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  • View in gallery

    (a) Time series of normalized spring [MAM(0)] AO index and the EASM index during 1948–2012. (b) Sliding correlations between the spring AO index and the EASM index displayed at the central year of a 23-yr window. The horizontal red line in (b) indicates the correlation coefficient is significant at the 5% level.

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    Anomalies of the following summer [JJA(0)] China station precipitation regressed upon the normalized spring AO index in (a) 1951–71 and (b) 1975–97. (c),(d) As in (a),(b), but the precipitation data are derived from PREC/L dataset. In (c), note that the analyzed time period is from 1949 to 1971. Dark (light) shading in the figures indicates the anomalies that are significantly different from zero at the 5% (10%) level. Contour intervals are 0.4 mm day−1 in (a)–(d).

  • View in gallery

    Anomalies of the following summer [JJA(0)] horizontal winds at 850 hPa regressed upon the normalized spring AO index in (a) 1949–71 and (b) 1975–97. Wind anomalies less than 0.3 m s−1 in both directions are not shown. Dark (light) shading in the figures denotes wind anomalies that are significantly different from zero at the 5% (10%) level for either direction.

  • View in gallery

    Anomalies of MAM(0) (a),(b) SLP; (c),(d) horizontal winds at 850 hPa; (e),(f) geopotential height at 500 hPa; and (g),(h) geopotential height at 300 hPa in (left) 1949–71 and (right) 1975–97 obtained by regressions on the normalized spring AO index. Dark (light) shading in (a),(b),(e)–(h) indicates the anomalies significant at the 5% (10%) level. Dark (light) shading in (c),(d) denotes the wind anomalies significant at the 5% (10%) level for either direction. Contour intervals are 0.25 hPa in (a),(b); 4 m in (e),(f); and 5 m in (g),(h). The wind vector scale (m s−1) is shown in the top right of (c),(d). Wind anomalies less than 0.3 m s−1 in both directions are not shown.

  • View in gallery

    Anomalies of MAM(0) SST regressed upon the normalized spring AO index in (a) 1949–71 and (b) 1975–97. Dark (light) shading in the figures indicates the anomalies that are significantly different from zero at the 5% (10%) level. Contour interval is 0.06°C.

  • View in gallery

    As in Fig. 5, but for precipitation anomalies. Contour interval is 0.2 mm day−1. Zero lines are omitted in both panels.

  • View in gallery

    Anomalies of (a),(b) AMJ(0); (c),(d) MJJ(0); and (e),(f) JJA(0) 850-hPa horizontal winds (vectors; m s−1) and SST (shadings; °C) in (left) 1949–71 and (right) 1978–98 obtained by regression upon the normalized spring AO index. Wind anomalies less than 0.3 m s−1 in both directions are not shown. Stippled areas in the figures denote the regions where SST anomalies are significantly different from zero at the 10% level.

  • View in gallery

    As in Fig. 7, but for 850-hPa horizontal winds (vectors; m s−1) and precipitation (shadings; mm day−1). Stippled areas in the figures denote the regions where precipitation anomalies are significantly different from zero at the 10% level.

  • View in gallery

    (a) The leading EOF mode of spring [MAM(0)] storm-track activity over the North Pacific (20°–70°N, 110°E–110°W). (b) The corresponding time series of the leading EOF mode. The dashed and solid horizontal lines indicate the average values in 1949–71 and 1975–97, respectively.

  • View in gallery

    (a) Storm-track activity (m) in spring [MAM(0)] average from 1949 to 1971. (b) Difference of the storm-track activity (m) in spring between the periods 1975–97 and 1949–71. Dark (light) shading indicates the anomalies that are significantly different from zero at the 5% (10%) level. Contour interval is 2 m in (b).

  • View in gallery

    Anomalies of spring [MAM(0)] 300-hPa extended EP flux (vectors; m2 s−2) and divergence of the EP flux (shading; 10−5 m s−2) in (a) 1949–71 and (b) 1975–97 obtained by regression on the normalized spring Pacific storm-track activity intensity index (multiplied by −1). The vector scale (m2 s−2) is shown in the top right of (a),(b). Extended EP fluxes less than 3 m2 s−2 in both directions are not shown.

  • View in gallery

    Difference of mean temperature (°C) between 1975–97 and 1949–71 at (a) 1000, (b) 925, (c) 850, and (d) 700 hPa. Stippled areas denote the regions where the difference is significant at the 10% level.

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An Interdecadal Change in the Relationship between Boreal Spring Arctic Oscillation and the East Asian Summer Monsoon around the Early 1970s

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  • 1 Center for Monsoon System Research, Institute of Atmospheric Physics, Chinese Academy of Sciences, and School of Earth Science, University of Chinese Academy of Sciences, Beijing, China
  • | 2 Center for Monsoon System Research, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China
  • | 3 Institute of Space and Earth Information Science, Chinese University of Hong Kong, Hong Kong, and Center for Monsoon System Research, Institute of Atmospheric Physics, Chinese Academy of Sciences, Beijing, China
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Abstract

Previous studies suggested that the boreal spring Arctic Oscillation (AO) exerts a pronounced influence on the following East Asian summer monsoon (EASM) variability. This study reveals that the relationship of spring AO with the following EASM experienced a significant interdecadal change in the early 1970s. The influence of spring AO on the following EASM is weak during the 1950s and 1960s but strong and significant during the mid-1970s through the mid-1990s. The spring AO-related sea surface temperature (SST), atmospheric circulation, and heating anomalies are compared between 1949–71 and 1975–97. Results show that the spring AO-related cyclonic circulation anomaly over the tropical western North Pacific is weaker and located more northward in the former epoch than in the latter epoch. Correspondingly, SST, atmospheric circulation, and heating anomalies over the tropical North Pacific are located more northeastward in the former than latter epoch from spring to summer. In the following summer, the spring AO-related cyclonic circulation anomalies over the tropical North Pacific are located farther away from East Asia in the former epoch. This interdecadal change in the AO–EASM connection may be attributed to a significant change in the intensity of spring North Pacific synoptic-scale eddy activity around the early 1970s from a weak regime to a strong regime, which induces a stronger eddy feedback to the low-frequency mean flow after the early 1970s. This may explain a stronger spring AO-related cyclonic circulation over the tropical western North Pacific and thus a closer relationship between the spring AO and the following EASM in the latter than former epoch.

