10. Indian Ocean Warming

The oceans have absorbed approximately 93% of the additional heat due to anthropogenic global warming since the 1950s (Cheng et al. 2017) and have resulted in a significant increasing trend in the global average ocean SST, as evidenced by modern instrumental records (Deser et al. 2010). Among the oceans, Indian Ocean stands out as one of the most rapidly warming ocean basins (Gnanaseelan et al. 2017; Beal et al. 2019). The global average rise in SST during 1951–2015 is 0.7 °C (0.11 °C/decade), while the TIO SST has risen by about 1.0 °C on average (0.15 °C/decade) (high confidence) (Fig. 10.1). A 138-year (1870–2007) monthly observed SST time series averaged along the ship track extending from the Gulf of Aden through the Malacca Strait also reveals a warming of 1.4 °C during the entire period (Chowdary et al. 2012). Coherently, historical simulations by climate models participating in the Coupled Model Inter-comparison Project (CMIP5), with CO₂ forcing at observed rates, show warming trends of 0.1–0.18 °C per decade during 1976–2005 (Fig. 10.1b), with maximum warming trends over the northern Arabian Sea.

Warming in the TIO has been basin-wide but spatially non-uniform, with the largest increasing trends seen in the central equatorial Indian Ocean and lowest warming trends off the Sumatra and Java coasts. While long-term warming (1900–2015) is maximum over the western Indian Ocean (Roxy et al. 2014), warming in recent decades is prominent over the central equatorial and subtropical south Indian Ocean (Fig. 10.1a). SST warming trends during the recent period (2000–2013) also exhibit an interhemispheric difference, with relatively weak warming to the north of 10° S and accelerated warming to the south of 10° S (Dong and McPhaden 2016).

The observed surface warming over the Indian Ocean has been linked to natural and anthropogenic causes. Climate model simulations show that over 90% of the SST trend since the 1950s is very likely due to increased anthropogenic emissions (Du and Xie 2008; Dong and Zhou 2014; Dong et al. 2014), while the remaining is due to internal variability (Dong et al. 2014). Among the anthropogenic causes, change in radiative forcing due to the increased greenhouse gas concentrations is the major factor (Du and Xie 2008). Over recent decades, increasing atmospheric pollutants, known as aerosols, have likely dampened the greenhouse gas forced warming of the Indian Ocean (Dong and Zhou 2014).

Changes in the tropical circulation, ocean-atmospheric interaction and dynamics play a role in the observed distribution of the warmer waters in the Indian Ocean. Redistribution of heat from the Pacific Ocean via the Walker circulation (Roxy et al. 2014) and the ITF (Dong and McPhaden 2016) is one of the reasons for the observed patterns of warming in the Indian Ocean. Other than the ITF, Indian Ocean may also be receiving a warming signal via the deep meridional overturning circulation, and from the Southern Ocean (Gille 2002). Contrary to expectation, there is a negative trend in the net heat flux despite the warming trend in the Indian Ocean SSTs (Rahul and Gnanaseelan 2013). This suggests that local ocean dynamics and ocean-atmosphere interaction also have a major role in the observed warming pattern in the Indian Ocean (Lau and Nath 2000; Du et al. 2009; Rahul and Gnanaseelan 2016; Pratik et al. 2019; Rao et al. 2012). The period 2000–2013 witnessed strong warming in the southern Indian Ocean south of 10° S inducing a north–south SST gradient. This interhemispheric gradient is forced primarily by an increased ITF, from the Pacific into the Indian Ocean, induced by stronger Pacific trade winds (Dong and McPhaden 2016; Lee et al 2015).

Future projections using CMIP5 simulations clearly show SST warming in the Indian Ocean with increasing anthropogenic emissions (high confidence). However, SST warming is projected to be non-uniform with regional variations in the Arabian Sea and the Bay of Bengal. Most models project a higher SST warming in the Arabian Sea than the Bay of Bengal. This is consistent with the changes that have been observed in the Indian Ocean in the last several decades (Zhao and Zhang 2016).

