ocean responses that have some bearing for life in tropical seas. The oceanic mixed layer, for instance, is a direct consequence of the surface cooling in convection centres of the MJO, and warming outside, fluctuations in SST propagate eastwards in tandem. SST differences can be about 0.5 °C. Strong surface wind of the MJO forces ocean currents and hence possible effects of horizontal advection. Strong winds force eastward equatorial currents of ≈ 1 m s−1 near the surface which may penetrate to 100 m depth, affecting the movements of pelagic organisms.
The pulse‐like structure of the MJO forces pulses of downwelling Kelvin waves (Zhang 2005). They propagate from their origin in the western Pacific to the eastern Pacific where the MJO is weak or absent. Vertical displacement of the thermocline thus occurs, typically to a depth of 20–30 m. This can affect ENSO events. That is, in the central Pacific near the eastern edge of the Indo‐Pacific Warm Pool, the eastward surface current of the Kelvin wave results in advection of warmer water eastward. In the eastern Pacific, the displacement of the thermocline associated with the downwelling Kevin waves weakens the cooling of equatorial upwelling, leading to warmer equatorial SSTs.
The MJO disturbs the upper ocean through surface fluxes of momentum, latent, and sensible heat, radiation and freshwater, with the latter three accounting for buoyancy flux. The net freshwater flux into the ocean (P‐E) is mainly controlled by rainfall, as strong evaporation in convective centres of the MJO compensates only slightly for the freshwater input. The net result is that perturbations in solar radiation flux (controlled by cloudiness) and latent heat flux (mostly controlled by surface winds) have similar amplitudes (25–30 W m−2). The intra‐seasonal amplitude of the net heat flux (mostly composed of radiation and latent heat) depends on the relative phase of different components of the MJO.
The Pacific Decadal Oscillation (PDO) is the dominant year‐round pattern of monthly North Pacific SST variability and is often described as a long‐lived El Niño‐like pattern in the tropical Pacific (Vishnu et al. 2018). The PDO is not a single phenomenon but is instead a complex aggregate of different atmospheric and oceanographic forcing spanning the extratropical and tropical Pacific (Newman et al. 2016). The PDO's amplitude is greatest from November–June, with weak maxima both in mid‐winter and late spring and a pronounced late summer‐early autumn minimum (Newman et al. 2016).
Positive (negative) phases of the PDO are associated with warming (cooling) of the tropical Pacific Ocean. The PDO modulates climate variability in various parts of the globe, such as drought frequency in the United States and summer monsoon rainfall in south China. Positive (negative) phases of the PDO are associated with the deficit (excess) Indian summer monsoon rainfall and enhance (suppress) the teleconnection between the rainfall in India and ENSO (Vishnu et al. 2018). The frequency of tropical cyclones over the western North Pacific also shows a decadal variability associated with the PDO. The number of tropical cyclones across the Pacific is less (high) in the warm (cold) phases of the PDO. There is also an out‐of‐phase variation in the number of monsoon depressions over the Bay of Bengal and the PDO. Vishnu et al. (2018) postulate that the variation in SSTs in the western equatorial Indian Ocean associated with the PDO could be one of the reasons for the changes in the moisture advection over the Bay of Bengal and hence the variation in the number of monsoon depressions on an interdecadal timescale.
The positive and negative phases of the PDO may have an impact on the expansion of the poorly oxygenated regions of the eastern Pacific Ocean (Duteil et al. 2018). During a ‘typical’ positive phase of the PDO, modelling indicates that the volume of the suboxic regions expanded by 7% over a 50 year period due to a slowdown of the large‐scale circulation related to the decrease in the intensity of the trade winds. The model suggested that the prevailing positive phase conditions of the PDO since 1975 may explain a significant part of the current deoxygenation of the eastern Pacific Ocean.
2.7 Climate Change: Physical Aspects
Humankind has had and is still having a direct impact on earth's climate. Since the beginning of the Industrial Age there has been an increase in the concentrations of carbon dioxide (Figure 2.10 top), methane and nitrous oxide in the atmosphere, the direct result of the burning of fossil fuels, and to a lesser extent, deforestation. The increases in atmospheric greenhouse gases have had a direct impact on the global ocean, warming the earth's seas which in turn has resulted in a rise in sea level via thermal expansion. The uptake of CO2 has resulted in a decrease in ocean pH (Figure 2.10 bottom) and CO32− concentration, a process termed ocean acidification. Other impacts of human‐induced climate change include increases in air and ocean temperature, an increase in OHC and changes in precipitation patterns.
Warming of the global ocean is the largest near the surface, the upper 75 m warmed by 0.11 °C per decade over the period 1971–2010 (IPCC 2014). It is likely that the ocean warmed from the 1870s to 1971. Regions of naturally high salinity have become more saline, while regions of low salinity have become fresher since the 1950s. Such regional trends in ocean salinity provide indirect evidence for changes in evaporation and precipitation over the ocean and for changes in the global water cycle (IPCC 2014). Since the beginning of the Industrial Era, oceanic uptake of CO2 has resulted in a decline in surface ocean pH (Section 2.7.2). There is medium confidence (IPCC 2014) that in parallel to warming, oxygen concentrations have declined in coastal waters and in the open ocean thermocline since the 1960s. OMZs are progressively expanding in the tropical Pacific, Atlantic, and Indian Oceans due to reduced ventilation and oxygen solubility in warmer, more stratified oceans (Stramma et al. 2011).
FIGURE 2.10 Trends in surface (< 50 m depth) ocean carbonate chemistry calculated from observations obtained during the Hawaii Ocean Time‐Series Program in the North Pacific from 1988 to 2017. The upper graph shows the concomitant increase in CO2 concentrations in both the atmosphere (red) and surface ocean (blue), presented as CO2 concentration in air (ppm). The bottom graph shows a decline in ocean pH (light blue, primary y‐axis) and carbonate ion (CO32−) concentration (green, secondary y‐axis on right).
Source: Doney et al. (2020), figure 1, p. 87. Licensed under CC BY 4.0. © Annual Reviews.
Over the period 1901–2010, global mean sea level rose by 0.19 m. The rate of sea‐level rise (SLR) since the mid‐nineteenth century has been larger than the mean rate during the previous two millennia (IPCC 2014). The mean rate of SLR was 1.7 mm a−1 between 1901 and 2010 with an increase to 3.2 mm a−1 between 1993 and 2010. Over this period, global mean SLR has been consistent with the sum of the observed contributions from ocean thermal expansion, the Greenland ice sheet, the Antarctic ice sheet, and land water storage. Rates of mean SLR vary over different regions due to fluctuations in ocean circulation. It is extremely likely (IPCC 2014) that more than half of the observed increase in global average surface temperature and mean SLR from 1951 to 2010 was caused by the anthropogenic increase in greenhouse gas concentrations and other anthropogenic forcing.
The recent climatological forecasts by the Intergovernmental Panel on Climate Change (IPCC) for until the end of this century (Church et al. 2013; Collins et al. 2013; Bindoff et al. 2019; Oppenheimer et al. 2019) predict that globally: (i) SSTs will rise by 1–3 ° C; (ii) oceanic pH will decline by 0.07–0.31 units; and (iii) mean atmospheric CO2 concentrations will increase to 441 ppm (from 391 ppm in 2011). Regional differences (Table
