NWARI) and the negative pressure vertical velocity at 500 hPa averaged over the tropical Atlantic region (58S–58N, 108–508W) is defined as the

FIG. 1. Observed DJFM rainfall trend (unit: mm yr21) based on the rainfall data provided by the NCC of the BoM over 1950–2007 (a) before and (b) after removing ENSO’s effect. The area in the northwest corner (solid line) is defined as NWA. (c) Variation and trend of NWA total rainfall before (black line) and after (red line) removing ENSO’s effect. The trend of NWA rainfall is 3.46 (3.83) mm yr21 before (after) removing ENSO’s effect. Here, the ENSO’s effect is removed by subtracting the linear component regressed against the DJFM Nin˜o-3.4 SST index. In (a) and (b), 11 rainfall stations in the NWA region listed in Table 1 are filled with red blank circles.

tropical Atlantic vertical velocity index (TAVVI). The TAVVI is used to depict the variability of tropical Atlantic atmospheric forcing and as a proxy of tropical Atlantic rainfall due to the absence of reliable rainfall data before 1980. The minus is added to the TAVVI for

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convenient comparison between the NWARI- and the TAVVI-related anomalies in a study field. Accordingly, when the TAVVI is positive, the atmospheric ascending motion is enhanced in the tropical Atlantic, and vice versa. Anomalies in this study are calculated by removing the associated climatology in DJFM over the ERA-40 period during 1957–2002.

3. Trend and variability of NWA rainfall a. DJFM rainfall trends Figure 1 shows the strong regional contrast of rainfall trend over Australia in DJFM. There is a clear dipole structure in the observed rainfall trend: rainfall is increasing in North West and central Australia with the center along the northwest coast of Australia, while rainfall tends to decrease in eastern Australia with the strongest decreasing trend located in the region along the northeast coast of Australia. The zonal-dipole variation of rainfall trend has also been identified in previous studies (e.g., Shi et al. 2008; Taschetto and England 2009). Some previous studies highlighted the impact of ENSO on rainfall in Australia (McBride and Nicholls 1983; Ropelewski and Halpert 1987; Lau and Nath 2000). The suppressed Australian rainfalls are mostly associated with El Nin˜o events. Consequently, the drought in eastern Australia may be partially linked to ENSO’s influence while the wet trend in NWA is reverse to ENSO’s effect because there is an upward trend in the Nin˜o-3.4 index during 1950–2007. Figure 1b shows that there are no remarkable changes in spatial distribution of DJFM rainfall trend by removing the ENSO’s effect (i.e., linearly removing the Nin˜o-3.4 index-related component from rainfall), compared with that including ENSO’s effect (Fig. 1a). Furthermore, the DJFM rainfall total series in NWA region (108–258S, 1108–1358E) shows a significant increasing trend with slope of 3.46 mm yr21, implying NWA DJFM rainfall has increased about 200 mm (;50% of the climatological value) during the period from 1950 to 2007 (Fig. 1c). After removing ENSO’s effect, the NWA rainfall increases more quickly, with a slightly stronger trend of 3.83 mm yr21. The stronger increasing rainfall trend is also noted by Shi et al. (2008). As indicated in the above analysis, the trend of ENSO during 1950–2007 is certainly unable to explain the observed wetting trend in the NWA rainfall in DJFM. Indeed, removing the influence of ENSO leads to an even stronger increasing rainfall trend. Therefore, in the rest of the paper, the effect of ENSO is removed from rainfall and other fields in order to highlight the dynamics

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of the variability and rainfall trend in NWA. The ENSOcorrected rainfall series in NWA in DJFM (i.e., NWARI; red curve in Fig. 1c), is used to represent the variation of DJFM rainfall in NWA after removing ENSO’s effect.

b. Local circulation anomalies related to NWA rainfall Figure 2 shows local circulation anomalies associated with the NWARI. Over Australia, the increased rainfall in NWA is related to an anomalous anticyclone in the upper troposphere, concurrent with anomalous divergences, with a northwest–southeast tilt, from NWA to South Australia (Fig. 2a). However, to its east and west sides, the anomalous winds converge. In the lower troposphere, an anomalous cyclone off the NWA coast induces an anomalous westerly to the north, which converges along the coast of NWA (Fig. 2c). The lowertroposphere convergence and upper-troposphere divergence over NWA is thus consistent with ascent in the midtroposphere (Fig. 2b) and in situ rainfall. On the other hand, the convergence in the upper troposphere and descent at 500 hPa along the east coast of Australia, associated with the NWARI, may favor less rainfall in eastern Australia, in contrast to that observed in NWA (Fig. 1b).

