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Article

Comparison of the Impacts of Sea Surface Temperature in the Western Pacific and Indian Ocean on the Asian Summer Monsoon Anticyclone and Water Vapor in the Upper Troposphere

by
Luyao Chao
,
Hongying Tian
*,
Xiaoxu Tu
,
Jiaying Jiang
and
Kailong Shen
Key Laboratory for Semi-Arid Climate Change of the Ministry of Education, School of Atmospheric Sciences, Lanzhou University, Lanzhou 730000, China
*
Author to whom correspondence should be addressed.
Remote Sens. 2024, 16(16), 2922; https://doi.org/10.3390/rs16162922
Submission received: 26 June 2024 / Revised: 2 August 2024 / Accepted: 7 August 2024 / Published: 9 August 2024
(This article belongs to the Special Issue Remote Sensing Approaches for the Detection and Analysis in Coastal and River Delta Regions)
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Abstract

:
The variation in the Asian summer monsoon anticyclone (ASMA) has long been of interest due to its effects on the weather and climate, as well as the vertical transport of pollutants in South Asia and East Asia. This study employs composite analysis to investigate the differences in the influences of sea surface temperature (SST) anomalies in the Western Pacific (WP) and the Indian Ocean (IO) on the ASMA and water vapor in the upper troposphere during summer. The underlying physical mechanisms were further explored. The results indicate that the warm SSTs in the WP have a greater impact on the intensity of the ASMA than those in the IO in summer. On the contrary, the cold SSTs in the IO have a greater impact on intensity of the ASMA than those in the WP in summer. The difference in the impact of SSTs in the WP and IO on the boundaries of the ASMA is relatively small. During positive SST anomalies in the WP, the increase in tropospheric temperature in South Asia and the strengthening of Walker circulation in the WP both contribute to the enhancement of the ASMA. The variations in tropospheric temperature and Walker circulation caused by positive SST anomalies in the IO are similar to those in the WP, except that the rising branch of the Walker circulation is located in the central and western IO. The decrease in SST in the WP region causes insignificant changes in the ASMA. During the cold SST period in the IO, the significant decrease in tropospheric temperature and the weakening of the Walker circulation in the IO region lead to a significant decrease in the intensity of the ASMA at the southern ASMA. When the SST in the WP and IO regions is warmer, the high value centers of water vapor in the troposphere generally coincide with the high value centers of temperature, accompanied by enhanced convection, significantly increasing the water vapor south of the ASMA. The anomalous sinking movement in the Western Pacific leads to relatively small changes in water vapor from the near-surface to 150 hPa over the southeast of the ASMA.

1. Introduction

Due to the sustained deep convective activity in South Asia and the dynamic and thermal effects generated by the topography of the Tibetan Plateau, the Asian summer monsoon anticyclone (ASMA) is the most prominent circulation that appears in the upper troposphere–lower stratosphere (UTLS) region of the Northern Hemisphere during summer [1,2,3,4,5,6]. As one of the key components of the Asian summer monsoon system, the ASMA has a significant influence on the atmospheric circulations, weather, and climate in the East Asian and the Northern Hemisphere [7,8,9,10]. The movement and area of the ASMA can influence the summer extreme precipitation over eastern China [11]. It is evident that the ASMA holds significant indicative value for precipitation forecasting. Huang et al. [12] constructed seasonal prediction of Afro-Asian summer monsoon precipitation (AfroASMP) based on machine learning and large set methods, and found that the model could cleverly predict the four monsoon precipitation indices of AfroASMP in 2011–2022, with correlation coefficients ranging from 0.58 to 0.90. On the other hand, the existence of the ASMA can lead to unique and regional characteristics in the distribution and changes in atmospheric composition in the UTLS region during the Asian summer monsoon. The ASMA has a strong trapping effect on the air inside, causing the air transported from deep convection to the UTLS region to remain in the anticyclone for a period of time, resulting in an increase in the tropospheric composition (such as H2O, CO and HCN) and a decrease in the stratospheric composition (such as O3) inside the anticyclone [13]. Subsequently, the air inside the ASMA can be further lifted to the lower stratosphere and transported by horizontal airflow throughout the stratosphere [14,15,16].
Many scholars studied factors that influence the ASMA in the past. Factors such as the thermal effect of the Tibetan Plateau [4,17,18,19,20], the latent heat from precipitation [21,22] and the Pacific and Indian Ocean SST [23] have important impacts on the formation of the ASMA and its intensity as well as its meridional and latitudinal movement. The ASMA is located over the Tibetan Plateau and the Iranian Plateau in summer, with the Indian Ocean to the south and the Pacific Ocean to the east. It can be argued that the ASMA is an important linkage in the interaction between the sea, land, and atmosphere and has a direct interaction with the tropical and subtropical atmospheric circulation and SST [23,24,25,26,27,28,29,30,31,32]. Huang et al. [23] pointed out that during the summer in the northern hemisphere, an increase in Indian Ocean SST can enhance the strength of the ASMA and cause its center to shift southward at 100 hPa, which is mainly related to the tropospheric warming caused by the increase in SST. Yang et al. [26] indicated that the increase in Indian Ocean SST can force the upper tropospheric temperature to exhibit Rossby wave patterns, where Kelvin and Rossby waves manifest as a Matsuno-Gill pattern [33,34], leading to the strengthening of the ASMA. Yang and Li [25] studied the impact of the tropical Pacific-Indian Ocean temperature anomaly mode on the ASMA, and found that the ASMA becomes stronger and shifts northwestward when there are negative SST anomalies in the western Indian Ocean and eastern Pacific Ocean, and positive SST anomalies in the eastern Indian Ocean and Western Pacific Ocean. Peng et al. [35] used numerical simulations to confirm that the interdecadal variations of the ASMA are closely related to the SST in the Indian Ocean in winter, while in summer, the SST in the South China Sea and Western Pacific plays an important role. It can be seen that there are regional differences in the impact of sea surface temperatures in Indian Ocean and Pacific Ocean on the ASMA, but there is still controversy over which region’s sea surface temperature dominates. It is necessary to study and compare the influence of sea surface temperature in the Western Pacific and Indian Ocean on the ASMA and identify the main controlling factors.
Previous studies elucidated that deep convective activities in South Asia and the ASMA are crucial for transporting air masses from the troposphere to the lower stratosphere [7,8,9,13,36,37,38,39,40,41,42]. The UTLS region influenced by the ASMA exhibits anomalously high water vapor [9,13]. Gettelman et al. [7] indicated that 75% of the total water vapor transported to the global tropical stratosphere occurs in the Asian monsoon and the Tibetan Plateau during summer. The water vapor entering the stratosphere can affect the rate of global warming through radiative processes [43], and the water vapor can also influence the activation of chlorine, leading to a decrease in stratospheric ozone and further causing radiation changes [44]. The vertical transport of water vapor in the UTLS region of the ASMA attracted widespread attention from scholars. Su et al. [45,46] highlighted that deep convective activity induced by elevated tropical SST plays a key role in increasing water vapor in the upper troposphere. Xie et al. [47] noted that the tropical Rossby–Kelvin wave response over the Northern Indian Ocean and Western Pacific leads to sustained warming of the tropical Atlantic SST, thereby raising the tropopause temperature and significantly increasing stratospheric water vapor. As the dominant mode of tropical interannual climate variability, ENSO also influences UTLS water vapor by affecting the tropopause or upper tropospheric temperature through tropical upwelling [48,49,50]. Overall, tropical SST can affect the transport of water vapor to the UTLS region, but the impact of SST in the Western Pacific and Indian Ocean on water vapor in the upper troposphere within the ASMA is still unclear. The aim of this study is to compare the impact of SST changes in the Western Pacific and Indian Oceans on the intensity and range of the ASMA, as well as water vapor in the upper troposphere. Then, the potential physical mechanisms behind these changes are explored.

