Editing IPCC:AR6/WGI/Chapter-8 (section)

===== 8.3.2.4.1 South and South East Asian Monsoon =====

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The AR5 reported a decreasing trend of global land monsoon precipitation over the last half-century, with primary contributions from the weakened summer monsoon systems in the Northern Hemisphere (NH). Since AR5, several studies have documented long-term variations and changes in the South and South East Asian summer monsoon (SAsiaM) rainfall. The SAsiaM strengthened during past periods of enhanced summer insolation in the NH, such as the early-to-mid Holocene warm period around 9000 to 6000 years before the present (BP) ( [[#Masson-Delmotte--2013|Masson-Delmotte et al., 2013]] ; [[#Mohtadi--2016|Mohtadi et al., 2016]] ; [[#Braconnot--2019|Braconnot et al., 2019]] ) and weakened during cold periods ( ''high confidence'' ), such as the Last Glacial Maximum (LGM) and Younger Dryas (Shakun et al. , 2007; Cheng et al. , 2012; Dutt et al. , 2015; Chandana et al. , 2018; Hong et al. , 2018; E. Zhang et al. , 2018). These long-time scale changes in monsoon intensity are tightly linked to orbital forcing and changes in high-latitude climate (Braconnot et al. , 2008; Battisti et al. , 2014; Araya-Melo et al. , 2015; Rachmayani et al. , 2016; Bosmans et al. , 2018; E. Zhang et al. , 2018). A weakening trend of the SAsiaM during the last 200 years has been documented based on tree ring oxygen isotope chronology from the northern Indian subcontinent ( [[#Xu--2018|Xu et al., 2018]] ) and South East Asia ( [[#Xu--2013|Xu et al., 2013]] ), oxygen isotopes in speleothems from northern India ( [[#Sinha--2015|Sinha et al., 2015]] ), and tree ring width chronologies from the Indian core monsoon region ( [[#Shi--2017|Shi et al., 2017]] ). Nevertheless, the detection of century-long decreases in regional monsoon rainfall is obscured by the presence of multi-decadal time scale precipitation variations ( [[#Turner--2012|Turner and Annamalai, 2012]] ; [[#Knutson--2018|Knutson and Zeng, 2018]] ) which are evident in long-term rain guage records extending back to the early 1800s ( [[#Sontakke--2008|Sontakke et al., 2008]] ) and emerge in long-term climate simulations ( [[#Braconnot--2019|Braconnot et al., 2019]] ).

A significant decline in summer monsoon precipitation is observed over India since the mid-20th century, which is accompanied by a weakening of the large-scale monsoon circulation (Mishra et al. , 2012; Abish et al. , 2013; Krishnan et al. , 2013, 2016; Saha et al. , 2014; Roxy et al. , 2015; Guhathakurta et al. , 2017; Samanta et al. , 2020). This precipitation decline is corroborated by a decreasing trend in the frequency of monsoon depressions that form over Bay of Bengal ( [[#Prajeesh--2013|Prajeesh et al., 2013]] ; [[#Vishnu--2016|Vishnu et al., 2016]] ), an increasing trend in the frequency and duration of monsoon breaks or ‘dry spells’ ( [[#Singh--2014|Singh et al., 2014]] ), significant decreases in soil moisture and increases in drought severity across different parts of India post-1950 (Niranjan Kumar et al. , 2013; Ramarao et al. , 2015, 2019; Krishnan et al. , 2016; Ganeshi et al. , 2020; Mujumdar et al. , 2020). While recent studies have reported an apparent recovery of the Indian summer monsoon over a relatively short period since 2003 ( [[#Jin--2017|Jin and Wang, 2017]] ; [[#Hari--2020|Hari et al., 2020]] ), long-term trends for the period 1951 – 2015 indicate an overall decrease in the regional monsoon precipitation ( [[#Kulkarni--2020|Kulkarni et al., 2020]] ; [[#Ayantika--2021|Ayantika et al., 2021]] ). A case study on the Indian summer monsoon is provided in [[IPCC:Wg1:Chapter:Chapter-10#10.6.3|Section 10.6.3]] .