4. Remote influence of tropical Atlantic atmospheric ascent Several studies have pointed out the likely relevance of the tropical Atlantic to that of Australia rainfall. For instance, the long-term rainfall decrease in the austral winter (June–July–August) over southwest Western Australia can be traced back to a decrease in the intensity of the African monsoon (Baines 2005), which has a strong interaction with SST and zonal winds in the equatorial Atlantic (Mitchell and Wallace 1992; Okumura and Xie 2004). In this section we explore if changes in tropical Atlantic atmospheric forcing may exert some influence on rainfall in NWA in the austral summer season.

a. Relationship between the NWA rainfall and tropical Atlantic atmospheric vertical motion Figure 3 shows geopotential height anomalies in the upper (200 hPa) and lower (850 hPa) tropospheres and MSLP anomalies associated with the NWARI. In the upper troposphere (Fig. 3a), the geopotential height anomalies are characterized by a teleconnection pattern along the westerly jet with a wavenumber-5 structure in the zonal direction circling around the Southern Hemisphere. The strong signals are mostly located south of 208S and the five centers with positive anomalies are located over off the west coast of subtropical South

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FIG. 2. (a) Anomaly patterns of regressed divergence (thin contour, unit: s21), geopotential height (thick contour, unit: gpm), and wind (vector, unit: m s21) at 200 hPa upon the NWARI. (b) As in (a), but for pressure vertical velocity (contour, unit: Pa s21) at 500 hPa. (c) As in (a), but at 850 hPa. The contour interval is 2 3 1027 in (a) and (c), but 5 3 1023 in (b), and the zero contours are omitted. Scaling for the arrows is given in the top-right corner (unit: m s21). The regression coefficients are obtained by regressing divergence and wind anomalies onto the NWARI time series. Shading denotes significance at the 0.05 level.

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America, the South Atlantic Ocean, the South Indian Ocean, Australia, and southwest Pacific Ocean near 1808, respectively. The positive geopotential height anomaly over Australia coincides with the anticyclonic anomaly shown in Fig. 2a. In the lower troposphere (Fig. 3b), the spatial distribution of anomalies closely resembles that in the upper troposphere, except for the subtropical center off the west coast of South America where there is a weak negative anomaly opposite to the positive one in the upper troposphere. In addition, the positive center over Australia shifts slightly southward and a negative anomaly off the coast of NWA is identified, consistent with the cyclonic anomaly revealed in Fig. 2c. The MSLP anomalies (Fig. 3c) are almost the same as the geopotential height anomalies in the lower troposphere (Fig. 3b). After detrending, it can be seen that spatial structures remain unchanged, though the intensity is reduced (Figs. 3d–f). Note that the anomaly pattern with decreased MSLP off the NWA coast and the increased MSLP to the west over the midlatitude South Indian Ocean (Fig. 3c) is similar to the anomaly pattern of MSLP associated with the first EOF mode of NWA rainfall identified by Shi et al. (2008, their Fig. 8a). Thus, anomaly patterns of MSLP in Shi et al. (2008) are probably MSLP responses over the South Indian Ocean and Australia to the teleconnection pattern shown in Fig. 3. In other words, we may expect the tropical Atlantic to exert its influence on the variability and rainfall trend in NWA via the teleconnection. Figure 4a shows the pressure vertical velocity anomalies at 500 hPa associated with the NWARI. Here the pressure vertical velocity is used as a proxy of rainfall due to the absence of reliable global rainfall data before 1980, and the associated uncertainty should be borne in mind. Note that the negative value of pressure vertical velocity depicts the ascending motion while the positive value for the subsiding motion. As shown in Fig. 4a, there are also remarkable ascending motion anomalies in the tropical Atlantic, South America, and the Indian Ocean besides Australia, corresponding to the increased rainfall in NWA. There is no significant signal in the tropical central and eastern Pacific since the ENSO’s effect was first removed. The anomalies in the tropical Indian Ocean where the easterly dominates in the upper troposphere (as shown in Fig. 8c), however, is difficult to trigger the upper-tropospheric teleconnection pattern in the midlatitudes related to the NWARI in Fig. 3. The anomalies over South America result from the effect of tropical Atlantic (this is discussed later in Fig. 6). Therefore, the teleconnection pattern related to the NWARI is probably attributed to the effect of tropical Atlantic anomalies.