Corresponding author address: Wen Chen, Institute of Atmospheric Physics, Chinese Academy of Sciences, 6 Beiertiao, Zhongguancun, Beijing 100190, China. E-mail: cw@post.iap.ac.cn

Abstract

Previous studies suggested that the boreal spring Arctic Oscillation (AO) exerts a pronounced influence on the following East Asian summer monsoon (EASM) variability. This study reveals that the relationship of spring AO with the following EASM experienced a significant interdecadal change in the early 1970s. The influence of spring AO on the following EASM is weak during the 1950s and 1960s but strong and significant during the mid-1970s through the mid-1990s. The spring AO-related sea surface temperature (SST), atmospheric circulation, and heating anomalies are compared between 1949–71 and 1975–97. Results show that the spring AO-related cyclonic circulation anomaly over the tropical western North Pacific is weaker and located more northward in the former epoch than in the latter epoch. Correspondingly, SST, atmospheric circulation, and heating anomalies over the tropical North Pacific are located more northeastward in the former than latter epoch from spring to summer. In the following summer, the spring AO-related cyclonic circulation anomalies over the tropical North Pacific are located farther away from East Asia in the former epoch. This interdecadal change in the AO–EASM connection may be attributed to a significant change in the intensity of spring North Pacific synoptic-scale eddy activity around the early 1970s from a weak regime to a strong regime, which induces a stronger eddy feedback to the low-frequency mean flow after the early 1970s. This may explain a stronger spring AO-related cyclonic circulation over the tropical western North Pacific and thus a closer relationship between the spring AO and the following EASM in the latter than former epoch.

Corresponding author address: Wen Chen, Institute of Atmospheric Physics, Chinese Academy of Sciences, 6 Beiertiao, Zhongguancun, Beijing 100190, China. E-mail: cw@post.iap.ac.cn

1. Introduction

The East Asian summer monsoon (EASM) is one of the most important and active components of the global climate system (Ding 1994; Chang 2004; Chang et al. 2011). The occurrence of drought and flood events caused by the unusual behavior of the EASM variability can exert substantial influences on the economy, industry, agriculture, and people’s daily life over the East Asian region (Huang and Zhou 2002; Huang et al. 2003, 2004). Hence, to skillfully forecast the EASM is of great importance for both society and economy development, which urges a better understanding of the EASM variability on both interannual and interdecadal time scales. On the interannual time scale, it has been demonstrated by previous studies that variability of the EASM is influenced by many factors, such as Eurasian snow cover (Wu and Qian 2003; Wu and Kirtman 2007; Zhao et al. 2007; Wu et al. 2009a), western Pacific subtropical high (Chang et al. 2000; Lu 2002), El Niño–Southern Oscillation (ENSO) cycle (Wang et al. 2000; Wu et al. 2003; Huang et al. 2004; Chen et al. 2013), western Pacific warm pool (Nitta 1987; Huang and Li 1987; Huang and Sun 1992), Indian summer monsoon (Wu 2002), and spring Arctic sea ice concentration (Wu et al. 2009b).

Several recent studies have found that the spring Arctic Oscillation (AO) is a very important precursory factor that could significantly influence the interannual variability of the following EASM (e.g., Gong et al. 2002; Gong and Ho 2003; Gong et al. 2011). The AO is a primary mode of climate variability over the extratropics of the Northern Hemisphere (Thompson and Wallace 1998). Gong et al. (2002) and Gong and Ho (2003) showed that interannual variability of AO in boreal late spring (i.e., April–May) is significantly connected with summer rainfall anomalies along the Yangtze River valley. They found that the East Asian summer jet stream and the rain belt tend to move northward (southward) in the positive (negative) phase of the late spring AO, which corresponds to a drier (wetter) condition in the Yangtze River valley. Using both observational data and numerical simulation outputs, Gong et al. (2011) analyzed the physical process of the influence of spring AO on the following EASM. They demonstrated that the atmospheric circulation and SST changes over the North Pacific play an important role in relaying the influence of spring AO on the following EASM. In the positive spring AO years, a significant cyclonic circulation anomaly develops over subtropical western North Pacific. This spring AO-related cyclonic circulation anomaly persists from spring to summer via a positive air–sea feedback mechanism and weakens the western North Pacific subtropical high, which subsequently reduces the northward transport of moisture from the oceans. As a result, significant positive precipitation anomalies are seen over a broad region south of 30°N extending from southern China to the western Pacific, and negative precipitation anomalies are observed in the lower valley of the Yangtze River and southern Japan.

On the other hand, it has been shown that the East Asian climate experienced a significant interdecadal change in the late 1970s (e.g., Hu et al. 1993; Yatagai and Yasunari 1994; Nitta and Hu 1996; Hu 1997; Wu and Chen 1998; Gong and Ho 2003; Qian and Qin 2008). For example, Hu (1997) reported that the rainfall and temperature in the southern and southwestern parts of China underwent a significant abrupt change in the middle and late 1970s. Gong and Ho (2003) found that the summer rainfall over the middle–lower valley of the Yangtze River experienced a notable regime shift around 1979. They showed that the rainfall before 1979 is relatively low, but afterward it increases steadily. Moreover, previous studies found that the relationship of the EASM with several climate systems (e.g., ENSO cycle and tropical Indian Ocean) experienced a significant interdecadal change in the late 1970s (Wang et al. 2008; Ding et al. 2008, 2009). The interdecadal change of the connection between the EASM and the spring AO in the late 1970s, however, has not been explored yet. Hence, an interesting question is whether there is a corresponding interdecadal change in the relation between the EASM and the spring AO. Gong et al. (2011) mentioned that correlation of spring AO with EASM is better after 1979. This implies a possible interdecadal change in the spring AO–EASM connection.

