Chapter 5. Sea-Air Interactions
Contributors: Jeremy T. Mathis (Convenor), Jose Santos, Renzo Mosetti Alberto Mavume, Craig Stevens, Regina Rodrigues, Alberto Piola, Chris Reason Patricio A. Bernal (Co-Lead member), Lorna Inniss (Co-Lead member)
1. Introduction
From the physical point of view, the interaction between these two turbulent fluids, th ocean and the atmosphere, is a complex, highly nonlinear process, fundamental to th motions of both. The winds blowing over the surface of the ocean transfer momentu and mechanical energy to the water, generating waves and currents. The ocean in tur gives off energy as heat, by the emission of electromagnetic radiation, by conduction and, in latent form, by evaporation.
The heat flux from the ocean provides one of the main energy sources for atmospheri motions. This source of energy for the atmosphere is affected by the turbulence at th air/sea interface, and by the spatial distribution of the centres of high and low energ transfer affected by the ocean currents. This coupling takes place through processe that fundamentally occur at small scales. The strength of this coupling depends on air sea differences in several factors and therefore has geographic and temporal scales ove a broad range. At these small scales on the sea-surface interface itself, waves, winds water temperature and salinity, bubbles, spray and variations in the amount of sola radiation that reaches the ocean surface, and other factors, affect the transfer o properties and energy.
In the long term, the convergence and divergence of oceanic heat transport provid sources and sinks of heat for the atmosphere and partly shape the mean climate of th earth. Analyzing whether these processes are changing due to anthropogenic influence and the potential impact of these changes is the subject of this chapter. Followin guidance from the Ad Hoc Working Group of the Whole, much of the informatio presented here is based on or derives from the very thorough analysis conducted by th Intergovernmental Panel on Climate Change (IPCC) for its recent Fifth Assessmen Report (ARS).
The atmosphere and the ocean form a coupled system, exchanging at the air-se interface gases, water (and water vapour), particles, momentum and energy. Thes exchanges affect the biology, the chemistry and the physics of the ocean and influenc its biogeochemical processes, weather and climate (exchanges affecting the water cycl are addressed in Chapter 4).
From a biogeochemical point of view, gas and chemical exchanges between the ocean and the atmosphere are important to life processes. Half of the Global Net Primar Production of the world is by phytoplankton and other marine plants, uptaking CO2 an releasing oxygen (Field et al., 1998; Falkowski and Raven, 1997). Phytoplankton is
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therefore also responsible for half of the annual production of oxygen by plants and through the generation of organic matter, is at the basis of most marine food webs i the ocean. Oxygen production by plants is a critical ecosystem service that keep atmospheric oxygen from otherwise declining. However, in many regions of the ocean phytoplankton growth is limited by a deficit of iron in seawater. Most of the iro alleviating this limitation reaches the ocean through wind-borne dust from the desert of the world.
Gas and chemical exchanges between the atmosphere and ocean are also important t climate change processes. For example, marine phytoplankton produces dimethy sulphide (DMS), the most abundant biological sulphur compound emitted to th atmosphere (Kiene et al., 1996). DMS is oxidized in the marine atmosphere to for various sulphur-containing compounds, including sulphuric acid, which influence th formation of clouds. Through this interaction with cloud formation, the massiv production of atmospheric DMS over the ocean may have an impact on the earth' climate. The absorption of CO2 from the atmosphere at the sea surface is responsibl for the fundamental role of the ocean as a carbon sink (see section 3 below).
2. Heat flux and temperature
2.1 Sea-Surface Temperature
Sea-surface temperature (SST) has been measured in surface waters by a variety o methods that have changed significantly over time. Furthermore the spatial patterns o SST change are difficult to interpret. Nevertheless a robust trend emerges from thes historical series after careful inspection and analysis of the datasets. Figure 1 shows th historical SST trend instrumentally observed using the best datasets of spatiall interpolated products, contrasted against the 1961 — 1990 climatology. Changes in SS are reported in this section and in Chapter 2 of the IPCC (Hartmann et al., 2013).
The IPCC in ARS concluded that ‘recent’ warming (since the 1950s) is strongly evident i SST at all latitudes of each ocean. Prominent spatio-temporal structures, including the E Nifio Southern Oscillation (ENSO), decadal variability patterns in the Pacific Ocean, and  hemispheric asymmetry in the Atlantic Ocean, were highlighted as contributors to th regional differences in surface warming rates, which in turn affect atmospheri circulation (Hartmann et al., 2013).