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FIG. 3. Geopotential height anomalies (unit: gpm) at (a),(d) 200 and (b),(e) 850 hPa and (c),(f) MSLP anomalies (with unit of hPa) regressed upon the NWARI. (d)–(f) As in (a)–(c), but obtained by using the detrended geopotential height, MSLP, and NWARI data. The thick solid line in (a),(d) depicts the westerly jet at 200 hPa with zonal wind exceeding 20 m s21. Shading denotes significance at the 0.05 level.

To further explore the connection between the NWARI and the tropical Atlantic ascending motion, we calculate the regionally averaged negative pressure vertical velocity at 500 hPa within the tropical Atlantic region (TAVVI; 58S–58N, 108–508W; red solid line in Fig. 4c) in which the ENSO effect has been removed. It is evident that the TAVVI and the NWARI are well coupled with correlation coefficient 0.62 during 1958–2002. Note that the link between these two indices remains (with correlation coefficient 0.51 between them, significant at the 0.01 level) after the data are detrended (Fig. 4c). This relationship between the tropical Atlantic and the NWA rainfall is confirmed by the upper-tropospheric divergence anomalies at 200 hPa associated with the NWARI (Fig. 5a). Consistent with the ascending motion anomaly, the upper-tropospheric divergence anomaly is identified over the tropical Atlantic corresponding to the enhanced rainfall in NWA. The relationship is also insensitive to the NWA rainfall region. As shown in Table 1, the TAVVI has a close relation to rainfall at these 11 stations in NWA, with correlations of 0.30– 0.50, significant at the 0.05 level. The significant

correlations remain after removing the linear trend from the data except for three stations. Moreover, the 11-station-averaged rainfall is well coupled with the TAVVI with a correlation coefficient of 0.65 (0.54) before (after) detrending, which is close to that of the box-averaged gridded rainfall (NWARI) with the TAVVI. To check the robust relationship between the TAVVI and rainfall, we use a bootstrap technique (e.g., Li and Smith 2009) to assess the correlations between the TAVVI and rainfall. This is done by resampling, 10 000 times, the TAVVI series and the DJFM rainfall with replacement, and then determining if the resulting correlation is significantly different from zero at the 0.05 level (i.e., if the TAVVI series has a strong positive link to rainfall in a specific station then the 95% confidence intervals estimated as the 5% and 95% percentiles of resampled correlations should be greater than zero). Table 1 shows 95% confidence intervals (correlations in parentheses) estimated from the 10 000 resampled correlation series between the TAVVI and rainfall from each of 11 stations and the average rainfall. It is evident that only one station (Station 14852, Victoria River Downs) yields a

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FIG. 4. Pressure vertical velocity anomalies (unit: Pa s21) at 500 hPa regressed upon (a) the NWARI and (b) the TAVVI. (c) Variation of the standardized TAVVI (red solid line), which is defined as negative pressure vertical velocity at 500 hPa averaged over the tropical Atlantic (58S–58N, 108–508W) depicted by the red rectangle in Figs. 4a,b, and the NWARI (blue dashed line) during 1958–2002. When the TAVVI is positive, the ascending motion is enhanced in the tropical Atlantic, and vice versa. The correlation coefficient between them is 0.62 (0.51) before (after) detrending. The contour interval is 5 3 1023 in (a) and (b), and zero contours are omitted. Shading denotes significance at the 0.05 level.

