the Southern Hemisphere in winter to the Northern Hemisphere in summer. The largest northward movement of the rain belt takes place over the Indian monsoon region where it moves from a mean position of about 5°S in winter to about 20°N in summer. Equatorial easterlies in the upper atmosphere are weak and confined between 5°N and 10°S, while the subtropical westerlies intrude up to 10°N during winter. The subtropical westerlies recede to north of 30°N during summer and a strong easterly jet characterises the equatorial upper atmosphere. Seasonal rainfall is highly variable with prolonged spells of dry and wet conditions often lasting for two–three weeks. Periods of wet conditions are generally associated with an increase in cyclonic activity and decrease in surface pressure over the central Indian monsoon trough region (northern Bay of Bengal to West India) and strengthening of the low‐level westerly jet over the Arabian Sea.
The East Asian monsoon features powerful breaks of cold air from the Eurasian continent during winter. During summer, the Tibetan Plateau acts as an elevated heat source and fosters an imposing subtropical monsoon front extending from the Bay of Bengal and the South China Sea all the way to the North Pacific, bringing abundant rainfall to East Asia, and transporting water and energy from the tropics to the mid‐latitudes. In contrast to warm trends elsewhere on earth, an interdecadal surface cooling has occurred in some regions of East Asia in the past few decades. Such interdecadal changes have been attributed to tropical ocean forcing. For example, in the late 1970s, summer rainfall anomalies in China were consistent with an intensification and southward extension of the western Pacific subtropical high over the subtropical regions of East Asia. Changes in SSTs over the Indian Ocean and tropical western Pacific may have partly accounted for these interdecadal changes.
2.4.2 The Indo‐Australian Monsoon
The Indo‐Australian monsoon is a component of the large‐scale Asian‐Australian monsoon system. The monsoon is associated with the seasonal migration of the ITCZ, which is located 10–15° north of the equator in austral winter and migrates southward over the north of the Australian continent during austral summer, driven by the development of intense heat lows over NW Australia and NW Queensland. North Australia experiences dry low‐level southeast trade winds in winter which undergo a reversal to dominant moist NW winds in summer. The onset of the monsoon usually occurs between late November and early January, but the monsoon season includes periods when southeast winds bring temporarily dry conditions.
The southern limit of the monsoon is defined by areas that receive >85% of the rainfall between November and April. Main rainfall events are associated with cyclones and other disturbances in the monsoon trough. Such conditions do not necessarily translate into high annual rainfall as there is a strong latitudinal gradient of rainfall with large regions along the northern and eastern coastlines having median rainfalls exceeding 1500 mm, but with most of the inland experiencing annual medians from 300 to 600 mm.
The Indo‐Australian wet monsoon season is characterised by excessively wet conditions usually referred to as ‘bursts’ interspersed with relatively dry ‘breaks’; both can sometimes extend for several weeks (Moise et al. 2020). Both Indonesia and Australia are linked in the summer monsoon. These wet and dry periods have a significant impact on recharge of freshwater ecosystems and river discharge on estuarine and marine ecosystems. A ‘burst’ in the Indo‐Australian monsoon is preceded by the development of a well‐defined extratropical wave packet in the Indian Ocean which propagates towards the Australian continent in the few days leading up to the onset of heavy rainfall in the tropics. These extratropical disturbances propagate equatorward over the continent and are accompanied by lower‐tropospheric air mass boundaries which also propagate into low latitudes. Ahead of these boundaries, relatively warm moist air is advected from the surrounding seas. Monsoon bursts are more likely to occur when the active phase of the Madden‐Julian Oscillation (Section 2.6) is in the vicinity of Australia (Duan et al. 2019).
2.4.3 The African Monsoons
The main two components of the African Monsoons are the West African Monsoon (WAM), which prevails during the Northern Hemisphere summer (June through September), and the East African Monsoon (EAM) with occurs during spring (March–May) and autumn (October–December). The combined influence of the Indo‐Pacific and the Atlantic Oceans drive the interannual and decadal monsoon variability over these regions.
The key feature of the WAM is low‐level SW flow from the Atlantic Ocean and the Intertropical Convergence Zone (ITCZ) north of the equator. The WAM is characterised by the migration of zonally banded rainfall from the Guinea coast to the Sahel (a 3860‐km arc‐like landmass immediately south of the Sahara Desert stretching east–west across the breadth of the African continent) and back again, resulting in two rainy seasons per year in the south and one in the north. The dynamics of the WAM circulation are linked to the existence of the African easterly jet, a mid‐tropospheric flow with peak amplitude of about 10–20 m s−1 in August at latitude 15°N. This jet is a basic part of the momentum balance of the WAM circulation and has a meteorological impact on the patterns of rainfall in the Sahel. Unlike the Australian monsoon, over West Africa the monsoon trough moves much further inland, some 15° or more in latitude from the Guinea coast, and therefore the pressure gradients over much of the Sahel are directed from the humid south towards the dry north. There is a seasonal cycle of the WAM with two phases: the ‘pre‐onset’ and the ‘onset’ (Janicot et al. 2011). In the ‘pre‐onset’ phase, the average (10°W − 10°E) of the intertropical front (the confluence of monsoonal winds from the south and dry winds from the north) crosses 15°N on its northward migration, while the ITCZ is still in the south. The first rainy season along the Guinean coast occurs between mid‐April and the end of June. In the ‘onset’ phase during the summer monsoon season, the ITCZ moves from 5°N to 10°N bringing rain over the Sahel.
The formation of the Atlantic cold tongue is the dominant seasonal SST signal in the eastern equatorial Atlantic and influences the onset of the WAM (Caniaux et al. 2011). From March to mid‐June (‘onset phase’), the cold tongue results from the intensification of the southeast trade winds associated with the anticyclone system located off the island of St. Helena. Steering of surface winds by the basin shape of the eastern equatorial Atlantic imparts optimal wind stress for generating the maximum upwelling south of the equator. From mid‐June to August, wind speeds north of the equator increase due to the northward progression of the intensifying trade winds and due to significant surface heat flux gradients produced by the differential cooling between the Atlantic cold tongue and the tropical waters circulating in the Gulf of Guinea. Thus, there is a close link between the eastern equatorial Atlantic and the West African monsoon.
The East African monsoon system corresponds with the Greater Horn of Africa region, extending from Tanzania in the south to Yemen and Sudan in the north. During winter, when NE trade winds flow across the NW Indian Ocean and equatorial moisture moves over the Indian Ocean exhibiting strong westerly flows over the equatorial Indian Ocean, East African rainfall is limited to a few highland areas (Funk et al. 2016). During spring, the Indian monsoon circulation transitions, the trade winds over the NW Indian Ocean reverse, and East African moisture convergence supports the ‘long’ rains. In summer, the SW Somali Jet intensifies over East Africa. As subsidence forms along the westward flank of this jet, precipitation shuts down over eastern portions of East Africa. In autumn, the Jet subsides but easterly moisture supports rainfall in limited regions of the eastern Horn of Africa. Thus, Indian Ocean SSTs drive East African rainfall variability by altering the local Walker circulation (see Section 2.6). The influence of the Pacific Ocean appears to be minimal (Tierney et al. 2013). SST anomalies were the dominant forcing of the severe drought of the 1970s and 1980s (Biasutti 2019). However, ENSO showed a positive relationship with the East African ‘short’ rain during the 1949–2016 period. Only during a recent period (2000–2016) was a significant relationship found between ENSO and East African ‘long’ rain. The strengthened interannual relationship