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—— COBE ——ERSST HadISS —-—HadSST3 ve HadNMAT2
Temperature anomaly (°C)
1850 1900 1950 2000
Figure 1. Global annual average sea surface temperature (SST) and Night Marine Air Temperature (NMAT relative to a 1961-1990 climatology from state of the art data sets. Spatially interpolated products ar shown by solid lines; non-interpolated products by dashed lines. From Hartmann et al. 2013, Fig. 2.18.
“It is certain that global average sea surface temperatures (SSTs) have increased sinc the beginning of the 20th century. (...) Intercomparisons of new SST data record obtained by different measurement methods, including satellite data, have resulted i better understanding of uncertainties and biases in the records. Although thes innovations have helped highlight and quantify uncertainties and affect ou understanding of the character of changes since the mid-20" century, they do not alte the conclusion that global SSTs have increased both since the 1950s and since the lat 19" century.” (Hartmann et al., 2013).
2.2 Changes in sea-surface temperature (SST) as inferred from subsurfac measurements.
Upper ocean temperature (hence heat content) varies over multiple time scales including seasonal, interannual (e.g., associated with El Nifio), decadal and centennia (Rhein et al., 2013). Depth-averaged (0 to 700 m) ocean-temperature trends from 197 to 2010 are positive over most of the globe. The warming is more prominent in th Northern Hemisphere, especially in the North Atlantic. This result holds true in differen analyses, using different time periods, bias corrections and data sources (e.g., with o without XBT or MBT data’) (Rhein et al. 2013). Zonally averaged upper-ocea temperature trends show warming at nearly all latitudes and depths (Figure 2a) However, the greater volume of the Southern Hemisphere ocean increases th contribution of its warming to the global heat content (Rhein et al., 2013). Stronges warming is found closest to the sea surface, and the near-surface trends are consistent
* XBT are expendable bathythermographs, probes that using electronic solid-state transducers registe temperature and pressure while they free fall through the water column. MBT are their mechanica predecessors, that lowered on a wire suspended from a ship, used a metallic thermocouple as transducer © 2016 United Nations  

with independently measured SST (Hartmann et al., 2013). The global average warmin over this period is 0.11 [0.09 to 0.13] °C per decade in the upper 75 m, decreasing t 0.015°C per decade by 700 m (Figure 2c) (Rhein et al 2013).
The globally averaged temperature difference between the ocean surface and 200  increased by about 0.25°C from 1971 to 2010. This change, which corresponds to a 4 pe cent increase in density stratification, is widespread in all the oceans north of abou 40°S. Increased stratification will potentially diminish the exchanges between th interior and the surface layers of the ocean; this will limit, for example, the input o nutrients from below into the illuminated surface layer and of oxygen from above int the deeper layers. These changes might in turn result in reduced productivity an increased anoxic waters in many regions of the world ocean (Capotondi et al., 2012).
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70 80°S 60°S 40°S 20°S 0°S 20° (c) Latitude
1960 1970 1980 1990 2000 2010
(a,b) Temp. trend (°C per decade (c) Temp. anom. (°C o
= 0.1 6.7 -0. -0.25
6.3 -0.3
d  g 6.  3d 61
Year
The boundaries and names shown and the designations used on this map do not imply official endorsement or acceptance by the United Nations.
Figure 2. (a) Depth-averaged (0 to 700) m ocean-temperature trend for 1971-2010 (longitude vs. latitude colours and grey contours in degrees Celsius per decade); (b) Zonally averaged temperature trend (latitude vs. depth, colours and grey contours in degree Celsius per decade) for 1971-2010 with zonall averaged mean temperature over-plotted (black contours in degrees Celsius). Both North (25-652N) an South (south of 30°S), the zonally averaged warming signals extend to 700 m and are consistent wit poleward displacement of the mean temperature field. Zonally averaged upper-ocean temperature trend show warming at nearly all latitudes and depths (Figure 2 (b). A relative maximum in warming appear south of 30°S. (c) Globally averaged temperature anomaly (time vs. depth, colours and grey contours i degrees Celsius) relative to the 1971-2010 mean; (d) Globally averaged temperature difference betwee the ocean surface and 200 m depth (black: annual values, red: 5-year running mean). All panels ar constructed from an update of the annual analysis of Levitus et al. (2009). From Rhein et al. (2013) Fig 3.1.