2. Materials and Methods

2.1. Materials

In this study, we utilize the ERA5 monthly mean reanalysis data set, which is made available by the European Centre for Medium-Range Weather Forecasts (ECMWF). ERA5 includes horizontal winds (u and v, units: m/s), geopotential (units: m2/s2), temperature (units: K), and vertical velocity (units: Pa/s). The ERA5 dataset’s horizontal resolution is 1° × 1°, with a vertical profile comprising 37 layers that range from 1000 hPa to 1 hPa. Outward longwave radiation (OLR, unit: W/s2) data on a monthly mean basis, with a resolution of 2.5° × 2.5°, are sourced from the National Centers for Environmental Prediction (NCEP) reanalysis data set of the National Oceanic and Atmospheric Administration (NOAA). The Extended Reconstructed Sea Surface Temperatures, unit: °C (ERSST), with 1° × 1° resolution, is extracted from the Hadley Centre Sea Ice and Sea Surface Temperature dataset (HadISST), an observational dataset created by the UK Met Office Hadley Center. The Microwave Limb Sounder (MLS) on the Aura satellite provides Level 2 water vapor data from 2005 to 2021 with a horizontal resolution along the orbit of 170–350 km below 4.6 hPa and a vertical resolution of 1.3–3.6 km from 316 to 0.22 hPa, spanning a total of 27 layers. Based on the data quality document, the water vapor data used ranges from 316 to 0.002 hPa, with an estimated precision in positive values, status code of even numbers, and quality greater than 1.45. The interpolated resolution of this data is 2° × 2°.
Earlier work demonstrated that there is an interdecadal variation in the relationship between the ASMA and the SSTs of the Indian Ocean, i.e., the correlation between the two was not high before 1979, whereas it became significant from 1979 onwards [44]. Comparable variations also occur in the connection of the ASMA with SST in adjacent areas, such as the Northwest Pacific Ocean [51,52]. Since the Indian Ocean and the Pacific Ocean are clearly connected and closely linked by Walker circulation [53], it can be considered that the Western Pacific and the ASMA have a comparable association. In addition, the intensity and extent of the ASMA also have significantly interdecadal variations. The intensity of the ASMA has been increasing since the late 1970s, accompanied by an expansion in its east–west span [35]. According to the results mentioned above, the period selected for the present study is from June to August (summer) over 1979–2021.

2.2. Methods

2.2.1. ASMA Index

Following the approaches proposed by Zhang et al. [24] and Zhou et al. [54], this study defines the intensity–area index of the ASMA as the total of the differences where the 100 hPa geopotential height is higher than 16,600 gpm, compared against 16,600 gpm, across all grid points within the region (0°~50°N, 0°~180°E). This index can reflect the comprehensive information of the intensity and size of the ASMA with obvious decadal and interdecadal variations.
As outlined by Zhang et al. [55], the ASMA intensity index is calculated using the average 150 hPa geopotential height during the summer over the area bounded by (20°~35°N, 65°~95°E).
To estimate the ASMA area, we referred to the definition of Zhu et al. [56]. The ASMA’s area index is determined by calculating the total area from the number of grid points with 100 hPa geopotential heights exceeding 16,800 gpm, each multiplied by the area of an individual grid cell.

2.2.2. Walker Circulation

According to Ge et al. [57], the Walker circulation is defined as the average vertical velocity over the range of 5°S to 5°N.

2.2.3. Composite Analysis

Composite analysis is a method frequently used in meteorological research. In this study, composite analysis is employed to examine years with high and low SST in the selected WP and IO regions. The regions exhibit strong positive correlation coefficients between SST and three characteristic indices of the ASMA. Differences between high and low SST years in the WP and IO regions are calculated, followed by significance testing. Significant areas identified through this testing are designated as key regions for further analysis.

2.2.4. Others

It is worth noting that in the present study, all variables had the least squares linear trend subtracted. The Student’s t-test is conducted to validate the significance of differences in the mean values of the samples.