Changes in SSTs projected for the end of the twenty-first century show regional and seasonal variability (Cai et al. 2013). The ensemble mean of CMIP5 RCP scenarios indicates stronger warming in the north-western Indian Ocean and weaker warming off the Sumatra and Java coasts (Zheng and Xie 2009). The CMIP5 RCP scenarios for the period 2040–2069 and 2070–2099 are shown in Fig. 10.2, where the stronger warming in the north-western Indian Ocean and weaker warming in the south-eastern Indian Ocean south of 15° S with respect to the base period of 1976–2005 are evident, and consistent with Zheng and Xie (2009). Patterns of warming are similar in both the RCP4.5 and RCP8.5 scenarios, although the magnitude of warming is much larger in the latter scenario (Fig. 10.2). RCP4.5 projects a warming rate of 0.13 °C/decade in the TIO (Fig. 10.2e), similar to the current rate of warming (Fig. 10.1c). Meanwhile, RCP8.5 indicates an accelerated warming at the rate of 0.35 °C/decade. CMIP5 (the ensemble mean) projected SST rise in the TIO in the near and far future for both the RCP4.5 and RCP8.5 scenarios are tabulated in Table 10.1.

The strong SST warming trend in the north-west basin accompanied by an increase in precipitation and weak SST warming trend in the south-east basin accompanied by decrease in precipitation drive strong surface easterly wind anomalies along the equatorial Indian Ocean (Li et al. 2016). Along with the east–west gradient in SST and precipitation under global warming scenarios, the thermocline in the east equatorial Indian Ocean also shoals (a favourable condition for the formation of IOD) as a result of a weakened Indian Ocean Walker cell and easterly wind change along the equator (Zheng and Xie 2009). The shoaling of the thermocline then strengthens the thermocline feedback in this region. Such a pattern of SST, precipitation and thermocline feedback would result in more frequent occurrences of extreme positive IOD (pIOD) events in the future, from one event every 17.3 years over the twentieth century to one event every 6.3 years by the end of twenty-first century (low confidence) (Cai et al. 2014). This suggests increased risks of climate and weather extremes in regions impacted by extreme pIOD events.

Upper OHC exhibits an increasing trend in the Indian Ocean since the 1950s (high confidence, Levitus et al. 2009; Xue et al. 2012; Han et al. 2014). The global OHC700 has risen at a rate of 2.8 × 10²² Joules per decade over the period 1955–2015 (Fig. 10.3a). Over the same period, the rate of rise in OHC700 in the tropical (30 °S to 30 °N and 40 °E to 115 °E) and north (5 °N to 30 °N and 40 °E to 100 °E) Indian Ocean basins were more muted at 0.62 × 10²² Joules per decade and 1.0 × 10²² Joules per decade, respectively (Fig. 10.3b). In other words, there was greater net heat content gain in the north Indian Ocean (NIO) than the TIO over 1955–2015 despite being a smaller basin in areal extent.

It is important to note that since the year 2000 both the TIO and NIO have experienced a steep rise in OHC (Cheng et al. 2017), which is absent in the global OHC signal (Fig. 10.3a, b). This abrupt increase of OHC in TIO has accounted for more than 70% of the global ocean heat gain in the upper 700 m during the same period (e.g. Lee et al. 2015), despite of representing only about 15% of the global ocean area. It is very likely that a significant portion of the heat due to greenhouse warming now resides in the upper 700 m of the Indian Ocean. Given the fact that the OHC700 in the Indian Ocean did not increase significantly during 1971–2000, the rapid increase during the recent period is striking.