In the present study, we will present evidences to show that the connection between the spring AO and the following EASM experienced a pronounced interdecadal change at the beginning of the 1970s and further investigate the possible reasons leading to this interdecadal change. The rest of the paper is organized as follows: Section 2 describes the datasets and analysis methods. Observational evidences for the interdecadal change in the spring AO–EASM relationship are presented in section 3. Section 4 compares the spring AO-related anomalies of atmospheric circulation, sea surface temperature (SST), and precipitation before and after the interdecadal change. Section 5 investigates the possible reasons leading to the interdecadal change in the spring AO–EASM relationship. Section 6 gives a summary and discussion.

2. Datasets and methods

The present study uses the monthly-mean SST data from the National Oceanic and Atmospheric Administration (NOAA) extended reconstructed SST version 3b (Smith et al. 2008; http://www.esrl.noaa.gov/psd/data/gridded/). This SST dataset has a horizontal resolution of 2° × 2° and is available from 1854 to the present. Two monthly-mean precipitation datasets over land are also employed in this study. They are 1) the 160 Chinese land stations obtained from the National Climate Center of the China Meteorological Administration (CMA) from 1951 to the present (http://www.bcc.cma.gov.cn/) and 2) the global land precipitation dataset [Precipitation Reconstruction over Land (PREC/L)] provided by the Climate Prediction Center of NOAA (Chen et al. 2002; http://www.esrl.noaa.gov/psd/data/gridded/). The monthly-mean 160 Chinese land station dataset covers the mainland China, with relatively evenly and densely distributed stations in eastern China but sparsely distributed stations in western China. This dataset has good continuity and integrity under strict data quality control since 1951, which is widely used in China’s routine climate prediction practice (http://www.bcc.cma.gov.cn/). The PREC/L dataset is constructed on a 1° × 1° grid covering the period from 1948 to the present. Monthly-mean precipitation data over ocean are taken from the National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis on 192 × 94 Gaussian grid from 1948 to the present (Kalnay et al. 1996; ftp://ftp.cdc.noaa.gov/Datasets/ncep.reanalysis.derived/).

In addition, this study uses the monthly-mean geopotential height at 500 and 300 hPa, horizontal winds at 850 hPa, sea level pressure (SLP), and the daily-mean geopotential height and horizontal winds at 300 hPa. These atmospheric data are all obtained from the NCEP–NCAR reanalysis available from 1948 to the present and have a horizontal resolution of 2.5° × 2.5° (Kalnay et al. 1996). We also use the daily-mean geopotential height from the 40-yr European Centre for Medium-Range Weather Forecasts Re-Analysis (ERA-40) from 1958 to 2001 (Uppala et al. 2005). The AO index is defined as the principle component corresponding to the leading empirical orthogonal function (EOF) mode of monthly SLP anomalies over the Northern Hemisphere (NH) extratropics (20°–90°N) (Thompson and Wallace 1998). The EASM index proposed by Wang and Fan (1999) is adopted in the present study to measure the intensity of the EASM. This EASM index is defined as the difference in region-averaged U850 between 5°–15°N, 90°–130°E and 22.5°–32.5°N, 110°–140°E, where U850 represents the zonal wind at 850 hPa averaged from June to August (JJA). Wang et al. (2008) found that the EASM index defined by Wang and Fan (1999) can well capture the total variance of the precipitation and three-dimensional atmospheric circulation over East Asia. In addition, the Pacific decadal oscillation (PDO) index used in this study is extracted from the Joint Institute for the Study of the Atmosphere and Ocean at their website (http://jisao.washington.edu/pdo/PDO.latest), which is available from 1900 to the present.

In this study, the analysis is focused on interannual variation. Hence, the monthly-mean data and the AO index are all subjected to a 7-yr high-pass Lanczos filter (Duchon 1979). Previous studies have found that ENSO events could exert substantial influences on the climate anomalies over East Asia (Wang et al. 2000), which may confuse the influence of spring AO on the following EASM. Because ENSO has a phase-locking feature with a mature peak in boreal winter, we subtract the ENSO signal in the variables by means of linear regression with respect to the Niño-3.4 index (area-averaged SST anomalies over the region of 5°S–5°N, 170°–120°W) during the period from the preceding November–December (ND) to the next January–March (JFM) [i.e., ND(−1)–JFM(0)]. Removal of signals associated with ND(−1)–JFM(0) Niño-3.4 index from the analysis data can also exclude the possibility that the significant correlation between spring AO and EASM is due to the influence of ND(−1)–JFM(0) Niño-3.4 index on both the spring AO and EASM. Here and for the remainder of this study, notations “(−1)” and “(0)” refer to the year before and during the boreal spring AO year, respectively.

3. Interdecadal change in the spring AO–EASM connection

The relationship between the spring AO and the EASM is not stable. Figure 1a shows the interannual anomalies of the normalized spring AO index and the following EASM index during 1948–2012. The spring AO index and the EASM index display both same-sign and opposite-sign anomalies during the analysis period. While same-sign anomalies appear to dominate the period from the late 1960s through the mid-1990s, opposite-sign anomalies are often seen during the 1950s through the mid-1960s and the late 1990s through the mid-2000s.

Fig. 1.
Fig. 1.