3. Results

3.1. Impact of SST Anomalies in the Western Pacific and the Indian Ocean on the ASMA

Figure 1 shows the horizontal distribution of correlation coefficients between the ASMA intensity–area index, intensity index, and area index with SST during the summers from 1979 to 2021. Linear trends and seasonal variations were removed from both the ASMA index and SST. The seasonal variation of each individual variable is removed by subtracting its climatological value in the corresponding month. Significant positive correlations between the ASMA intensity–area index, intensity index, and area index with Indian Ocean SST are evident over the region spanning from 40°E to 100°E and 25°S to 25°N. In the region spanning from 110°E to 150°E and 0° to 35°N, there is a significant positive correlation between the ASMA intensity–area index, intensity index, and area index with Western Pacific SST. The two regions showing positive correlation are adjacent to the ASMA at 100 hPa and situated in the south and southeast of the ASMA, respectively. Since the air–sea interaction is more significant in the same or adjacent regions, heat exchanges between the oceans and the air also have a significant impact on the ASMA [24,35]. Therefore, Area 1 (depicted as the black box on the right side of Figure 1, 110°–150°E, 0°–35°N), which is denoted as the Western Pacific (WP) region, is selected as the region with the most obvious positive relationship between SST in the Western Pacific Ocean and the ASMA. Area 2 (depicted as the black box on the left side of Figure 1, 40°–100°E, 25°S–25°N), namely the Indian Ocean (IO) region, is chosen as the region with the most pronounced positive relationship between SST in the Indian Ocean and the ASMA. The normalized time series of SST averaged over the WP and IO regions for the summers from 1979 to 2021 are displayed in Figure 2. Abnormally warm and cold SST years are identified based on normalized SSTs greater than 0.5 and less than −0.5. The WP region is found to have 11 warm SST years (1983, 1987, 1988, 1991, 1995, 1998, 2001, 2010, 2016, 2017, and 2020) and 14 cold SST years (1982 1985, 1986, 1992, 1993, 1997, 2006, 2008, 2009, 2011, 2012, 2014, 2015, and 2018). The IO region has 12 warm SST years (1983, 1987, 1988, 1991, 1998, 2003, 2009, 2010, 2015, 2017, 2019, and 2020) and 14 cold SST years (1981, 1984, 1985, 1989, 1993, 1994, 1996, 1999, 2000, 2004, 2008, 2013, 2018, and 2021). Composite analyses are conducted based on these identified warm and cold SST years to comparatively study the impacts of SST anomalies in the Western Pacific and the Indian Ocean on the ASMA. In this study, the warm and cold SST years identified for the WP and IO regions partially overlap, which means that their impacts on the ASMA can be interwoven.
The ASMA is strongest at 150 hPa. To assess the impact of regional SST variations on both the intensity and boundaries of the ASMA, this study selects the pressure levels of 200 hPa and 100 hPa, which are adjacent to 150 hPa. Figure 3a–d displays the horizontal distributions of geopotential height anomalies and climatological mean at 200 hPa and 100 hPa during the warm and cold SST in the WP region. It is clear that the intensity of the ASMA at 200 hPa and 100 hPa anomalously enhanced over the entire ASMA area as well as to its south during the warm SST in the WP region (Figure 3a,c). During the cold SST years in the WP region, changes in geopotential height at both 100 hPa and 200 hPa are insignificant, while geopotential height only significantly decreases in the southwestern part of the ASMA. In addition, the east and west boundaries of the ASMA are also different at the above two pressure levels, but the north and south boundaries do not change significantly. At 200 hPa (Figure 3a), when the SST in the WP region is abnormally warm, the east and west boundaries of the ASMA significantly expand outward compared to the climatological anticyclone boundary, and the ASMA covers a large area from northeastern Africa in the west to the eastern coast of China in the east. It is important to note that the eastern boundary expanded much more extensively than the western boundary. When the SST in the WP region is cold, the east and west boundaries of the ASMA retreat inward, and the ASMA covers a smaller area from the northwest of the Arabian Peninsula in the central-west part of China in the east. In this situation, the extent of the west boundary retreat is larger than that of the east boundary. Corresponding to abnormally warm or cold SST in the WP region, variations in the east and west boundaries of the anticyclone at 100 hPa resemble to those at 200 hPa, although the changes at 100 hPa are relatively small (Figure 3c,d). Overall, the north boundary of the anticyclone at 100 hPa shifts northward by about 5° compared to that at 200 hPa. Meanwhile, the anticyclone becomes wider in morphology and controls a larger area, which is consistent with the characteristics of the warm center of the ASMA. When the SST in the WP region is abnormally warm, the east and west boundaries of the ASMA expand eastward and westward, respectively, at both 200 and 100 hPa. Compared with the climatological mean state, geopotential height anomalously enhances throughout the entire area of the ASMA as well as to its south, while the changes of the south and north boundaries of the anticyclone are relatively small. To some extent, these phenomena indicate that SST increase in the WP region significantly influences the intensity of the ASMA and its east and west boundaries, while exerting less impact on the southern and northern boundaries of the anticyclone.
Figure 3e–h illustrates horizontal distributions of differences in geopotential height between the climatological mean and the abnormally warm and cold SST years in the IO region at 200 hPa and 100 hPa, respectively. Comparing with Figure 3a and Figure 3c, it is found that the changes in the geopotential height within the ASMA range remain consistent with that in the WP region during warm SST years in the IO region (Figure 3e,g). However, the magnitude of the geopotential height changes within the ASMA under warm SST conditions in the IO region is comparatively lower than that in the WP region. This suggests that SST warming in the WP region has a more pronounced impact on the ASMA than in the IO region. On the contrary, during cold SST periods in the IO region, a notable decrease in the geopotential height is observed at the southern boundary of the ASMA and the area to its south (Figure 3f,h). This suggests that cold SST in the IO region has a more pronounced impact on the ASMA compared to cold SST in the WP region, which only induces relatively insignificant changes in the ASMA (Figure 3b,d). When SST varies in the IO region, changes in the boundary of the ASMA at 200 hPa and 100 hPa are basically consistent with those in the WP region. However, during cold SST periods in the IO region, the retreat of the ASMA’s west boundary is relatively weak, while the retreat of the ASMA’s east boundary is relatively strong. In summary, warm SSTs in the WP region may exert a stronger impact on the ASMA than warm SSTs in the IO region, whereas the opposite may be true for cold SSTs in the two regions.