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2.3 Upper Ocean Heat Content (UOHC)
The ocean’s large mass and high heat capacity allow it to store huge amounts of energy more than 1000 times that found in the atmosphere for an equivalent increase i temperature. The earth is absorbing more heat than it is emitting back into space, an nearly all this excess heat is entering the ocean and being stored there.
The upper ocean (0 to 700 m) heat content increased during the 40-year period fro 1971 to 2010. Published rates range from 74 TW to 137 TW (1 TW = 10” watts), whil an estimate of global upper (0 to 700 m depth) ocean heat content change, using ocea statistics to extrapolate to sparsely sampled regions and estimate uncertaintie (Domingues et al., 2008), gives a rate of increase of global upper ocean heat content o 137 TW (Rhein, et al. 2013). Warming of the ocean accounts for about 93 per cent of th increase in the Earth’s energy inventory between 1971 and 2010 (high confidence) Melting ice (including Arctic sea ice, ice sheets and glaciers) and warming of th continents and atmosphere account for the remainder of the change in energy (Rhein e al. 2013). Global integrals of 0 to 700 m upper ocean heat content (UOHC) (Figure 3. estimated from ocean temperature measurements all show a gain from 1971 to 201 (Rhein et al. 2013).
(a)
  o
0-700 m OHC (ZJ °
155 Levitu 5 Ishii  8 Domingue 9 Palme = Smith
 6
-100 T T 7 7  1950 1960 1970 1980 1990 2000 2010
Year
Figure 3. Observation-based estimates of annual global mean upper (0 to 700 m) ocean heat content in Z (1Z= 107 Joules) updated from (see legend): Levitus et al. (2012), Ishii and Kimoto (2009), Domingues e al. (2008), Palmer et al. (2009; O©American Meteorological Society. Used with permission.) and Smith an Murphy (2007). Uncertainties are shaded and plotted as published (at the one standard error level, othe than one standard deviation for Levitus, with no uncertainties provided for Smith). Estimates are shifte to align for 2006-2010, 5 years that are well measured by the ARGO Program of autonomous profilin floats, and then plotted relative to the resulting mean of all curves for 1971, the starting year for tren calculations.
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2.4 The ocean’s role in heat transport
Solar energy is unevenly distributed over the earth’s surface, leading to excess hea reaching the tropics and a heat deficit in latitudes poleward of about 40° in eac hemisphere. The heat balance, and therefore a relatively stable climate, is maintaine through the meridional redistribution, or flux, of heat by the atmosphere and the ocean Quantification and understanding of this heat content and its redistribution have bee achieved through diverse methods, including international programmes maintainin instrumented moorings, transoceanic lines of XBTs, satellite observations, numerica modelling and, more recently, the ARGO Program of autonomous profiling instrument (Abraham et al., 2013; von Schuckmann and Le Traon, 2011).
In the latitude band between 25°N and 25°S, the atmospheric and oceanic contribution to the meridional heat fluxes are similar, and the atmosphere dominates at highe latitudes. In the ocean, the heat flux is accomplished by contributions from the wind driven circulation in the upper ocean, by turbulent eddies, and by the Meridiona Overturning Circulation (MOC). The MOC is a component of ocean circulation that i driven by density contrasts, rather than by winds or tides, and one which exhibits  pronounced vertical component, with dense water sinking at high latitudes, offset b broadly distributed upwelling at lower ones. As distinct circulation patterns characteriz each of the ocean basins, their individual contributions to the meridional heat flux diffe significantly. Estimates indicate that, on a yearly average, the global oceans carry 1- PW (1PW=10"°W) of heat from the tropics to higher latitudes, with somewhat highe transports to the northern hemisphere (Fasullo and Trenberth, 2008).
Most of the heat excess due to increases in atmospheric greenhouse gases goes into th ocean (IPCC, 2013). Although all ocean basins have warmed during the last decades, th increase in heat content is not uniform; the increase in heat content in the Atlanti during the last four decades exceeds that of the Pacific and Indian Oceans combine (Levitus et al., 2009; Palmer and Haines, 2009). Enhanced northward heat flux in th subtropical South Atlantic, which includes heat driven from the subtropical Indian Ocea through the Agulhas Retroflection, may have contributed to the larger increase in hea content in the Atlantic Ocean compared with other basins (Abraham et al., 2013; Lee e al., 2011).
Numerical simulations also indicate that changes in ocean heat fluxes are the mai mechanism responsible for the observed temperature fluctuations in the subtropica and subpolar North Atlantic (Grist et al., 2010).