(a) Time series of normalized spring [MAM(0)] AO index and the EASM index during 1948–2012. (b) Sliding correlations between the spring AO index and the EASM index displayed at the central year of a 23-yr window. The horizontal red line in (b) indicates the correlation coefficient is significant at the 5% level.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

The interdecadal changes in the relationship of the spring AO with the EASM is clearly seen in the sliding correlations between the spring AO index and the EASM index with a window of 23 yr (Fig. 1b). Years shown in Fig. 1b are labeled according to the central year of the 23-yr window. The correlation is positive and significant at the 5% level during 1971–94, while the positive correlation before 1970 and after 1995 is below the 5% level (Fig. 1b). The increase in the positive correlation from 1970 to 1971 (Fig. 1b) is attributed to the replacement of the large opposite-sign anomalies of the spring AO index and the EASM index in 1959 by the large same-sign anomalies in 1982 in the 23-yr sliding window (Fig. 1a). In addition, the decrease in the positive correlation from 1994 to 1995 is due to the opposite-sign anomalies in 2006 replacing the same-sign anomalies in 1983 (Fig. 1a). Change in the spring AO–EASM correlation around the early 1970s and the mid-1990s can also be obtained when using different EASM indices (e.g., Wang et al. 1998; Li and Zeng 2002; Zhang et al. 2003). It should be mentioned that change in the spring AO–EASM correlation around the mid-1990s is consistent with the finding of Gao et al. (2014). Gao et al. (2014) showed that the relationship between interannual variations of spring [averaged from March to May (MAM)] AO and EASM [averaged from May to July (MJJ)] experienced a remarkable interdecadal change around 1997. They indicated that North Pacific plays a dominant role in memorizing and relaying the influence of spring AO on EASM during 1979–97, whereas the spring AO-related signals are mainly memorized over the Indian Ocean during 1998–2007.

In this study, we focus on investigating the interdecadal change in the spring AO–EASM correlation around the early 1970s. Based on the sliding correlation with a 23-yr window (Fig. 1b), we select the periods of 1949–71 and 1975–97 in the contrasting analysis throughout this study. These two periods are selected because the difference in the correlation between these two 23-yr periods is the largest for the analysis period.

The distribution of spring AO-related precipitation anomalies in the following summer display remarkable differences over East Asia between the two epochs. Figures 2a and 2b show anomalies of China station precipitation in the following summer [JJA(0)] regressed upon the normalized spring AO index for the periods 1951–71 and 1975–97, respectively. Note that the 160 Chinese station precipitation is only available starting from 1951. During 1975–97, significant and positive precipitation anomalies are observed over southern China, and significant negative precipitation anomalies are seen over the mei-yu rain belt (along the Yangtze River) (Fig. 2b). During 1951–71, there is no significant signal over southern China or mei-yu rain-belt region (Fig. 2a). The above striking contrast of spring AO-related precipitation anomalies during these two epochs is further confirmed based on the PREC/L data (Figs. 2c,d). During 1949–71, no significant precipitation anomalies can be found over southern China, the Yangtze River, or southern Japan. During 1975–97, significant negative precipitation anomaly is seen over the Yangtze River and the anomalous rain belt extends northeastward to southern Japan (Fig. 2d). Because the EASM variability is closely related to summer precipitation anomalies along the Yangtze River to southern Japan (Wang et al. 2001, 2008), the pronounced contrast in the spring AO-related precipitation anomalies along the Yangtze River to southern Japan (Fig. 2) implies that the spring AO–EASM connection is weak during 1949–71. In contrast, during 1975–97, the spring AO could exert pronounced influences on the EASM variability.

Fig. 2.
Fig. 2.

Anomalies of the following summer [JJA(0)] China station precipitation regressed upon the normalized spring AO index in (a) 1951–71 and (b) 1975–97. (c),(d) As in (a),(b), but the precipitation data are derived from PREC/L dataset. In (c), note that the analyzed time period is from 1949 to 1971. Dark (light) shading in the figures indicates the anomalies that are significantly different from zero at the 5% (10%) level. Contour intervals are 0.4 mm day−1 in (a)–(d).

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

Note that both station and PREC/L data show significant positive precipitation anomalies in the central part of north China in the early period, with the PREC/L precipitation anomalies extending farther northward to Mongolia (Figs. 2a,c). The mechanism underlying the link between spring AO and above-mentioned positive precipitation anomalies is still unclear and should be further investigated.

Changes consistent with precipitation anomalies are identified in summer atmospheric circulation anomalies over East Asia and western North Pacific. Figures 3a and 3b contrast anomalies of JJA(0) winds at 850 hPa as obtained by regression on the normalized spring AO index between 1949–71 and 1975–97. During 1975–97, a significant cyclonic circulation anomaly is observed over tropical western North Pacific, accompanied by significant easterly wind anomalies in the region spanning from the Pacific to southeastern China along 25°–30°N and significant westerly wind anomalies to the south of 15°N, extending from the South China Sea to the tropical western–central Pacific (Fig. 3b). This distribution of 850-hPa wind anomalies during 1975–97 is generally similar to that obtained by Gong et al. (2011) using the NCEP–NCAR reanalysis data for the period 1951–2008 (see their Fig. 2). The northeasterly wind anomalies over southeastern China (Fig. 3b) would reduce the water vapor transportation from the oceans to the Yangtze River region and lead to negative precipitation anomalies there and positive precipitation anomalies over southeastern China (Fig. 2b). During 1949–71, a pronounced cyclonic circulation anomaly can also be observed over subtropical western North Pacific (Fig. 3a), but its position shifts northeastward compared to that during 1975–97 (Fig. 3b). In particular, wind anomalies over eastern China, the South China Sea, and the Maritime Continent are very weak (Fig. 3a). This implies that the influence of the spring AO on the following EASM variability may be very weak during this period, consistent with the result obtained from Fig. 2. Hence, the stronger connection between the spring AO and the following EASM in 1975–97 compared to 1949–71 may be attributed to the fact that the spring AO-related low-level cyclonic circulation anomaly over the western North Pacific in JJA(0) is located closer to southeastern Asia during the later period (Fig. 3b), while it shifts northeastward far away from the Asian continent in the earlier period (Fig. 3a).

Fig. 3.
Fig. 3.