3.2. Physical Mechanisms of SST Anomalies in the Western Pacific and Indian Oceans Affecting the ASMA

Prior research indicated that changes in the ASMA are associated with changes in tropospheric temperature [23]. In this study, we adopt the definition of Huang et al. [23], which defines the mean value of 850–100 hPa temperature as the tropospheric temperature. Figure 4 shows the horizontal distribution of tropospheric temperature differences during the periods of warm and cold SSTs in the WP region and the IO region with respect to the climatological state. During the warm SST years in the WP region (Figure 4a), the tropospheric temperature within the ASMA and over areas to its south is significantly warmer than normal, particularly in the western part of the ASMA and the central-west Indian Ocean. This anomalous warming subsequently results in a notable increase in local geopotential height and intensifies the ASMA (Figure 3a). However, the tropospheric warming in the eastern part of the ASMA is not significant, which differs somewhat from the variation in geopotential height at 200 hPa (Figure 3a). During the years of cold SST in the WP region (Figure 4b), changes in the tropospheric temperature are insignificant. It only decreases significantly on the southwest side of the ASMA, which is consistent with the change in the 200 hPa geopotential height (Figure 3b). In contrast, when SST in the IO region is warm (Figure 4c), the tropospheric temperature also increases significantly within the ASMA and over areas to its south, with relatively high values occurring in the west-central Indian Ocean. The increases in the tropospheric temperature in the eastern part of the ASMA are more significant compared with those in the WP region, which aligns with the variation in the geopotential height (Figure 3e). The anomalous tropospheric temperature increase causes an anomalous increase in local geopotential height, thereby enhancing the ASMA. During cold SST years in the IO region (Figure 4d), the tropospheric temperature decreases significantly at the southern boundary of the ASMA and over areas to its south. Meanwhile, there is a relatively small change in the west-central Indian Ocean, which is more pronounced compared to the decrease in tropospheric temperature in the WP region.
Figure 5a–d illustrates the latitude pressure and longitude pressure profiles of wind differences and air temperature differences between warm and cold SST years and climatological averages in the WP region. During the positive SST years in the Western Pacific, there is an abnormal rise in the entire tropospheric temperature. Anomalous downdrafts prevail in the eastern and central portions of the WP region below 150 hPa, accompanied by anomalous southwesterly flow in the lower troposphere. In addition, anomalous updrafts are dominant in the southern and western sectors of the WP region, coupled with anomalous northeasterly flow that prevails mainly from 250 to 150 hPa. Consequently, the positive temperature anomaly extends northward and westward, entering the ASMA area and reaching up to 100 hPa. This extension brings about a notable positive anomaly in local geopotential height, leading to the ASMA’s intensification (Figure 3a,b). When SST in the WP region is cold, a significant decrease in lower tropospheric temperature is observed. Since the lower troposphere is mainly influenced by abnormally western downdrafts, temperature changes above 500 hPa are not significant (Figure 5d). Figure 5e–h displays latitude pressure and longitude pressure cross sections of temperature and wind changes in warm and cold SST years compared to their climatological values in the IO region. Warm SST in the IO region (Figure 5e,f) induces an anomalous temperature increase throughout the troposphere, along with anomalous updrafts over the IO region and prevailing northeasterly flow from 250 to 100 hPa. These factors lead to a northward extension of the positive temperature anomaly, which enters the ASMA area, and results in an anomalous temperature increase in the ASMA. On the contrary, a decrease in SST in the IO region results in a significant temperature decrease over the region, which can extend to the center of the ASMA and reach up to 100 hPa. In summary, increasing SSTs in the WP and IO regions can cause an increase in tropospheric temperature, extending into the ASMA. The circulation also changes accordingly, resulting in an abnormal enhancement of the ASMA. During the cold SST years in the WP region, there is an insignificant cooling observed in the upper troposphere due to relatively small atmospheric circulation changes. However, decreasing SST in the IO region results in tropospheric temperature cooling that extends into the ASMA, leading to anomalous ASMA weakening.
SST variations influence atmospheric circulations not only by altering tropospheric temperature but also by affecting meridional Walker circulation, zonal Hadley circulation, and other circulation patterns through sensible and latent heat transport. Figure 6 shows the vertical profiles of Walker circulation anomalies and wind field anomalies over the WP and IO regions in the warm and cold SST years. Under warm SST conditions in the WP region (Figure 6a), the Walker circulation notably intensifies, with anomalous westward ascending motion in the central and western WP region and divergence at 300 hPa. Meanwhile, anomalous descending motions dominate in the eastern sector. Concurrently, OLR weakens significantly over the Indian Peninsula and the eastern Indian Ocean (Figure 7a), indicating enhanced convective activities. In contrast, OLR increases notably over the Western Pacific, suggesting weakened convective activities there. Strengthened convective activities can lead to greater precipitation, which in turn amplifies the release of latent heat through condensation. Earlier research demonstrated the critical influence of Indian Ocean precipitation on ASMA variability [58]. The significant surge in latent heat from condensation over the eastern Indian Ocean raises the geopotential height at 200 hPa to an unusually high level. The elevated geopotential height anomalies then expand northward and reach the area of the ASMA, leading to abnormal intensification of the ASMA (Figure 3a). During periods of positive SST in the Indian Ocean (Figure 6c), the Walker circulation also strengthens. However, this strengthening is not as pronounced as that during warm SST periods in the WP region. Over the western Indian Ocean, anomalous westward ascending motion prevails, while anomalous descending motion dominates over the Western Pacific. Easterly wind anomalies dominate above 150 hPa over the Indian Ocean, while westerly wind anomalies prevail above 150 hPa outside the Indian Ocean. Meanwhile, OLR anomalies weaken over the area from the western Indian Ocean to the vicinity of the Indian Peninsula, corresponding to enhanced convective anomalies. OLR significantly increases over the Western Pacific, indicating a substantial decrease in convection there. Similarly, the notable strengthening of convection over the central-western Indian Ocean leads to an abnormally high geopotential height at 200 hPa over the region (Figure 3e). Compared to the impact of warm SSTs in the WP region, the anomalous Walker circulation induced by warm SSTs in the IO region tends to exhibit a more westward orientation, leading to a more westward distribution of very high geopotential height anomalies. During cold SST periods in the WP region (Figure 6b), the Walker circulation anomalies are generally opposite to those during warm SST periods. The anomalies are characterized by anomalies of convergence at 150–200 hPa over the WP region and its western area, corresponding to anomalous eastward downdrafts. Anomalous updrafts prevail east of the WP region. This pattern aligns with the significantly weakened convection over the central and eastern Indian Ocean and the significantly enhanced convection over the Western Pacific (Figure 7b). During cold SST periods in the IO region (Figure 6d), the Walker circulation anomalies are opposite to those during warm SST periods. Anomalous downdrafts prevail over the central and western Indian Ocean, resulting in reduced precipitation over the Indian Ocean and subsequently a significant decrease in geopotential height, thereby weakening the ASMA. Near the Indian Peninsula and the eastern Indian Ocean, convective activity intensifies simultaneously (Figure 7d). This change is opposite to the changes observed during cold SST periods in the WP region. This opposite change partially offsets the weakening effect of convective activity caused by cold SSTs in the WP region and explains, to a certain degree, why there is no significant decrease in geopotential height during cold SST periods in the WP region.
Following the Gill pattern [34], anomalous cyclonic circulation will occur in the lower troposphere during warm SST periods in the WP region. Figure 8a–d shows the 850 hPa geopotential height anomalies in summer for warm and cold SST conditions in the WP and IO regions, compared to the average climatological state. Positive geopotential height anomalies and anomalous anticyclonic circulation in the lower troposphere are evident over the Western Pacific Ocean during warm SST periods in both the WP and IO regions. Additionally, during the period of warm SST in the IO region, anomalous geopotential height over the Western Pacific exhibits a broader enhancement. Previous studies indicated that warm SST in the Indian Ocean can generate anomalous anticyclonic circulation over the Western Pacific Ocean [59]. Given the overlap of positive and negative SST years between the WP and IO regions, repetitive years of the positive and negative SST conditions in the WP and IO regions are excluded before plotting Figure 8e–h. It is observed that during warm SST periods in the WP region, the abnormal anticyclonic circulation over the Western Pacific Ocean notably weakens and fails to pass the significance test. Therefore, one can conclude that the anomalous anticyclonic circulation over the Western Pacific Ocean, generated by elevated SST in the IO region, counterbalances the unusual upwelling in the Western Pacific driven by warm SST in the WP region.