In addition, higher variability in the TIO and NIO upper OHC compared to the global OHC is evident from Fig. 10.3. Changes in OHC700 in the NIO from 1955 to 2015 (Fig. 10.3c) show wide spatial variations with warming in the Arabian Sea, Bay of Bengal, equatorial Indian Ocean and in the region south of 20° S (Fig. 10.3c) (e.g. Nagamani et al. 2016; Anandh et al. 2018), but a zonally extended cooling from 20° S to 5° S. Trends in TIO OHC2000 (ocean heat content in the upper 2000 m) during 2005–2015 are similar to that of OHC700 (Fig. 10.3b).

There are significant differences in the recent OHC700 trends between the major ocean basins, particularly the Pacific and Indian Oceans. Despite of large natural variability in OHC700, increasing trends due to anthropogenic influence are evident (high confidence). The abrupt increase in the OHC700 during 2003–2012 was not due to surface heating, but due almost entirely to horizontal heat convergence in the form of an enhanced ITF (Lee et al. 2015). This Indian Ocean OHC increase corresponds to a concurrent Pacific Ocean OHC decrease in the 0–100 m, suggesting a transfer of heat from Pacific Ocean to Indian Ocean (Liu et al. 2016). The net surface heat flux into the Pacific Ocean increased greatly during 2003–2012, consistent with the La Niña-like condition across the Pacific Ocean. However, the anomalous surface heat uptake was completely masked by horizontal heat divergence in the tropical Pacific, which was mainly traced to the increased heat transport to the Indian Ocean. This has, in fact, inhibited the increase in tropical Pacific OHC700, which instead shows a slight decrease during this period.

Among the global oceans, the ensemble-mean trends for the 700–6000 m layer show the largest increasing trends in the North Atlantic sub-polar gyre, north Indian Ocean and Southern Ocean (e.g. Palmer et al. 2017). As there are fewer observations available in the Indian Ocean between 3000 and 5000 m compared to the other oceans, deep OHC estimates in the Indian Ocean, particularly in the north-western part, are relatively less accurate (Purkey and Johnson 2010). Although the Atlantic and Indian oceans contributed to the global OHC increase at 3500 m depth, the OHC increases in these two basins have been weak below 4000 m during the past two decades. The smaller contribution of the Indian Ocean to the global OHC increase might be due to the smaller bottom layer volume of the Indian Ocean than that of the Pacific Ocean (e.g. Kouketsu et al. 2011).

The observed upper OHC of the Indian Ocean reveals significant decadal variations (Han et al. 2014; Mohapatra et al. 2020). Model experiments suggest that the observed multidecadal trends in OHC are partially associated with anthropogenic forcing. The observed decadal variability in the basin-wide distribution of sea level and thermal structure in the Indian Ocean results primarily from forcing by Indian Ocean winds (Deepa et al. 2019; Srinivasu et al. 2017), with a significant contribution from the ITF transport to the interior of the south Indian Ocean after 1990 (Han et al. 2014). For the Pacific and Indian Oceans, decadal shifts are primarily observed in the upper 350 m, likely due to shallow subtropical circulation, leading to an abrupt increase of OHC in the Indian Ocean carried by the ITF from the Pacific Ocean over the last decade (Cheng et al. 2017). During the slowdown in Pacific SST warming, there was anomalous warming in the Indian Ocean and an accelerated OHC rise below 50 m, which is associated with a La Niña-like climate shift, and an enhanced heat transport of the ITF (Liu et al. 2016). Transmission of the multidecadal signal occurs via an oceanic pathway through the ITF and is manifested across the Indian Ocean centred along 12° S as westward-propagating Rossby waves modulating the thermocline and subsurface heat content variations (Rahul and Gnanaseelan, 2016; Deepa et al. 2019). Changes in Pacific wind forcing in recent decades and associated rapid increases in Indian Ocean subsurface heat content can thus affect the basin’s leading mode of variability (Ummenhofer et al. 2017). Jin et al (2018) argued that the western Indian Ocean subsurface heat content is influenced by Interdecadal Pacific oscillation (IPO) through wind driven Ekman pumping via the atmospheric bridge, whereas the eastern Indian Ocean is largely affected through the oceanic pathway via ITF. The recent negative phase of IPO (1998–2012) enhanced the Pacific easterlies, which eventually led to export of anomalous heat from the Pacific towards the Indian Ocean (Gastineau et al 2018). The OHC300-based definition of PDO takes into account variations throughout the upper ocean and is better suited to capture various characteristics of PDO variability (Kumar and Wen 2016). Based on multiple observational datasets, ocean reanalysis products and an ocean model simulation, Li et al. (2018) reported the presence of prominent multidecadal variations in the Indian Ocean OHC400. They suggested that the upper Indian Ocean first experienced a heat content increase at the rate of 5.9 ± 2.5 × 10²¹ J decade⁻¹ during 1965–79, followed by a decrease at the rate of −5.2 ± 2.5 × 10²¹ J decade⁻¹ during 1980–96, and subsequently an enhanced rate of increase of 13.6 ± 1.1 × 10²¹ J decade⁻¹ from 2000 to 2014. This suggests that the Indian Ocean OHC underwent tremendous decadal variations which might continue into the future. CMIP5 simulations have limited skill in capturing decadal variability in upper OHC during the past 45 years (Collins et al. 2013; Cheng et al. 2015).