Anomalies of the following summer [JJA(0)] horizontal winds at 850 hPa regressed upon the normalized spring AO index in (a) 1949–71 and (b) 1975–97. Wind anomalies less than 0.3 m s−1 in both directions are not shown. Dark (light) shading in the figures denotes wind anomalies that are significantly different from zero at the 5% (10%) level for either direction.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

4. Changes of spring AO-related atmospheric circulation, SST, and precipitation anomalies between 1949–71 and 1975–97

In this section, we investigate in detail the differences in atmospheric circulation, SST, and atmospheric heating anomalies from spring to the following summer between the periods 1949–71 and 1975–97. We examine first the anomalies in simultaneous spring and then the evolution of anomalies from spring to summer.

a. Atmospheric circulation, SST, and precipitation anomalies in simultaneous spring

Figure 4 shows anomalies of SLP, winds at 850 hPa, and geopotential height at 500 and 300 hPa in MAM(0) as obtained by regression on the normalized spring AO index for 1949–71 and 1975–97, respectively. The spring AO-related SLP anomalies exhibit significant differences over the North Pacific and East Asia between the two periods (Figs. 4a,b). During 1949–71, significant positive SLP anomalies are observed along the eastern coast of Eurasia and northern North Pacific (Fig. 4a), and significant negative SLP anomalies are seen over the high latitudes of Asia and tropical central North Pacific (Fig. 4a). Positive SLP anomalies over northern North Pacific are stronger and more significant during 1975–97 than during 1949–71 (Fig. 4b). Around the Korean Peninsula, significant negative SLP anomalies are present during 1975–97 (Fig. 4b), whereas SLP anomalies are significantly positive during 1949–71 (Fig. 4a). Furthermore, the regions with significant negative SLP anomalies over tropical central North Pacific extend eastward and southward during 1975–97 compared to those during 1949–71 (Figs. 4a,b).

Fig. 4.
Fig. 4.

Anomalies of MAM(0) (a),(b) SLP; (c),(d) horizontal winds at 850 hPa; (e),(f) geopotential height at 500 hPa; and (g),(h) geopotential height at 300 hPa in (left) 1949–71 and (right) 1975–97 obtained by regressions on the normalized spring AO index. Dark (light) shading in (a),(b),(e)–(h) indicates the anomalies significant at the 5% (10%) level. Dark (light) shading in (c),(d) denotes the wind anomalies significant at the 5% (10%) level for either direction. Contour intervals are 0.25 hPa in (a),(b); 4 m in (e),(f); and 5 m in (g),(h). The wind vector scale (m s−1) is shown in the top right of (c),(d). Wind anomalies less than 0.3 m s−1 in both directions are not shown.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

Pronounced differences also appear in the spring AO-related 850-hPa wind anomalies over the North Pacific between the two periods (Figs. 4c,d). During 1975–97, a significant cyclonic circulation anomaly is observed over the tropical central North Pacific and a pronounced anticyclonic circulation anomaly is seen over the midlatitudes of North Pacific, together with significant westerly wind anomalies over the high latitudes and tropical western–central North Pacific and significant easterly wind anomalies over the midlatitudes of North Pacific around 35°–50°N (Fig. 4d). During 1949–71, the anticyclonic circulation anomalies over the midlatitudes of the North Pacific and the cyclonic circulation anomalies over the tropical North Pacific are weaker compared to those in 1975–97. Correspondingly, the spring AO-related westerly wind anomalies over the tropical western–central Pacific are much weaker and shift northward in the earlier period compared to those in the later period (Figs. 4c,d).

In the middle–upper troposphere, the spring AO-related geopotential height anomalies over the North Pacific also display remarkable differences between the two periods (Figs. 4e–h). During 1975–97, significant positive geopotential height anomalies are seen over the midlatitudes of the North Pacific and significant negative geopotential height anomalies are observed over subtropical North Pacific (Figs. 4f,h). The spring AO-related geopotential height anomalies over the North Pacific display a barotropic structure (Figs. 4b,d,f,h), resembling the traditional boreal wintertime AO structure obtained by previous study (Thompson and Wallace 1998, 2000).During 1949–71, positive geopotential height anomalies over the midlatitudes northeastern Pacific and negative geopotential height anomalies over the subtropical northern Pacific are much weaker and less significant compared to those during the period 1975–97 (Figs. 4e–h).

It can be concluded from Fig. 4 that the most remarkable difference of the spring AO-related atmospheric circulation anomalies between the two periods appears in the tropical and subtropical North Pacific. During 1975–97, significant westerly wind anomalies can be observed in the tropical western–central North Pacific and significant negative geopotential heights anomalies can be seen in the subtropical North Pacific (Figs. 4d,f,h). In contrast, during 1949–71, the westerly wind anomalies over tropical Pacific and the negative geopotential height anomalies over subtropical Pacific are much weaker and less significant compared to those during 1975–97 (Figs. 4c,e,g).

Corresponding to the above different atmospheric circulation anomalies, the SST anomaly pattern displays notable differences as well. Figures 5a and 5b show the SST anomalies in MAM(0) obtained by regression upon the normalized spring AO index during the two periods, respectively. During 1949–71, significant negative SST anomalies are seen in the equatorial central Pacific, tropical eastern Indian Ocean, the South China Sea and subtropical western North Pacific (Fig. 5a). In addition, significant positive SST anomalies are observed in the midlatitudes of northwestern Pacific and in the equatorial western Pacific around 150°E extending northeastward to subtropical northeastern Pacific (Fig. 5a). During 1975–97, pronounced negative SST anomalies are seen in the tropical eastern Indian Ocean and the subtropical central Pacific between 150°E and 160°W (Fig. 5b). The most pronounced differences between the two periods appear in the equatorial central Pacific (Figs. 5a,b). During 1975–97, positive SST anomalies cover the equatorial central Pacific (Fig. 5b). However, during 1949–71, the equatorial central Pacific is dominated by negative SST anomalies (Fig. 5a).

Fig. 5.
Fig. 5.

Anomalies of MAM(0) SST regressed upon the normalized spring AO index in (a) 1949–71 and (b) 1975–97. Dark (light) shading in the figures indicates the anomalies that are significantly different from zero at the 5% (10%) level. Contour interval is 0.06°C.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

Using ocean model simulations, Gong et al. (2011) and Hu et al. (2013) demonstrated that the spring AO-related SST anomalies in simultaneous spring are triggered by atmospheric circulation changes. During 1975–97, significant westerly wind anomalies over the tropical western–central Pacific would reduce the climatological easterly winds and subsequently weaken the evaporation, vertical water mixing, and upwelling of cool water and lead to SST warming there (Figs. 4d, 5b). In contrast, during 1949–71, the spring AO-related westerly wind anomalies over the tropical western Pacific are very weak. In addition, easterly wind anomalies can be seen over the tropical central–eastern Pacific, which would enhance the climatological easterly trade winds and result in SST cooling via changing surface heat fluxes and upwelling of cool water (Figs. 4c, 5a).