3.3. Impact of SST Anomalies in the Western Pacific and the Indian Ocean on the Water Vapor in the Upper Troposphere

This study utilizes Level 2 water vapor data from the MLS on the Aura satellite, while other variables are derived from ERA5 reanalysis data. To avoid the decrease in data reliability caused by interpolation, the study uses MLS water vapor data at 216 hPa and corresponding ERA5 reanalysis variables at 200 hPa. The following content demonstrates the response of upper troposphere water vapor content within the ASMA to SST changes in the WP and IO regions. Figure 9 depicts the horizontal distribution of MLS water vapor differences at 216 hPa and ERA5 water vapor differences at 200 hPa between warm SST and climate state in the WP and IO regions. Since MLS data are only available after 2004, the significance of the water vapor differences is weaker than that of ERA5, but the areas of increased water vapor are substantially consistent between the two datasets. When the SST in the WP region is warm, there is a significant increase in water vapor at 200 hPa on the north side of the Tibetan Plateau, eastern part within the ASMA, and south of the ASMA (Figure 9a,c). Compared to the locations of abnormal increases in geopotential height, there is no significant increase in water vapor in the western part within the ASMA and southeastern side of the ASMA. High values of water vapor appear over the eastern part within the ASMA and the Indian Peninsula. When the SST in the IO region is warm, the water vapor changes at 200 hPa are generally consistent with those when the SST in the WP region is warm (Figure 9d). Within the ASMA, the range of significant increase in water vapor is smaller than when the SST in the WP region is warmer. In the eastern part of the ASMA, the range of significant increase in water vapor is greater than when the SST in the WP region is warmer. High values of water vapor occur in the eastern part within the ASMA, the Arabian Sea, and the Bay of Bengal.
Based on Figure 9, we selected the relative center position of the ASMA (30°N, 80°E) and the region with high water vapor values at 15°N to further observe the vertical distribution of water vapor. Figure 10 shows the vertical profiles of water vapor differences along 80°E, 30°N, and 15°N between warm SST and climate state in the WP and IO regions. The water vapor changes for the cooler years are not shown here because the water vapor changes in cooler years are essentially the opposite of those in warmer years. When the SST in the WP region is warmer, water vapor significantly increases on the southern side of the ASMA, which can reach up to about 125 hPa. There is a relatively high value of water vapor at 400–150 hPa. These changes correspond to the significant increase in water vapor at 200 hPa on the southern side of the ASMA, as shown in Figure 9c. In addition, most of this water vapor does not enter the stratosphere (Figure 10a). In the northern part of the Tibetan Plateau, there is also a relative high value of water vapor. Due to the lower tropopause in this region, the water vapor can extend upward and northward into the stratosphere (Figure 10a). Directly above and to the east of the Tibetan Plateau, the water vapor extends upward into the eastern part within the ASMA, reaching up to 125 hPa (Figure 10b). From Figure 10c, it can be seen that water vapor in the troposphere increases significantly, with high values of water vapor appearing between 600 and 200 hPa over the eastern Arabian Sea. The anomalous increase in water vapor extends upward and eastward, bypassing the region of water vapor decrease from the near-surface to 150 hPa over 110°–140°E. Then the water vapor enters the stratosphere, leading to a significant increase in water vapor in the stratosphere over the Western Pacific. When the SST in the IO region is warmer, the water vapor changes are generally consistent with those in the WP region (Figure 10d–f). The water vapor significantly increases on the southern side of the ASMA, with a high value of water vapor around 500–350 hPa and another high value around 350–150 hPa (Figure 10d). Compared to the case of warm SST in the WP region, the high values of water vapor appear with two centers here (Figure 10a). In the eastern part of the Tibetan Plateau, the water vapor extends upward into the eastern part within the ASMA, reaching the stratosphere (Figure 10e). In Figure 10f, it can be seen that water vapor over the Arabian Sea to the Bay of Bengal also extends upward. Most of the water vapor enters the stratosphere, leading to a significant increase in water vapor in the stratosphere over 10°–140°E.

3.4. Physical Mechanisms of SST Anomalies in the Western Pacific and Indian Oceans Affecting the Water Vapor in the Upper Troposphere

Based on Figure 10, we plotted the vertical profiles of temperature and wind fields (Figure 11). Additionally, we plotted the horizontal distribution of divergence differences at 500 hPa and 200 hPa to further study the changes in water vapor (Figure 12). When the SST in the WP region is warmer, the high value regions of water vapor in the troposphere generally match with the high value regions of temperature. They are accompanied by anomalous upward airflow towards the east, allowing the water vapor to reach around 125 hPa (Figure 10a,b and Figure 11a,b). This indicates that the warm centers in the troposphere are conducive to the increase in water vapor. There is anomalous low-level convergence and high-level divergence over the eastern Arabian Sea (Figure 12a,b), with a decrease in OLR anomalies (Figure 7a). The region is mainly controlled by anomalous upward airflow (Figure 11c), resulting in the upward transport of water vapor and the appearance of high water vapor values over the eastern Arabian Sea (Figure 9c). Meanwhile, there is an increase in OLR anomalies over the Western Pacific (Figure 7a), with anomalous downward airflow dominating the troposphere (Figure 11c). This corresponds to the anomalous anticyclone circulation at 850 hPa and anomalous convergence in the southeast of the ASMA at 200 hPa over the Western Pacific (Figure 12a,b), which might be the reason for the lack of a significant increase in water vapor in the southeast of the ASMA. Figure 8 explains that this anomalous anticyclonic circulation is triggered over the Western Pacific by the warm sea surface temperature in the IO region. When the SST in the IO region is warmer, the changes in temperature and wind fields are generally consistent with those when the SST in the WP region is warmer (Figure 11d,e). Anomalous upward airflow also controls the region between 60° and 90°E (Figure 11f). The significant decrease in OLR over the Arabian Sea corresponds to a significant increase in convection (Figure 7c), resulting in high water vapor values at 200 hPa over the Arabian Sea (Figure 9d). However, the low-level convergence and high-level divergence over the eastern Arabian Sea are stronger compared to those when the SST in the WP region is warmer (Figure 12c,d), allowing more water vapor to enter the stratosphere (Figure 10f). In summary, when the SST in the WP and IO regions is warmer, the high-value centers of water vapor in the troposphere generally match with the high-value centers of temperature and are accompanied by anomalous increases in convection. These changes significantly increase the water vapor on the south side of the ASMA. The anomalous downward airflow over the Western Pacific results in insignificant changes in water vapor from the near-surface to 150 hPa.