Meridional heat flux estimates inferred from the residual of heat content variation suggest that the heat transferred northward throughout the Atlantic is transferred t the atmosphere in the subtropical North Atlantic (Kelly et al., 2014). Observations fro the Rapid/Mocha instrument array at 26°N in the North Atlantic indicate that the mea Atlantic meridional heat flux at this latitude is 1.33 PW, with substantial variability du to changes in the strength of the MOC (Cunningham et al., 2007; Kanzow et al., 2007 Johns et al., 2011; McCarthy et al., 2012). Moreover, recent studies show tha interannual changes in the MOC (and the associated heat flux measured at 26°N) lead t temperature anomalies in the subtropical North Atlantic which, in turn, can have a
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strong impact on the northern hemisphere climate (Cunningham et al., 2013; Buchan e al., 2014).
2.5. Air-sea Heat fluxes
Heat uptake by the ocean can be substantially altered by natural oscillations in th earth’s ocean and atmosphere. The effects of these large-scale climate oscillations ar often felt around the world, leading to the rearrangement of wind and precipitatio patterns, which in turn substantially affect regional weather, sometimes wit devastating consequences.
The ENSO is the most prominent of these oscillations and is characterized by a anomalous warming and cooling of the central-eastern equatorial Pacific. The war phase is called El Nifio and the cold, La Nifia. During El Nifio events, a weakening of th Pacific trade winds decreases the upwelling of cold waters in the eastern equatoria Pacific and allows warm surface water that generally accumulates in the western Pacifi to flow east.
As a consequence, El Nifios release heat into the atmosphere, causing an increase i globally averaged air temperature. However, the “recharge oscillator theory” (Ren an Jin, 2013) indicates that a buildup of upper-ocean heat content is a necessar precondition for the development of El Nifio events. La Nifias are associated with  strengthening of the trade winds, which leads to a strong upwelling of cold subsurfac water in the eastern Pacific. In this case, the ocean uptake of heat from the atmospher is enhanced, causing the global average surface temperature to decrease (Roemmic and Gilson, 2011).
The cycling of ENSO between El Nifio and La Nifia is irregular. In some decades El Nifi has dominated and in other decades La Nifia has been more frequent, also seen i phase shifts of the Interdecadal Pacific Oscillation (Meehl et al., 2013), which is relate to build up and release of heat. A strengthening of the Pacific trade winds in the pas two decades has led to a more frequent occurrence of La Nifias (England et al., 2014) Consequently, the heat uptake by the subsurface ocean was enhanced, leading to  slowdown of the surface warming (Kosaka and Xie, 2013). This is one of the factor affecting the global mean temperature, expected to increase by 0.21°C per decade fro 1998 to 2012, but which instead warmed by just 0.04°C (the so-called recent warmin hiatus, IPCC, 2013). Although there are several hypotheses on the cause of the globa warming hiatus, the role of ocean circulation in this negative feedback is certain Drijfhout et al. (2014) have shown that the North Atlantic, Southern Ocean and Tropica Pacific all play significant roles in the ocean heat uptake associated with the warmin hiatus.
Chen and Tung (2014) analyzed the historical and recent record of sea surfac temperature and Ocean Heat Content (OHC), and found distinct patterns at the surfac and in deeper layers. On the surface, the patterns conform to the El Nifio/La Nifi patterns, with the Pacific Ocean playing a dominant role by releasing heat during an E Nifio (or capturing heat during La Nifia). At depth, the dominant pattern shows heating
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taking place in the Atlantic Ocean and in the Circumpolar Current region. Coinciding i time, changes in OHC could help to explain the observed slowdown in global warming. I is anticipated that the mechanisms involved may at some point reverse, releasing larg amounts of heat to the atmosphere and accelerating global warming (e.g., Levermann et al., 2012).
Many other naturally occurring ocean-atmosphere oscillations in the Pacific, Atlantic and Indian Oceans have also been recognized and named. The ENSO as a globa phenomenon, has an expression in the Atlantic basin called the Atlantic Nifio. In the las six decades, this mode has weakened, leading to a warming of the equatorial easter Atlantic of up to 1.5°C (Tokinaga and Xie, 2011). Although the role of the Atlantic Nifi on the global heat budget is not significant, this Atlantic warming trend has led to a increase in precipitation over the equatorial Amazon, Northeast South America Equatorial West Africa and the Guinea coast, and a decrease in rainfall over the Sahe (Gianinni et al., 2003; Tokinaga and Xie, 2011; Marengo et al., 2011; Rodrigues et al. 2011). Moreover, recent studies have shown that the Atlantic Nifio can have an effec on ENSO (Rodriguez-Fonseca et al., 2009; Keenlyside et al., 2013).