Monitoring and understanding OHC change and the role of circulation in shaping the patterns of increase remain key to predicting global and regional climate change, and sea-level rise (Zanna et al. 2019). Under the low-to-medium (RCP4.5) emissions scenario, half of the energy taken up by the ocean by the end of the twenty-first century will be in the uppermost 700 m, and 85% will be in the uppermost 2000 m (low confidence). Future changes in wind and air-sea fluxes, and ocean transport, likely to have serious implications for regional sea-level rise and coastal flood risk. Additionally, the spatial patterns of OHC change under global warming contribute to the regional sea-level projections (e.g. Slangen et al 2014). There is a large spread among CMIP5 models in projections of future changes in OHC, suggesting an urgent need for further refinement (Cheng et al 2016; Allison et al. 2019).

Rapid warming of the Indian Ocean during 1950–2015, along with substantial changes in land use and anthropogenic aerosols have altered the Indian summer monsoon (Singh et al. 2019, Chaps.  3 and 6). During 1950–2015, there has been a significant decline in the summer monsoon rainfall over central India and parts of north India due to a reduction in the tropospheric thermal contrast that is associated with the rapid warming of the Indian Ocean (Mishra et al. 2012; Saha et al. 2014; Roxy et al. 2016). At the same time, rapid warming in the Arabian Sea has resulted in a rise in widespread extreme rains over Western Ghats and central India, since warming induces increased fluctuations in the monsoon winds, with ensuing episodes of enhanced moisture transport from the Arabian Sea towards the Indian subcontinent (Roxy et al. 2017). Indian Ocean warming is also found to reduce rainfall over India during the onset phase and increase it during the withdrawal phase (Chakravorty et al. 2016).

In terms of tropical cyclones, the Bay of Bengal region witnesses more than 80% of the global fatalities associated with tropical cyclones, while only accounting for 5% of these storms globally (Beal et al. 2019). Since tropical cyclones primarily draw their energy from evaporation at the ocean surface, SST and OHC strongly constrain cyclone intensity (Rajeevan et al. 2013). Global warming appears to have increased the intensity of tropical cyclones during the post-monsoon period in the Bay of Bengal (Chap. 8) and the pre-monsoon period in the Arabian Sea. For example, Cyclone Nilofar in 2014 was the first severe tropical cyclone to be recorded in the Arabian Sea in the post-monsoon season. Though the cyclone did not make landfall, it produced heavy rainfall along the western coast of India. Future projections suggest a likely increase in the number of extremely severe tropical cyclones in response to Indian Ocean warming, particularly in the Arabian Sea, while changes in frequency remain uncertain (Chap. 8).