4. Conclusions

This study utilized the ERA5 reanalysis dataset, OLR data, SST data from 1979 to 2021, as well as the water vapor data provided by MLS from 2004 to 2021. The composite analysis method was used to investigate the impact of SST anomalies in the WP and IO regions on the ASMA and the water vapor in the upper troposphere around the ASMA. The detailed physical mechanism is summarized schematically in Figure 13.
The results reveal that abnormal increases in geopotential height occur within and to the south of the ASMA under warm SST conditions in the WP, resulting in a significant strengthening of the ASMA and outward expansion of its east and west boundaries. During cold SST periods in the WP region, there is insignificant variation in the intensity of the ASMA at 200 and 100 hPa, although a notable decrease in geopotential height can be found at the southwest boundary of the ASMA. The eastern and western boundaries of the ASMA contract inward, although to a weak extent. When SST in the IO is warmer, significant increases in geopotential height are observed within the ASMA and over areas to its south at 200 and 100 hPa. The intensity and area variations of the ASMA are comparable to that during warm SST periods in the WP region. However, the enhancement in geopotential height within the ASMA is less significant compared to that under warm SST conditions in the WP region. Moreover, the area of anomalous high geopotential height shifts westward over the Indian Ocean. During cold SST periods of IO, different from the effect of cold SST in the WP region, there is a significant decrease in geopotential height in the south boundary of the ASMA and its south. The ASMA notably weakens with inward contraction of its east and west boundaries. Overall, the warm SSTs in the WP region may have a greater impact on the intensity of the ASMA than those in the IO region. On the contrary, the cold SSTs in the IO region may have a greater impact on the intensity of the ASMA than those in the WP region. The difference in the impact of SST in the WP and IO on the boundaries of the ASMA is relatively small.
The warming of SST in the WP region induces a notable warming of the lower troposphere in that area. This warming is accompanied by anomalous updrafts in the southwest of the WP region and extends westward and northward into the ASMA area. Consequently, the tropospheric temperature within and to the south of the ASMA experiences significant increases, aligning with the range of anomalous geopotential height increases. Meanwhile, the Walker circulation over the WP region has notably enhanced and convection in the eastern IO has anomalously increased. These led to an anomalous increase in latent heat release over this region, which coincides with relatively high geopotential height. The rising tropospheric temperature in South Asia and increased latent heat release over the IO collectively contribute to the increase in geopotential height within and to the south of the ASMA. As a result, the ASMA is markedly strengthened and its east and west boundaries expand outward.
When the SST in the IO region is warm, the atmosphere in the lower troposphere above the IO region is abnormally warming, and the positive temperature anomaly extends northward to the ASMA with the abnormally rising airflow. This results in a significant increase in tropospheric temperature within and to the south of the ASMA. Meanwhile, the Walker circulation in the IO region undergoes significant strengthening, but its degree of enhancements is weaker than that during the warm SST in the WP region. Anomalous updrafts prevail over the western Indian Ocean, while an anomalous anticyclonic circulation occurs in the lower troposphere over the Western Pacific Ocean, resulting in increased latent heat release from precipitation. The combined effects of the above factors result in a notable enhancement of geopotential height within and to the south of the ASMA. During the cold SST in the IO, the lower troposphere over the IO region significantly cools down. This cooling extends with the anomalous westward and upward motion over the eastern Indian Ocean into the ASMA region and ultimately reaches up to 100 hPa. Concurrently, the weakening of the Walker circulation in the IO region results in reduced convective activities over the western Indian Ocean. This results in a notable decline in tropospheric temperature at the south boundary of the ASMA and its south. It leads to a significant decrease in the intensity of the ASMA at its south boundary and to its south. This is significantly different from the results during cold SST periods in the WP region.
When the SST in the WP region is warm, an abnormal anticyclonic circulation appears over the Western Pacific Ocean. However, according to the Gill pattern, the Western Pacific as a heat source will trigger an anomalous cyclonic circulation. This is primarily attributed to the dominance of an anomalous anticyclonic system triggered by the rise in SST in the Indian Ocean. When the SST in the Western Pacific is cold, the changes in anticyclones are not significant. This is due to the significant enhancement of convective activity in the eastern Indian Ocean during the cold SST period in the IO region. At the same time, there is an overlap in cold SST years between the WP and IO regions. This reduces the degree of abnormal weakening of convective activity in the eastern Indian Ocean and leads to smaller changes in the intensity of the ASMA when WP SST decreases.
The water vapor at 200 hPa significantly increases within the east and south of the ASMA during the SST warming in the WP and IO regions. The high centers of water vapor in the troposphere generally coincide with the high-value centers of temperature, accompanied by anomalous enhanced convection, which significantly increases the water vapor south of the ASMA. The anomalous downward movement over the Western Pacific leads to insignificant changes in water vapor from the near-surface to 150 hPa. When the SST in the WP region is warmer, the anomalous upward movement within the eastern ASMA is located further west compared to the warmer SST in the IO, causing the center of high water vapor to move eastward. Compared to the warmer SST in the IO, the anomalous upward movement south of the ASMA is located further east, resulting in an eastward shift of the high water vapor center.

Author Contributions

Conceptualization, H.T.; data curation, L.C., X.T., J.J. and K.S.; formal analysis, L.C. and H.T.; investigation, L.C., X.T., J.J. and K.S.; methodology, H.T.; project administration, L.C., X.T., J.J. and K.S.; resources, H.T.; software, L.C. and H.T.; supervision, H.T.; validation, H.T.; visualization, L.C.; writing—original draft, L.C. and H.T.; writing—review and editing, H.T., L.C., X.T., J.J. and K.S. All authors have read and agreed to the published version of the manuscript.