In the Indian Ocean, the dominant basin-wide oscillation is the Indian Dipole Mode (Saj et al., 1999). A positive phase is characterized by cool surface-temperature anomalies i the eastern Indian Ocean, warm-temperature anomalies in the western Indian Ocean and easterly wind-stress anomalies along the equator. Similarly to ENSO, meridiona heat transport and the associated buildup of upper-ocean heat content are a possibl precondition for the development of the Indian Ocean Dipole event (McPhaden an Nagura, 2014). The warm surface temperatures in the western Indian Ocean ar associated with an increase in subsurface heat content and vice-versa for the east (Fen et al. 2001; Rao et al., 2002). This zonal contrast of ocean heat content is induced b anomalies of zonal wind along the equator and the resulting variability in zonal mas and heat transport (Nagura and McPhaden 2010). The warm surface temperatures i the western Indian Ocean are associated with an increase in subsurface heat conten and vice-versa for the east; the positive dipole causes above-average rainfall in easter Africa and droughts in Indonesia and Australia (Behera et al., 2005; Yamagata et al. 2004; Ummenhofer et al., 2009; Cai et al., 2011; Section 5 below). Although th phenomena discussed here are global, many of the most significant impacts are on th coastal environment (see following Section).
2.6 Environmental, economic and social impacts of changes in ocean temperatur and of major ocean temperature events
Coastal waters are valuable both ecologically and economically because they support  high level of biodiversity. They act as nursery areas for many commercially importan fish species, and are the marine areas most accessible to the public. Because inshor habitats are shallow, water temperatures in coastal areas are closely linked to th regional climate and its seasonal and long-term fluctuations. Coastal waters also hos some of the most vulnerable marine habitats, because they are intensively exploited b (including, but not limited to) the fishing industry and recreational craft, and because of
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their proximity to outlets of pollution, such as rivers and sewage outfalls. Coasta development and the threat of rising sea level may also impinge upon these valuabl habitats (Halpern et al., 2008). Ecological degradation can lower the socio-economi value of coastal regions, with negative impacts on commercial fisheries, aquacultur facilities, damage to coastal infrastructure, problems with power-station cooling, an exert a dampening effect on coastal tourism from degraded ecological services.
It has been recently shown that when compared with estimates for the global ocean decadal rates of SST change are higher at the coast. During the last three decades approximately 70 per cent of the world’s coastline has experienced significant increase in SST (Lima and Wethey, 2012). This has been accompanied by an increase in th number of yearly extremely hot days along 38 per cent of the world’s coastline, an warming has been occurring significantly earlier in the year along approximately 36 pe cent of the world’s temperate coastal areas (defined as those between latitudes 30° an 60° in both hemispheres) at an average rate of 6.1 + 3.2 days per decade (Lima an Wethey, 2012).
The warming of coastal waters can have many serious consequences for the ecologica system (Harley et al., 2006). This can include changes in the distribution of importan commercial fish and shellfish species, particularly the movement of species to highe latitudes due to thermal stress (Perry et al., 2005). Warming of coastal waters also ca lead to more favourable conditions for many organisms, among them marine invasiv species that can devastate commercial fisheries and destroy marine ecosyste dynamics (Occhipinti-Ambrogi, 2007). Water quality might also be impacted by highe temperatures that can increase the severity of local outbreaks by pathogenic bacteria o the occurrence of Harmful Algal Blooms (HABs). These in turn would cause harm t seafood, consumers and marine organisms (Bresnan et al., 2013). Increased coral ree bleaching and mortality from warming seas (combined with ocean acidification, see nex sections) will lead to the loss of important marine habitats and associated biodiversity.
Changes in ocean temperatures have global impacts. As ocean temperatures warm species that prefer specific temperature ranges may relocate — as has been observed for instance, in copepod assemblages in the North Atlantic (Hays et al., 2005). Som organisms, like corals, are sedentary and cannot relocate with changing temperatures. I the water becomes too warm, they may experience a bleaching event. Higher sea leve and warmer ocean temperatures can alter ocean circulation and current flow an increase the frequency and intensity of storms, leading to changes in the habitat o many species worldwide.