insignificant correlation for detrended data, with a 95% confidence interval (20.07, 0.39). As a consequence, the TAVVI has a strong positive link to rainfall over NWA. In addition, Figs. 4a and 5a also indicate that the increased NWA rainfall is associated with the ascending motion and upper-tropospheric divergence anomalies in the Amazon basin, and the subsiding motion and uppertropospheric convergence anomalies in the subtropical South Atlantic. These features are also identified in the pressure vertical velocity and upper-tropospheric divergence anomalies associated with the TAVVI (Figs. 4b and 5b). The anomalies in the Amazon basin are consistent with the upper-tropospheric anticyclonic anomaly off the west coast of South America and lowertropospheric cyclonic anomaly over eastern part of South America, as a Gill-type response (Gill 1980) to enhanced ascending motion in the tropical Atlantic (Fig. 6). Compared with the center of anticyclonic anomaly in the upper troposphere, the eastward-located cyclonic anomaly in the lower troposphere in Fig. 6 is probably due to the effect of topography. The anomalies in the subtropical Atlantic, however, may attribute to the descent branch of

meridional circulation that relates to the anomalous ascent in the tropical Atlantic. We also regress geopotential height and MSLP onto the TAVVI (figures are not shown). Anomaly patterns of geopotential height and MSLP associated with the TAVVI closely resemble those associated with the NWARI (Fig. 3). These results further suggest the enhanced tropical Atlantic atmospheric ascent may link to the variability and trend of NWA rainfall. This linkage is also demonstrated by the anomalies patterns of the longitude–pressure cross section for zonal wind and vertical velocity and horizontal divergence averaged over 108–208S regressed onto the detrended TAVVI (Fig. 7). The anomalies of ascending motion, upper-tropospheric divergence, and lower-tropospheric convergence are identified over NWA region, which favors in situ rainfall in NWA.

b. Dynamical processes of tropical Atlantic effect on the NWA rainfall In this section, we investigate the dynamical process of the remote effect of the tropical Atlantic on NWA

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FIG. 5. Divergence anomalies at 200 hPa regressed upon (a) the NWARI and (b) the TAVVI. The contour interval is 2 3 1027 s21, and zero contours are omitted. Shading denotes significance at the 0.05 level.

rainfall via the wave activity flux. Following Takaya and Nakamura (2001), the zonal and meridional components of a wave activity flux for stationary Rossby waves (W) is defined as " # 2 1 u(c9x 2 c9c9xx ) 1 y(c9x c9y 2 c9c9xy ) W5 , 2juj u(c9 c9 2 c9c9 ) 1 y(c9y2 2 c9c9 ) x y xy yy

where u 5 (u, y) denotes wind and c is the streamfunction at 200 hPa; the overbar represents the climatological mean in DJFM averaged during 1958–2002; and the subscript and prime signify partial derivatives and the anomalies associated with the TAVVI (Fig. 8) and the NWARI (Fig. 9) with the zonal mean removed, respectively. Figure 8 shows the wave activity flux at 200 hPa associated with the TAVVI. The ascent over the tropical Atlantic triggers a strong wave activity flux emanating from the positive geopotential height anomaly over the west coast of subtropical South America, southeastward through midlatitude South America into the South Atlantic, and then eastward along the latitude about 508S. The eastward wave activity flux, at the exit of the strong upper-troposphere westerly jet whose core is located over the South Atlantic and the South Indian Oceans (Fig. 8c), is divided into two branches: one continuing eastward into the South Pacific and the other diverting northeastward into Australia. The latter accumulates wave activity over Australia and forms an in situ anticyclonic anomaly. After being detrended, the eastward wave activity flux is still evident, though the amplitude is decreased (Fig. 8b). Previous studies (Hoskins and Karoly 1981; Sardeshmuhk and Hoskins 1988) demonstrated that the divergent winds caused by strong ascending motion can generate Rossby wave trains in westerlies. In the tropical eastern Pacific, the westerly (Fig. 8c) favors a poleward propagated Rossby wave into the midlatitude South Atlantic (Fig. 8a), which can be triggered by the TAVVI-related upper-tropospheric atmospheric forcing over the tropical Atlantic and South America (Fig. 5b).

TABLE 1. Correlation between the TAVVI and rainfall gauged over 11 stations in NWA (Fig. 1) in 1958–2002. The 95% confidence intervals for the correlation between the TAVVI and rainfall are shown in parentheses in the final two columns. Detrended Station No. 2014 2019 3017 3023 3028 5014 7169 14 015 14 825 15 528 15 593 Ave a b c

Station Kimberly Res Station Margaret River Mount House Station Roebuck Plains Anna Plains Mount Florance Rhodes Ridge Darwin Airport Victoria River Downs Yuendumu Alcoota ...

Statistical significance is p , 0.01. Statistical significance is p , 0.05. Statistical significance is p , 0.1.