Evidence from several climate modelling studies indicates that the observed decrease in the regional monsoon precipitation during the second half of the 20th century is dominated by the radiative effects of NH anthropogenic aerosols, with smaller contributions due to volcanic aerosols from the Mount Pinatubo (1991) and El Chichón (1982) eruptions (Bollasina et al. , 2011; Polson et al. , 2014; Sanap et al. , 2015; Krishnan et al. , 2016; Liu et al. , 2016; [[#Lau--2017|Lau and Kim, 2017]] ; Lin et al. , 2018; Takahashi et al. , 2018; Undorf et al. , 2018a, b; Patil et al. , 2019; M. Singh et al. , 2020; see Box 8.1, Figure 1 and Figure 8.11). Land-use changes over South and South East Asia and the rapid warming trend of the equatorial Indian Ocean during the recent few decades also appear to have contributed to the observed decrease in monsoon precipitation ( [[#Roxy--2015|Roxy et al., 2015]] ; [[#Krishnan--2016|Krishnan et al., 2016]] ; [[#Singh--2016|Singh, 2016]] ). Overall, the magnitude of the precipitation response to anthropogenic forcing exhibits large spread across CMIP5 models pointing to the strong internal variability of the regional monsoon ( [[#Saha--2014|Saha et al., 2014]] ; [[#Salzmann--2014|Salzmann et al., 2014]] ; [[#Sinha--2015|Sinha et al., 2015]] ), including variations linked to phase changes of the Pacific Decadal Variability (Section AVI.2.6; X. [[#Huang--2020|Huang et al., 2020]] a), uncertainties in representing aerosol – cloud interactions ( [[#Takahashi--2018|Takahashi et al., 2018]] ), and the effects of local compared with remote aerosol forcing (Bollasina et al. , 2014; Polson et al. , 2014; Undorf et al. , 2018b). CMIP3 and CMIP5 models do not accurately reproduce the observed seasonal cycle of precipitation over the major river basins of South and South East Asia, limiting the attribution of observed regional hydroclimatic changes ( [[#Hasson--2014|Hasson et al., 2014]] , 2016; [[#Biasutti--2019|Biasutti, 2019]] ). While warm rain processes and organized convection are known to dominate the heavy orographic monsoon rainfall over the Western Ghats mountains ( [[#Shige--2017|Shige et al., 2017]] ; [[#Choudhury--2018|Choudhury et al., 2018]] ), in various parts of India ( [[#Konwar--2012|Konwar et al., 2012]] ) and East Asia ( [[IPCC:Wg1:Chapter:Chapter-11#11.7.3.1|Section 11.7.3.1]] ), there are uncertainties in representing the regional physical processes of the monsoon environment, including cloud – aerosol interactions ( [[#Sarangi--2017|Sarangi et al., 2017]] ), land – atmosphere (e.g., Bartonet al., 2020) and ocean – atmosphere coupling ( [[#Annamalai--2017|Annamalai et al., 2017]] ), in state-of-the-art climate models (see also [[#8.5.1|Section 8.5.1]] ).

In summary, there is ''high confidence'' in observational evidence for a weakening of the SAsiaM in the second half of the 20th century. Results from climate models indicate that anthropogenic aerosol forcing has dominated the recent decrease in summer monsoon precipitation, as opposed to the expected intensification due to GHG forcing ( ''high confidence'' ). On paleoclimate time scales, the SAsiaM strengthened in response to enhanced summer warming in the NH during the early-to-mid Holocene, while it weakened during cold intervals ( ''high confidence'' ). These changes are tightly linked to orbital forcing and changes in high-latitude climate ( ''medium co'' ''nfidence'' ).

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