Funding

This research was funded by the National Natural Science Foundation of China (No. 42130601, 42175088).

Data Availability Statement

The original contributions presented in the study are included in the article, further inquiries can be directed to the corresponding author.

Acknowledgments

The authors would like to thank the support for the ERA5 monthly mean reanalysis product provided by the European Center for Medium-Range Weather Forecasts, NCEP reanalysis product provided by the National Oceanic and Atmospheric Administration, and SST data extracted from the Hadley Centre Sea Ice and Sea Surface Temperature dataset.

Conflicts of Interest

The authors declare no conflicts of interest.

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Remotesensing 16 02922 g001
Figure 1. Horizontal distributions of correlation coefficients (color shaded) of the (a) ASMA intensity–area index, (b) intensity index, and (c) area index with SST in the summer (June–August) of 1979–2021. The red contours show the JJA climatological geopotential height enclosed by the 16,750 gpm isoline at 100 hPa. The left black box indicates the selected study region over the Indian Ocean (IO: 40°–100°E, 25°S–25°N), and the right black box indicates the selected study region over the Western Pacific (WP: 110°–150°E, 0°–35°N). The thin black line indicates the Tibetan Plateau boundary (the same hereafter), and the black dots represent the corresponding values are significant at the 95% confidence level.
Figure 1. Horizontal distributions of correlation coefficients (color shaded) of the (a) ASMA intensity–area index, (b) intensity index, and (c) area index with SST in the summer (June–August) of 1979–2021. The red contours show the JJA climatological geopotential height enclosed by the 16,750 gpm isoline at 100 hPa. The left black box indicates the selected study region over the Indian Ocean (IO: 40°–100°E, 25°S–25°N), and the right black box indicates the selected study region over the Western Pacific (WP: 110°–150°E, 0°–35°N). The thin black line indicates the Tibetan Plateau boundary (the same hereafter), and the black dots represent the corresponding values are significant at the 95% confidence level.
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Figure 2. Standardized time series of SST (unit: °C) averaged over the (a) WP region and the (b) IO region in the summers of 1979–2021. The red dots represent years with warm SST, the blue dots represent years with cold SST, and the black dots represent years with normal SST.
Figure 2. Standardized time series of SST (unit: °C) averaged over the (a) WP region and the (b) IO region in the summers of 1979–2021. The red dots represent years with warm SST, the blue dots represent years with cold SST, and the black dots represent years with normal SST.
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Figure 3. Horizontal distribution of anomalies in geopotential height (unit: gpm) between warm SST and climate state at (a,b,e,f) 200 hPa and (c,d,g,h) 100 hPa in the (a,c) WP region and (e,g) IO region during summer, as well as differences in geopotential height between cold SST and climate state in the (b,d) WP region and (f,h) IO region. Green, purple, and black contours in (a,b,e,f) indicate the 12,520 gpm geopotential height contours at 200 hPa for warm SST years, cold SST years, and the climatological state, respectively. Green, purple, and black contours in (c,d,g,h) indicate the 16,750 gpm geopotential height contours at 100 hPa for warm SST years, cold SST years, and the climatological state, respectively. Dotted regions represent that the anomalies are statistically significant at the 90% confidence level (the same hereafter).
Figure 3. Horizontal distribution of anomalies in geopotential height (unit: gpm) between warm SST and climate state at (a,b,e,f) 200 hPa and (c,d,g,h) 100 hPa in the (a,c) WP region and (e,g) IO region during summer, as well as differences in geopotential height between cold SST and climate state in the (b,d) WP region and (f,h) IO region. Green, purple, and black contours in (a,b,e,f) indicate the 12,520 gpm geopotential height contours at 200 hPa for warm SST years, cold SST years, and the climatological state, respectively. Green, purple, and black contours in (c,d,g,h) indicate the 16,750 gpm geopotential height contours at 100 hPa for warm SST years, cold SST years, and the climatological state, respectively. Dotted regions represent that the anomalies are statistically significant at the 90% confidence level (the same hereafter).
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Figure 4. Horizontal distribution of anomalies in tropospheric temperature (defined as the mean air temperature from 850 to 100 hPa, color shaded, unit: K) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region. Green, purple, and black contours indicate the 12,520 gpm geopotential height contour at 200 hPa for years of warm, cold, and climate mean SSTs, respectively.
Figure 4. Horizontal distribution of anomalies in tropospheric temperature (defined as the mean air temperature from 850 to 100 hPa, color shaded, unit: K) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region. Green, purple, and black contours indicate the 12,520 gpm geopotential height contour at 200 hPa for years of warm, cold, and climate mean SSTs, respectively.
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Figure 5. Vertical profiles of regional averaged temperature anomalies between warm SST and climate state in the (a,b) WP region (average along the selected WP region) and (e,f) IO region (average along the selected IO region), as well as differences between cold SST and climate state in the (c,d) WP region and (g,h) IO region. In panels (a,c,e,g), purple arrows represent meridional wind and vertical velocity passing the significance test, while green contours represent zonal wind passing the significance test. In panels (b,d,f,h), purple arrows represent zonal wind and vertical velocity passing significance test, while green contours represent meridional wind passing significance test. Positive values are displayed as solid green contours, and negative values as dashed green lines. Magenta contour lines indicate the ASMA, and black dashed lines denote the selected WP region and IO region.
Figure 5. Vertical profiles of regional averaged temperature anomalies between warm SST and climate state in the (a,b) WP region (average along the selected WP region) and (e,f) IO region (average along the selected IO region), as well as differences between cold SST and climate state in the (c,d) WP region and (g,h) IO region. In panels (a,c,e,g), purple arrows represent meridional wind and vertical velocity passing the significance test, while green contours represent zonal wind passing the significance test. In panels (b,d,f,h), purple arrows represent zonal wind and vertical velocity passing significance test, while green contours represent meridional wind passing significance test. Positive values are displayed as solid green contours, and negative values as dashed green lines. Magenta contour lines indicate the ASMA, and black dashed lines denote the selected WP region and IO region.
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Figure 6. Meridional-vertical profiles of Walker circulation (averaged between 5°S and 5°N) anomalies (color shaded) and winds field anomalies (purple arrows indicate latitudinal wind and vertical velocity) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region. Black dashed lines delineate the selected WP region in panel (a,b), and the selected IO region in panel (c,d).
Figure 6. Meridional-vertical profiles of Walker circulation (averaged between 5°S and 5°N) anomalies (color shaded) and winds field anomalies (purple arrows indicate latitudinal wind and vertical velocity) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region. Black dashed lines delineate the selected WP region in panel (a,b), and the selected IO region in panel (c,d).
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Figure 7. Horizontal distributions of outward longwave radiation anomalies (color shaded, unit: W/m2) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region.
Figure 7. Horizontal distributions of outward longwave radiation anomalies (color shaded, unit: W/m2) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region.
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Figure 8. Horizontal distribution of anomalies in geopotential height (color shaded, unit: gpm) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region. The (e–h) are similar to (a–d), but the repetitive years were removed from the warm and cold years of SST in the WP region and IO region. The green arrows indicate abnormal horizontal winds (u and v). The area shaded in black represents the Tibetan Plateau.
Figure 8. Horizontal distribution of anomalies in geopotential height (color shaded, unit: gpm) between warm SST and climate state in the (a) WP region and (c) IO region during summer, as well as differences between cold SST and climate state in the (b) WP region and (d) IO region. The (e–h) are similar to (a–d), but the repetitive years were removed from the warm and cold years of SST in the WP region and IO region. The green arrows indicate abnormal horizontal winds (u and v). The area shaded in black represents the Tibetan Plateau.
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Figure 9. Horizontal distribution of anomalies in MLS water vapor (unit: ppmv) at 216 hPa between warm SST and climate state in (a) WP region and (b) IO region during summer, as well as differences in ERA5 water vapor at 200 hPa in (c) WP region and (d) IO region. Green, blue, and black contours indicate the 12,520 gpm geopotential height contours at 200 hPa for warm SST years, cold SST years, and the climatological state, respectively. Dotted regions represent that the anomalies are statistically significant at the 90% confidence level.
Figure 9. Horizontal distribution of anomalies in MLS water vapor (unit: ppmv) at 216 hPa between warm SST and climate state in (a) WP region and (b) IO region during summer, as well as differences in ERA5 water vapor at 200 hPa in (c) WP region and (d) IO region. Green, blue, and black contours indicate the 12,520 gpm geopotential height contours at 200 hPa for warm SST years, cold SST years, and the climatological state, respectively. Dotted regions represent that the anomalies are statistically significant at the 90% confidence level.
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Figure 10. Vertical profiles of water vapor anomalies (color shaded, unit: 10−8%) along 80°E, 30°N, and 15°N between warm SST and climate state in (a–c) WP region and (d–f) IO region during summer. The green line represents the height of the tropopause (2.5 PVU contour line), the dotted area indicates passing the 90% significance test, and the black contour line represents the range of the ASMA. Black shadow represents topography.
Figure 10. Vertical profiles of water vapor anomalies (color shaded, unit: 10−8%) along 80°E, 30°N, and 15°N between warm SST and climate state in (a–c) WP region and (d–f) IO region during summer. The green line represents the height of the tropopause (2.5 PVU contour line), the dotted area indicates passing the 90% significance test, and the black contour line represents the range of the ASMA. Black shadow represents topography.
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Figure 11. Vertical profiles of temperature anomalies (color shaded, unit: K) along 80°E, 30°N, and 15°N between warm SST and climate state in the (a–c) WP region and (d–f) IO region during summer. Purple arrows represent the vertical velocity differences. The green line represents the height of the tropopause (2.5 PVU contour line), the dotted area indicates passing the 90% significance test, and the black contour line represents the range of the ASMA. Black shadow represents topography.
Figure 11. Vertical profiles of temperature anomalies (color shaded, unit: K) along 80°E, 30°N, and 15°N between warm SST and climate state in the (a–c) WP region and (d–f) IO region during summer. Purple arrows represent the vertical velocity differences. The green line represents the height of the tropopause (2.5 PVU contour line), the dotted area indicates passing the 90% significance test, and the black contour line represents the range of the ASMA. Black shadow represents topography.
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Figure 12. Horizontal distribution of anomalies in divergence (unit: 10−6 s−1) at 500 and 200 hPa between warm SST and climate state in (a,b) WP region and (c,d) IO region during summer. Purple arrows represent the horizontal winds (u and v, units: m/s). Green, blue, and black contours indicate the 12,520 gpm geopotential height contours at 200 hPa for warm SST years, cold SST years, and the climatological state, respectively.
Figure 12. Horizontal distribution of anomalies in divergence (unit: 10−6 s−1) at 500 and 200 hPa between warm SST and climate state in (a,b) WP region and (c,d) IO region during summer. Purple arrows represent the horizontal winds (u and v, units: m/s). Green, blue, and black contours indicate the 12,520 gpm geopotential height contours at 200 hPa for warm SST years, cold SST years, and the climatological state, respectively.
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Figure 13. Summary schematic.
Figure 13. Summary schematic.
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MDPI and ACS Style