Lat (8N) 215.65 218.63 217.05 217.93 219.25 221.79 223.10 212.42 216.40 222.26 222.82 ...

Lon (8E) 128.71 126.86 125.70 122.47 121.49 117.86 119.37 130.89 131.01 131.80 134.45 ...

No a

0.43 (0.19, 0.65) 0.49a (0.30, 0.65) 0.35b (0.16, 0.53) 0.46a (0.20, 0.68) 0.50a (0.24, 0.69) 0.41a (0.15, 0.62) 0.50a (0.26, 0.68) 0.38a (0.12, 0.58) 0.34b (0.17, 0.50) 0.30b (0.14, 0.48) 0.41a (0.20, 0.58) 0.65a (0.49, 0.77)

Yes a

0.40 (0.22, 0.56) 0.36a (0.11, 0.59) 0.22c (0.00, 0.44) 0.37a (0.06, 0.65) 0.39a (0.10, 0.63) 0.34b (0.12, 0.55) 0.29b (0.02, 0.53) 0.30b (0.05, 0.51) 0.17 (20.07, 0.39) 0.22c (0.08, 0.37) 0.41a (0.08, 0.62) 0.54a (0.34, 0.69)

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FIG. 6. (a) Anomaly pattern of regressed geopotential height (contour, unit: gpm) and wind (vector, unit: m s21) at 200 hPa upon the TAVVI. (b) As in (a), but at 850 hPa. Scaling for the arrows is given in the top-right corner (unit: m s21). Shading denotes significance at the 0.05 level for geopotential height anomalies. In (b), the thick black contour depicts the region whose topography is exceeding 1500 m.

The upper-tropospheric divergence anomalies over South America and tropical Atlantic (Fig. 5b) are consistent with the positive geopotential height anomaly off the west coast of South America. In the midlatitudes, the strong westerly jet over the South Atlantic and the South Indian Oceans, acting as a waveguide (Hoskins and Ambrizzi 1993), delivers the wave energy eastward. The stationary Rossby wavenumber of climatological 200-hPa basic flows in DJFM is presented in Fig. 8d. In mid- to high latitudes the Rossby wavenumber varies from 2 and 4, which is in agreement with the observed zonal wavenumber-3 structure with a wavelength across about 1208 at 508S where the strong westerly jet lies. In the subtropical region at 308S the Rossby wavenumber is between 4 and 8, in good agreement with the observed wavenumber-6 structure with a wavelength across about 608 over the subtropical South America, Australia, and subtropical South Pacific. The observed wavenumber-5 structure, actually, combines two waves in the mid- to high latitude over the South Atlantic and the South Indian Ocean and three waves in the subtropical regions over Australia, the South Pacific, and South America. Thus, the teleconnection pattern identified in the last section (Fig. 3) can be attributed to this wavenumber-5 pattern triggered by the atmospheric ascent anomaly over the tropical Atlantic (Fig. 8). To test the influence of tropical Atlantic atmospheric ascent on the NWA rainfall more directly, the wave

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FIG. 7. Longitude–pressure cross section averaged over 108–208S for zonal wind and vertical pressure velocity multiplied with negative 100 (vector) and horizontal divergence (contour) regressed onto the TAVVI (detrended). The contour interval is 1 3 1027 s21. The shading and black thick arrows denote significant divergence and wind anomalies at the 0.05 level, respectively.

activity flux associated with the NWARI is shown in Fig. 9. It is evident that patterns of geopotential height anomalies and the associate wave activity flux are similar to those associated with the TAVVI (Fig. 8). That is, the wave activity flux, associated with the teleconnection pattern demonstrated by eddy geopotential height anomalies in Fig. 9, emanates from off the west coast of subtropical South America at about 208S southeastward into the midlatitude South Atlantic and then eastward to Australia. After detrending, the wave train still remains with reduced intensity (Fig. 9b). However, when we further remove the effect of the tropical Atlantic atmospheric ascent by subtracting the linear component regressed against the TAVVI, the southeastward wave activity flux emanating from off the west coast of subtropical South America disappears (Fig. 9c). These results support the potential role of the tropical Atlantic on the NWA rainfall. On the other hand, the midlatitude eastward propagated wave activity flux along the strong westerly jet remains. It indicates the independence of the eastward-propagated Rossby wave train along the midlatitude westerly jet from the southeastwardpropagated Rossby wave train excited by the tropical Atlantic ascent. Note that the amplitude of the geopotential height anomaly over Australia is reduced about one-half in Fig. 9c, compared with the origin anomaly associated with the NWARI in Fig. 9a. Putting all these results together, we find that the tropical Atlantic anomalous ascent has a strong link to DJFM rainfall in NWA. This link manifests as the