We have further examined the spring AO-related atmospheric heating anomalies in MAM(0) during the two periods. To a large extent, precipitation anomalies can be used to denote the atmospheric heating anomalies (e.g., Yu and Zwiers 2007; Chen et al. 2014). Figure 6 shows anomalies of precipitation in MAM(0) in association with the spring AO during 1949–71 and 1975–97, respectively. During 1949–71, significant positive precipitation anomalies are observed in the region extending from the western Pacific to the eastern Pacific along 10°–20°N and significant negative precipitation anomalies are seen over the equatorial central Pacific and to extend from east of Taiwan northeastward to south of Japan (Fig. 6a). During 1975–97, significant positive precipitation anomalies are also seen over the tropical North Pacific (Fig. 6b). The regions with significant negative precipitation anomalies over the tropical central North Pacific during the period 1975–97 are smaller and shift northward compared to those in 1949–71 (Figs. 6a,b). Significant negative precipitation anomalies near the equatorial central Pacific during 1949–71 may be attributed to negative SST anomalies there (Figs. 5a, 6a).

Fig. 6.
Fig. 6.

As in Fig. 5, but for precipitation anomalies. Contour interval is 0.2 mm day−1. Zero lines are omitted in both panels.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

b. Evolution of atmospheric circulation, SST, and precipitation anomalies

Figure 7 displays evolutions of 850-hPa winds and SST anomalies from April–June(0) [AMJ(0)] to JJA(0) obtained by regression on the normalized spring AO index for 1949–71 and 1975–97, respectively. Correspondingly, Fig. 8 displays evolutions of atmospheric heating (denoted by precipitation) anomalies. Also included in Fig. 8 are 850-hPa wind anomalies as in Fig. 7. Generally speaking, evolutions of SST, 850-hPa wind, and atmospheric heating anomalies during 1975–97 (Figs. 7b,d,f and 8b,d,f) are similar to the results obtained by Gong et al. (2011). The wind anomalies display a dipole structure over the North Pacific with an anomalous anticyclonic circulation over the midlatitudes of the North Pacific and a cyclonic circulation anomaly over the tropical western North Pacific (Fig. 4d). Westerly wind anomalies to the south of the anomalous cyclone produce SST warming between 150°E and 180° near the equator via reduction of upward net surface heat flux and suppression of upwelling of cool water (Figs. 4d, 5b, and 7b). The atmospheric heating anomalies near the equator induced by the positive SST anomalies (Figs. 8b,d,f), in turn, intensify the anomalous cyclonic circulation over the tropical western North Pacific via a Gill-type response of the atmosphere (Matsuno 1966; Gill 1980). In addition to the baroclinic Gill-type response of the atmosphere to tropical heating, a stationary barotropic wave response can also contribute to the maintenance of the anomalous cyclonic circulation over the tropical western North Pacific (e.g., Lee et al. 2009). Through above-mentioned positive air–sea feedback mechanism, the spring AO-related cyclonic circulation anomaly persists from spring to summer (Figs. 7b,d,f and 8b,d,f). Subsequently, it influences the western North Pacific subtropical high and modifies the EASM precipitation. During 1949–71, the tropical central Pacific near the equator from spring to summer is dominated by negative SST anomalies and negative atmospheric heating anomalies (Figs. 7a,c,e and 8a,c,e). The positive SST anomalies and atmospheric heating anomalies are located more northeastward over the North Pacific from AMJ(0) to JJA(0) compared to those during the period 1975–97 (Figs. 7, 8). Hence, coupled with the positive SST and atmospheric heating anomalies, the corresponding cyclonic circulation anomalies over western North Pacific from spring to summer are located more northeastward in the earlier period than in the later period (Figs. 7, 8).

Fig. 7.
Fig. 7.

Anomalies of (a),(b) AMJ(0); (c),(d) MJJ(0); and (e),(f) JJA(0) 850-hPa horizontal winds (vectors; m s−1) and SST (shadings; °C) in (left) 1949–71 and (right) 1978–98 obtained by regression upon the normalized spring AO index. Wind anomalies less than 0.3 m s−1 in both directions are not shown. Stippled areas in the figures denote the regions where SST anomalies are significantly different from zero at the 10% level.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

Fig. 8.
Fig. 8.

As in Fig. 7, but for 850-hPa horizontal winds (vectors; m s−1) and precipitation (shadings; mm day−1). Stippled areas in the figures denote the regions where precipitation anomalies are significantly different from zero at the 10% level.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

From the above analysis, the differences in the location of cyclonic circulation anomaly in JJA(0) between the two periods are attributed to the differences in the distribution of SST and atmospheric heating anomalies over the tropical central Pacific. The positive SST anomalies are located more northeastward from spring to summer during 1949–71 than during 1975–97 (Figs. 7, 8). On the other hand, the differences in the SST anomalies in MAM(0) may be mainly resulted from the differences in atmospheric circulation anomalies in simultaneous spring over the tropical North Pacific in response to AO (Gong et al. 2011; Hu et al. 2013). Thus, the SST anomalies in the tropical central Pacific play a role in connecting spring AO to EASM variability. In the following section, the possible reasons responsible for the changes in the spring AO-related cyclonic circulation anomalies in MAM(0) over the tropical North Pacific in the two periods will be discussed.