Indian Ocean SSTs have a role in regulating the surface air temperatures over the Indian subcontinent (Chowdary et al. 2014). Associated with the basin-wide warming and frequent El Niños, the frequency and duration of heat waves have increased over the Indian subcontinent (Rohini et al. 2016). Rising ocean temperatures have also resulted in instances of marine heat waves in the Indian Ocean. Marine heat waves are similar to heat waves over the land, with periods of extremely high ocean temperatures that persist for days to months (Collins et al. 2019). Recent marine heat waves, including the one in 2016 that co-occurred with an extreme El Niño event of 2015–16, resulted in mass bleaching of coral reefs and adversely impacted aquaculture industries along the Indian Ocean rim countries (Collins et al. 2019). Satellite observations reveal that the intensity of marine heat waves have increased and that they have very likely doubled in frequency over 1982–2016. Climate projections indicate that at the high emissions RCP 8.5 GHG scenario, a one-in-100-day marine heat wave event (with pre-industrial CO₂ levels) is very likely to become a one-in-four-day event by 2031–2050 and a one-in-two-day event by 2081–2100 (Collins et al. 2019).

The impact of Indian Ocean warming is reflected in the sea-level changes, as thermal variations have dominated these changes in recent decades (Chap. 9). The largest sea-level changes were observed along the northern and eastern coasts of the Bay of Bengal (Chap. 9). Tide gauge data corroborates the increasing sea level among the coastal regions of the Indian Ocean, such as Mumbai, Kochi, Visakhapatnam on the Indian coast and Durban, Fremantle, Port Hedland on the Australian coast. The consistent increase in sea level is attributed to the thermal expansion of sea water, due to a basin-wide surface warming in the Indian Ocean. Besides, the observed changes in the wind circulation in the Indian Ocean has modulated the ocean heat transport, and distributed the heat across the basin, resulting in a large thermosteric response in the sea level (Chap. 9).

Beyond the regional climate, Indian Ocean warming has global and remote impacts also. The Indian Ocean has contributed to more than 21% of the global oceanic heat uptake over the last two decades and contributed strongly to the temporary slowdown in global warming during 1998–2013 (Lee et al. 2015; Cheng et al. 2017). Climate model experiments indicate that the rapid warming of the Indian Ocean is strengthening the Atlantic meridional overturning circulation (Hu and Fedorov 2019; Cherchi 2019). The basin-wide warming could modulate the Pacific climate, affect the North Atlantic oscillation and enhance the positive Southern Annular Mode, and may cause West Sahel and Mediterranean droughts (Beal et al. 2019). Atmospheric blocking triggered by tropical convection in the Indian and Pacific oceans can cause persistent anticyclonic circulation that not only leads to severe drought but also generates marine heat waves in the adjacent ocean. Warming in the Indian Ocean and associated deep convection is found to trigger droughts in South America and marine heat waves in the adjoining South Atlantic (Rodrigues et al. 2019). A strong negative IOD event in 2016 strongly impacted East African rainfall, with some regions recording below 50% of normal rainfall, leading to devastating drought, food insecurity and unsafe drinking water for over 15 million people in Somalia, Ethiopia and Kenya (Collins et al. 2019).

The biogeochemical properties of the Indian Ocean are distinct from other ocean basins for mainly two reasons—the land boundary in the north due to Indian subcontinent and the large amplitude of seasonally reversing monsoon cycle. Climate driven physical fluctuations are expected to impact the marine ecosystem substantially by modifying the biotic and abiotic environments, which can lead to severe repercussions for the oceanic primary production. The warming of SSTs in the western Indian Ocean leads to increased stratification in the basin. Roxy et al. (2016) show a declining trend in marine phytoplankton in the western Indian Ocean (high confidence), by 30% in the observations during 1998–2013 and 20% in the CMIP5 simulations during 1950–2015 (medium confidence). This significant decline in phytoplankton is attributed to enhanced stratification of the oceanic water column as a result of rapid surface warming, thereby suppressing the mixing of nutrients from subsurface layers into the surface. Downward trends in primary production over the Indian Ocean can be detrimental to the marine food web and the fishing industry, especially the economically valuable tuna industry (Lee et al. 2005). However, careful gathering of in situ observations in the open ocean is needed to build a substantial understanding of the bio-physical interactions in the Indian Ocean.