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FIG. 9. (a),(b) As in Figs. 8a,b, but associated with the NWARI. (c) As in Fig. 9b, but with the detrended data, which also removed the tropical Atlantic effect by subtracting the linear component regressed against the TAVVI. The scaling of the arrow is given in the top-right corner (unit: m2 s22).

FIG. 8. (a) Eddy geopotential height anomalies (contour with unit of gpm) regressed upon the TAVVI and associated wave activity flux (vector). The scaling of the arrow is given in the top-right corner (unit: m2 s22). (b) As in (a), but obtained by using detrended data. (c) Climatology of zonal wind (unit: m s21) at 200 hPa averaged during 1958–2002. (d) The stationary Rossby wavenumbers (KS), defined in Hoskins and Ambrizzi (1993), in DJFM based on climatological 200-hPa zonal flow.

teleconnection pattern (Figs. 3, 8, and 9), which, in the upper troposphere, begins with a southeastwardpropagated Rossby wave train triggered by the tropical Atlantic ascent and the associated anomalies in the Amazon basin. It propagates from off the west coast of subtropical South America into the South Atlantic and joins the eastward-propagated Rossby wave train embedded in the strong westerly jet over the South Atlantic and the South Indian Oceans. Thus, through the southeastwardpropagated Rossby wave train, the tropical Atlantic disturbances enhance the eastward-propagated Rossby wave train and influence the NWA rainfall.

5. Possible connection to tropical Atlantic SST The above analysis shows that there exists a wellcoupled relationship between the TAVVI and NWARI (Fig. 4c), which is supported by the teleconnection instigated by the enhanced tropical Atlantic ascent. Hence, the observed upward trend in NWA rainfall may be attributed to the increasing trend in tropical Atlantic atmospheric ascent. An intuitive question is what drives

the enhanced atmospheric ascent over the tropical Atlantic. In this section, we will discuss possible SST variations related to the tropical Atlantic atmospheric ascent. Figure 10a shows the tropical Atlantic SST anomalies associated with the TAVVI. It can be seen that there is a significant positive correlation between the TAVVI and SSTs in the tropical Atlantic (Fig. 10a), and the correlation pattern does not change too much with detrended data (Fig. 10b). Together with the significant warming trend in SSTs in the tropical Atlantic (Fig. 10c), these results suggests that warming SSTs in the tropical Atlantic may cause the enhanced atmospheric ascending motion over the tropical Atlantic. Meanwhile, since the tropical Atlantic atmospheric ascending motion is positively correlated with the NWA rainfall as identified in the last section, we then expect that a warm SST anomaly in the tropical Atlantic leads to the enhanced rainfall in NWA. Figures 10d,e show the SST anomalies regressed upon the NWARI. As expected, the NWARIrelated SST anomalies do show a warm SST anomaly in the tropical Atlantic, similar to spatial pattern associated with the TAVVI, though its significant region is reduced.

6. Conclusions and discussion In this study, remote influence of the tropical Atlantic on the rainfall variation in NWA in austral summer season has been investigated. It is found that the increased NWA rainfall may relate to a midlatitude teleconnection pattern associated with enhanced atmospheric ascending

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FIG. 10. Correlation coefficients of SST in the tropical Atlantic with the TAVVI (a) before and (b) after detrending. (c) The trend of raw SST (contour) during 1958–2002 (unit: K decade21). (d),(e) As in (a),(b), but correlated with the NWARI. Shading indicates significance at the 0.05 level in (a)–(c) and 0.1 level in (d),(e).