5. Possible causes of the changes in circulation anomalies

It has been demonstrated by previous studies that the interaction between synoptic-scale eddy activity and the low-frequency flow is an important internal source of the low-frequency variability of the atmospheric circulation (Jin et al. 2006a,b). Gong et al. (2011) and Chen et al. (2014) demonstrated that feedback of synoptic eddies to the low-frequency mean flow plays an important role in the formation of the spring AO-related cyclonic circulation anomalies over the tropical western North Pacific. As suggested by Gong et al. (2011) and Chen et al. (2014), in the positive spring AO years, the westerly jet stream weakens significantly over the midlatitudes of the North Pacific. The weakening of the westerly jet stream is accompanied by weakened synoptic-scale eddy activity, as well as negative geopotential height tendency immediately to its south and positive geopotential height tendency to its north (Lau 1988). Hence, the eddy feedback to low-frequency mean flow may explain the formation of the spring AO-related cyclonic circulation anomalies over the tropical western North Pacific. The reverse is true when the spring AO is in its negative phase. During 1949–71, the spring AO-related cyclonic circulation anomaly over the tropical western North Pacific is weaker than that in 1975–97 (Figs. 4c,d). Hence, it is reasonable to hypothesize that the strength of the synoptic-scale eddy feedback to low-frequency flow may be stronger during 1975–97 than during 1949–71. Previous studies have demonstrated that strength of the synoptic-scale eddy feedback to the low-frequency mean flow is determined by several factors, including the intensity of the synoptic-scale eddy activity, eddy life time scale, and eddy spatial length scale (e.g., Jin et al. 2006a,b). In particular, the intensity of synoptic-scale eddy activity is a key component among those factors in determining the strength of the eddy feedback. Specially, when the magnitude of the low-frequency mean flow remains the same, the synoptic-scale eddy feedback strength is proportional to the intensity of the eddy activity (Jin 2010). Therefore, it is hypothesized that the intensity of the boreal springtime synoptic-scale eddy activity is stronger in 1975–97 than in 1949–71. In the following, we will present evidence to demonstrate that this is indeed the case.

Following previous studies, the synoptic-scale eddy is defined as the root-mean-square of the 2–8-day bandpass-filtered geopotential height at standard pressure levels (e.g., Chang and Fu 2002; Lee et al. 2012). In the present study, we only show the results at 300 hPa because results at other pressure levels bear a close resemblance (not shown). An EOF analysis is performed to obtain the dominant mode of the variability of the boreal spring Pacific storm-track activity during the period 1948–97. The region selected to conduct EOF analysis is from 110°E to 110°W and from 20° to 70°N. The data field was weighted to account for the decrease of area toward the pole before performing EOF analysis (North et al. 1982a).

Figure 9 displays the leading EOF mode of boreal spring synoptic-scale eddy activity over the North Pacific and the corresponding time series. It explains about 30% of the total variance and is separated from the other eigenvalues based on the criterion of North et al. (1982b). The leading EOF mode is characterized by a monopole pattern (Fig. 9a) and it depicts the variability in the intensity of the synoptic-scale eddy activity (Lau 1988; Lee et al. 2012). The corresponding principal component (PC) time series (Fig. 9b) is used as the intensity index of the boreal spring synoptic eddy activity over the North Pacific. The index shows a clear transition in the early 1970s. Before the early 1970s, the North Pacific synoptic eddy activity is generally weak, while the synoptic eddy activity becomes obviously strong after the early 1970s. As there are potential data problems before the satellite era (1979) in the NCEP–NCAR reanalysis (Kistler et al. 2001), we utilize the daily-mean geopotential height from ERA-40 (not shown) to confirm the results based on the NCEP–NCAR reanalysis. Results show that the transition from weak to strong storm-track activity regime in spring over North Pacific during the early 1970s can also be captured by ERA-40. It should be mentioned that this interdecadal change in the springtime North Pacific storm-track activity has also been found in the boreal cold season. For instance, utilizing radiosonde observations and the NCEP–NCAR reanalysis, Chang and Fu (2002) showed that the North Pacific storm-track activity averaged from December to January was weak before the early 1970s and became strong afterward. Using the NCEP–NCAR reanalysis and ERA-40, Lee et al. (2012) reported that the North Pacific storm-track intensity averaged from November to March experienced a significant transition around the early to mid-1970s.

Fig. 9.
Fig. 9.

(a) The leading EOF mode of spring [MAM(0)] storm-track activity over the North Pacific (20°–70°N, 110°E–110°W). (b) The corresponding time series of the leading EOF mode. The dashed and solid horizontal lines indicate the average values in 1949–71 and 1975–97, respectively.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

Figure 10a displays climatology of the spring North Pacific synoptic eddy activity during 1949–71, and Fig. 10b shows the difference of the synoptic eddy activity between 1975–97 and 1949–71. The difference of the North Pacific synoptic eddy activity between the earlier and later period bears some resemblances to the spatial pattern of the leading EOF mode shown in Fig. 9a. This confirms that the spring North Pacific synoptic eddy activity has experienced a dramatic interdecadal change, with stronger eddy activities extending from East Asia to North Pacific along 30°–50°N during 1975–97 than during 1949–71 (Fig. 10b). This implies that the strength of the synoptic-scale eddy feedback is stronger in the later period than in the earlier period.

Fig. 10.
Fig. 10.

(a) Storm-track activity (m) in spring [MAM(0)] average from 1949 to 1971. (b) Difference of the storm-track activity (m) in spring between the periods 1975–97 and 1949–71. Dark (light) shading indicates the anomalies that are significantly different from zero at the 5% (10%) level. Contour interval is 2 m in (b).