Approximately 30% of the historical anthropogenic CO₂ emissions have been absorbed by the oceans since the pre-industrial era (e.g. Canadell et al. 2007). The increasing oceanic uptake of CO₂ has changed seawater chemistry and resulted in ocean acidification, with profound impacts on biological ecosystems in the upper ocean. Long-term increasing trends in ocean acidification, consistent with the increase in atmospheric CO_2, are evident over the past several decades (Dore et al. 2009).

Recent work by Sreeush et al. (2019a) has found that the Indian Ocean is acidifying due to the accumulation of anthropogenic CO₂ from the atmosphere, and this storage of anthropogenic carbon in the Indian Ocean is comparable (after normalizing for the size of the basins) with the other major oceans (Sabine et al. 2004). Surface ocean pH over the Indian Ocean has declined by about 0.1 unit (current mean is 8.1) relative to pre-industrial levels and is larger over the western Indian Ocean (e.g. Sreeush et al. 2019a). This increase in ocean acidity may be responsible for the functional collapse of reef building corals. The western Arabian Sea has undergone more rapid acidification than the rest of the TIO basin due to strong upwelling in this region drawing up anthropogenic CO₂ embedded in the deeper ocean. Moreover, SST warming also accelerates acidification due to the endothermic nature of CO₂ dissolution in water. An ocean biogeochemical model-based simulations from 1960 to 2009 show that western Arabian Sea has acidified by 108% due to dissolved inorganic carbon, −36% due to buffering due to alkalinity, 16% due to SST warming, 6% due to salinity changes and remaining due to changes in other minor ions (Sreeush et al. 2019a). Considering that the western Arabian Sea is a highly productive zone of the Indian Ocean (Roxy et al. 2016), the role of SST warming in exacerbating acidification needs to be monitored carefully.

Dissolved O₂ is a major determinant of the abundance and distribution of the marine habitat. The Indian Ocean contains one of the oceans’ most pronounced oxygen minimum zones (OMZs) encompassing more than 50% of the area containing OMZs (e.g. Helly and Levin 2004). The tropical open ocean O₂ concentration has decreased at a rate of 0.1–0.3 µ mol kg⁻¹ year⁻¹ during the past five decades (Stramma et al. 2008). Long-term measurements over the TIO also show a pronounced decrease (at a rate of 20–30 mol m⁻² per decade) in O₂ concentration (Koslow et al. 2011). This oxygen decline is a result of warming-induced decline in oxygen solubility and reduced ventilation of deep ocean due to enhanced ocean surface stratification (Keeling et al. 2010).

Future projections based on the CMIP5 models indicate further warming of the Indian Ocean and a resultant decline in marine primary productivity (Fig. 10.4, Bopp et al. 2013). CMIP5 models also project a continuing decline in open-ocean surface pH from the current mean of 8.1–7.8 by the end of the twenty-first century, as a response to the projected rise in CO₂ emissions (Bopp et al. 2013). Most models project decreasing oxygen concentrations in the ocean by 1–7% from present-day concentrations by 2100, with global warming (Keeling et al. 2010; Bopp et al. 2013).

Interactions of rising temperatures, ocean acidification and oxygen minimum zones narrow thermal ranges and enhance sensitivity to temperature extremes in marine organisms ranging from corals to fish. These negative effects of increase in warming are projected to be most significant in developing nations in tropical regions (high confidence, IPCC AR5 2014). However, understanding the combined effect of both low O₂ and low pH on marine ecosystems in the Indian Ocean demands more active research in the Indian Ocean.