motion in the tropical Atlantic. The teleconnection pattern is characterized by a wavenumber-5 structure in geopotential height anomalies in the upper troposphere, centered off the west coast of subtropical South America, the South Atlantic, the southwest Indian Ocean, Australia, and the southwest Pacific, respectively. Diagnosis of wave activity flux supports the teleconnection by showing that the tropical Atlantic disturbance excites a southeastward-propagated Rossby wave train from off the west coast of subtropical South America into the South Atlantic in the upper troposphere, enhancing the eastward-propagated Rossby wave train trapped in the strong westerly jet whose core is located over the South Atlantic and the South Indian Oceans and leading to the formation of a positive geopotential height anomaly over Australia in the upper troposphere. This indicates that there are anomalous divergences in the upper troposphere, resulting in ascent and convergence in the lower troposphere over Australia and the associated increased rainfall in NWA. We find that there exists a well-coupled relationship between the NWA rainfall and the tropical Atlantic vertical velocity whose upward trend is associated with the warming trend of SST in the tropical Atlantic. Thus, the increasing trend in atmospheric ascent induced by the warming of SST in the tropical Atlantic may partially explain the observed rainfall trend in NWA. It is worth mentioning that Shi et al. (2008) investigated the dynamics of the increased rainfall trend in

NWA in DJF. They showed the rainfall trend attributes to a low MSLP anomaly off the NWA coast, which is related to the enhancement of the zonal gradient of SST over the South Indian Ocean toward the West Australian (WA) coast (Fig. 11c in their paper). But it is not clear what drives the strong regional contrast of SST trend. Here we provide a mechanism that the observed regional contrast of SST trend may be due to changes in SST over the South Indian Ocean manifesting as responses to atmospheric circulations associated with the teleconnection pattern identified in this study. To see this viewpoint more clearly, we calculate the TAVVIrelated trend by projecting raw SST anomalies onto the time series of the TAVVI using the same methods as Shi et al. (2008). Figure 11 shows that the TAVVI-related SST trend is consistent with SST trend resulting from the sum of trends associated with DJF NWA rainfall EOF1 and EOF2 revealed by Shi et al. (2008, their Fig. 11a), with the warming of SST near the WA coast and the cooling over the central South Indian Ocean. In addition, as the MSLP responses to the teleconnection pattern, it shows a high MSLP anomaly over the central South Indian Ocean and a low MSLP anomaly off the NWA coast (Fig. 3). This leads to a surface poleward airflow to the east of the low MSLP anomaly, which attributes to the warming SST near the WA coast. Therefore, the teleconnection pattern in the midlatitude Southern Hemisphere enhances the zonal gradient of SST over the South Indian Ocean toward

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CSIRO. Research was partially supported by the National Basic Research Program of China (Grant 2010CB950403). Yun Li was supported by the Indian Ocean Climate Initiative Project and CSIRO Climate Adaptation Flagship. REFERENCES FIG. 11. SST trend (unit: 8C) associated with the TAVVI. The trend is calculated by projecting raw DJFM SST anomalies onto time series of the TAVVI.

the WA coast, which drives the increased rainfall in NWA (Shi et al. 2008). That is, the NWA rainfall-related zonal gradient of SST anomalies over the South Indian Ocean (Shi et al. 2008) may be linked to the enhanced atmospheric ascent induced by warming SST in the tropical Atlantic. Finally, we point out that after removing the linear trend and the tropical Atlantic atmospheric forcing effect, the wave train embedded within the westerly jet waveguide in the midlatitude still remains, though it becomes much weaker (Fig. 9). The result suggests that, in addition to the tropical Atlantic effect, there are still other possible drivers on the variation of rainfall in NWA. For example, the midlatitude storm track (or extratropical cyclone) over the South Atlantic and South Indian Oceans (Simmonds and Keay 2000) may influence the rainfall variability in NWA through eddy flux transportation feedback on the westerly jet (Nakamura and Shimpo 2004), from which the wave activities gain energy and develop eastward. Davidson et al. (2007) proposed that northeastward Rossby wave propagation from the cyclogenesis region over the South Indian Ocean toward Australia in the upper troposphere is the possible mechanism triggering the development of monsoon trough during onset of the Australian monsoon. On the other hand, in our current study, only the linear component related to ENSO is removed. Some previous studies have noticed asymmetry in the Australian rainfall response to ENSO (Power et al. 2005; Cai et al. 2010), which may exert influence on the NWARI-related teleconnection. However, those are beyond the current study. Further research on the dynamical process of other possible drivers on the variation of rainfall in NWA by using observation data and model simulation is needed in future studies. Acknowledgments. We thank three anonymous reviewers for their constructive comments. Zhongda Lin received a visiting scholar funding support through China Scholarship Council for conducting research at

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