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

To substantiate the increase in the strength of the synoptic eddy feedback to low-frequency mean flow, we show anomalies of extended Eliassen–Palm (EP) flux and the corresponding divergence at 300 hPa in spring obtained by regression on the normalized spring North Pacific synoptic eddy intensity index during the two periods in Fig. 11. Following Trenberth (1986), the horizontal flux components of the extended EP flux is written as , where u′ and υ′ denote synoptic-scale zonal and meridional winds, respectively, and φ is latitude. The synoptic eddy fields were obtained by applying a Butterworth bandpass filter to retain fluctuations with a 2–8-day period (Murakami 1979). The overbar indicates the time mean from March to May. As suggested by previous studies (e.g., Trenberth 1986; Lau 1988; Cai et al. 2007), divergence of the extended EP flux can be used to qualitatively describe the low-frequency atmospheric circulation anomalies induced by the synoptic-scale eddy feedback because of the barotropic process. A local EP flux divergence (convergence) is accompanied by the forcing of cyclonic (anticyclonic) vorticity to the north of the divergence (convergence) region and anticyclonic (cyclonic) forcing to the south (Lau 1988). From Fig. 11, the weakening of the North Pacific synoptic eddy activity is accompanied by the convergence of EP flux over the midlatitude of North Pacific around 40°N during the two periods. However, the anomalous EP flux convergence over the midlatitudes of North Pacific is stronger during 1975–97 than during 1949–71. In addition, during the latter period, the EP flux convergence region shows an obvious expansion. These results demonstrate that the strength of the synoptic eddy feedback to low-frequency mean flow is stronger during 1975–97 than during 1949–71.

Fig. 11.
Fig. 11.

Anomalies of spring [MAM(0)] 300-hPa extended EP flux (vectors; m2 s−2) and divergence of the EP flux (shading; 10−5 m s−2) in (a) 1949–71 and (b) 1975–97 obtained by regression on the normalized spring Pacific storm-track activity intensity index (multiplied by −1). The vector scale (m2 s−2) is shown in the top right of (a),(b). Extended EP fluxes less than 3 m2 s−2 in both directions are not shown.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

The above results suggest that the difference in the eddy feedback strength between the two periods may explain the differences in the strength of the spring AO-related cyclonic circulation anomalies over the tropical western North Pacific. The eddy feedback strength is stronger during 1975–97 than during 1949–71, leading to stronger spring AO-related cyclonic circulation anomaly over the tropical western North Pacific as well as stronger westerly wind anomalies to the south of this cyclonic circulation.

6. Summary and discussion

Previous studies have demonstrated that the spring AO could exert a significant influence on the subsequent EASM variability (e.g., Gong et al. 2011). In this study, we present evidences to demonstrate that the influence of spring AO on the following EASM experienced a significant interdecadal change in the early 1970s. During 1949–71, the influence of spring AO on EASM is insignificant, whereas the spring AO is able to exert a pronounced influence on the EASM during 1975–97.

The interdecadal change in the spring AO–EASM connection in the early 1970s is related to the interdecadal change in the spring AO-related simultaneous cyclonic circulation anomalies over the tropical western North Pacific. During 1975–97, a significant cyclonic circulation anomaly in association with the spring AO is seen over the tropical western North Pacific. The pronounced westerly wind anomalies to the south of the spring AO-related cyclonic circulation anomalies lead to SST warming over the tropical western–central Pacific near the equator via reducing the upward net surface heat flux and upwelling of cool water. Then, the spring AO-related cyclonic circulation anomaly over the tropical western North Pacific persists from spring to summer through a positive air–sea feedback mechanism and subsequently weakens the western North Pacific subtropical high and finally influences the EASM precipitation. In contrast, during 1949–71, the spring AO-related cyclonic circulation anomaly over the tropical western North Pacific and the associated westerly wind anomalies to the south are weaker and shift northward compared to those in 1975–97. Hence, the induced SST warming and atmospheric heating anomalies over the tropical North Pacific in spring are located more northeastward compared to those in the later period. As a result, the spring AO-related cyclonic circulation anomalies in the following summer over the North Pacific are located far away from East Asia, leading to insignificant spring AO–EASM relationship in this period.

The differences of spring AO-related cyclonic circulation over the tropical North Pacific between the two periods are related to the different strengths of synoptic-scale eddy feedback, weak in the earlier period and strong in the later period. Indeed, we found that the springtime Pacific storm-track activity underwent a significant interdecadal change around the early 1970s from a weak regime to a strong regime. Therefore, it is concluded that the interdecadal change in the spring AO-related cyclonic circulation anomalies in the simultaneous spring over the tropical western North Pacific may be attributable to the interdecadal change in the intensity of the boreal spring North Pacific synoptic eddy activity.

Lee et al. (2012) showed that intensification of the North Pacific storm-track activity during boreal cold season after the middle to early 1970s is mainly attributed to the enhancement of mean meridional temperature gradient. It has been demonstrated by previous studies that enhancement of mean meridional temperature gradient could provide a favorable condition for baroclinic eddy growth (e.g., Lindzen and Farrell 1980; Hoskins and Valdes 1990; Nakamura and Shimpo 2004; Penny et al. 2010). Figure 12 shows difference in mean temperature between 1975–97 and 1949–71 at 1000, 925, 850, and 700 hPa. From Fig. 12, mean temperature at low levels significantly increased to the south and decreased to the north of about 30°N in the North Pacific. Hence, the meridional gradient of the mean temperature is enhanced in the North Pacific. This indicates that intensification of the North Pacific storm-track activity during boreal spring may be associated with the enhanced meridional thermal gradient.

Fig. 12.
Fig. 12.

Difference of mean temperature (°C) between 1975–97 and 1949–71 at (a) 1000, (b) 925, (c) 850, and (d) 700 hPa. Stippled areas denote the regions where the difference is significant at the 10% level.

Citation: Journal of Climate 28, 4; 10.1175/JCLI-D-14-00409.1

It should be noted that mean temperature anomaly pattern over North Pacific region in Fig. 12 bears some resemblances to that in association with the PDO (Mantua et al. 1997). Would the PDO contribute to the interdecadal change in the spring AO–EASM connection around the early 1970s? To address this issue, we have examined the PDO index averaged from March to May (not shown). A change in the PDO index can be found at the middle to late 1970s. It implies that change in the PDO index occurred at a time different from the change in the spring AO–EASM correlation. This indicates that PDO does not have a significant modulation effect on the connection between the spring AO and EASM.

Acknowledgments

We thank the three anonymous reviewers for their constructive suggestions and comments, which helped to improve the paper. This study is supported jointly by the National Natural Science Foundation of China Grants 41025017 and 41230527 and the Jiangsu Collaborative Innovation Center for Climate Change.

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