Chao, L.; Tian, H.; Tu, X.; Jiang, J.; Shen, K. Comparison of the Impacts of Sea Surface Temperature in the Western Pacific and Indian Ocean on the Asian Summer Monsoon Anticyclone and Water Vapor in the Upper Troposphere. Remote Sens. 2024, 16, 2922. https://doi.org/10.3390/rs16162922

AMA Style

Chao L, Tian H, Tu X, Jiang J, Shen K. Comparison of the Impacts of Sea Surface Temperature in the Western Pacific and Indian Ocean on the Asian Summer Monsoon Anticyclone and Water Vapor in the Upper Troposphere. Remote Sensing. 2024; 16(16):2922. https://doi.org/10.3390/rs16162922

Chicago/Turabian Style

Chao, Luyao, Hongying Tian, Xiaoxu Tu, Jiaying Jiang, and Kailong Shen. 2024. "Comparison of the Impacts of Sea Surface Temperature in the Western Pacific and Indian Ocean on the Asian Summer Monsoon Anticyclone and Water Vapor in the Upper Troposphere" Remote Sensing 16, no. 16: 2922. https://doi.org/10.3390/rs16162922

APA Style

Chao, L., Tian, H., Tu, X., Jiang, J., & Shen, K. (2024). Comparison of the Impacts of Sea Surface Temperature in the Western Pacific and Indian Ocean on the Asian Summer Monsoon Anticyclone and Water Vapor in the Upper Troposphere. Remote Sensing, 16(16), 2922. https://doi.org/10.3390/rs16162922

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