Changes in ocean temperatures affect not only marine ecosystems, but also the climat over land, with devastating economic and social implications. Many natural oceani oscillations are known to have an impact on (terrestrial) climate, but these oscillation and the response of the climate to them are also changing during recent decades. Fo instance, an El Nifio phase of ENSO (see previous Section for more details on ENSO displaces great amounts of warm water from the western to the eastern Pacific, leadin to more evaporation over the latter. As a consequence, western and southern Sout America and parts of North America experience wetter conditions. At the same time Australia, Brazil, India, Indonesia, the Philippines, parts of Africa and the United State © 2016 United Nations 1 

of America suffer droughts. La Nifia events usually cause the opposite patterns However, in the last several decades, ENSO events have changed their spatial an temporal characteristics (Yeh et al., 2009; McPhaden, 2012).
During recent decades, the warm waters of El Nifio events have been displaced to th central Pacific instead of to the eastern Pacific. It is not clear yet whether these change are linked to anthropogenic climate change or natural variability (Yeh et al., 2011). I any case, the effects on climate of an ENSO event centred in the central Pacific (a centra Pacific ENSO) are in sharp contrast to that associated with one centred in the easter Pacific.
For instance, northeastern and southeastern Australia experience a reduction in rainfal during the eastern Pacific El Nifios and there is a decrease in rainfall over northwester and northern Australia during central Pacific events (Taschetto and England, 2009 Taschetto et al., 2009). The Indian monsoon fails during eastern Pacific El Nifios, but i enhanced during central Pacific El Nifios (Kumar et al., 2006). Over the semi-arid regio of northeast Brazil, eastern Pacific El Niftios/La Nifias cause dry/wet conditions; centra Pacific El Nifios have the opposite effect, with the worst drought in the last 50 year associated with the strong 2011/12 La Nifia and not with El Nifios as in the pas (Rodrigues et al., 2011; Rodrigues and McPhaden, 2014). This drought caused th displacement of 10 million people and economic losses on the order of 3 billion Unite States dollars in relation to agriculture and cattle raising alone. In contrast to drought i Brazil, the 2011/12 La Nifia caused floods across southeastern Australia.
In other ocean basins, changes in oceanic oscillations and temperatures have also ha an impact on climate. For instance, in the Indian Ocean, a positive phase of the India Dipole Mode (warm/cold temperatures in the western/eastern equatorial Indian Ocean leads to flooding in east Africa and droughts in Indonesia, Australia, and India (Saji et al. 1999; Ashok et al., 2001; Gadgil et al., 2004; Yamagata et al., 2004; Behera et al., 2005 Ummenhofer et al., 2009; Cai et al., 2011). The counterpart of ENSO in the Atlanti (Atlantic Nifio) has weakened during the last six decades, leading to an increase in SST i the eastern equatorial Atlantic. As a consequence, rainfall has been enhanced over th equatorial Amazon and West Africa (Tokinaga and Xie, 2011). On the other hand, a unusual warming of the tropical North Atlantic in 2005 was responsible for one of th worst droughts in the Amazon River basin and a record Atlantic hurricane season Hurricanes Rita and Katrina caused the loss of almost 2000 lives and an estimate economic toll of 150 billion —135 billion US dollars from Katrina and 15 billion U dollars from Rita. (http://www.datacenterresearch.org/data-resources/katrina/facts for-impact/). Anomalous warm conditions also occurred in the tropical North Atlantic i 2010 leading to two once-in-a-century droughts in less than five years in the Amazo River basin (Marengo et al., 2011).
Ocean warming will stress species both through thermic changes in their environmenta envelope and through increased interspecies competition. These shifts become all th more important in shelf seas once they reach terrestrial boundaries, i.e., the shiftin species runs out of shelf. For example, changes in the coastal currents in south-easter Australia cause changes to primary production through to fisheries productivity. Thi then feeds through to local and regional socio-economic impacts (Suthers et al., 2011).
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The IPCC ARS concluded that “it is unlikely that annual numbers of tropical storms hurricanes and major hurricanes counts have increased over the past 100 years in th North Atlantic basin. Evidence, however, is for a virtually certain increase in th frequency and intensity of the strongest tropical cyclones since the 1970s in that region (Hartmann et al. 2013, Section 2.6.3). Moreover, the IPCC ARS states that “it is difficul to draw firm conclusions with respect to the confidence levels associated with observe trends prior to the satellite era and in ocean basins outside of the North Atlantic (Hartmann et al. 2013, Section 2.6.3). Although a strong scientific consensus on th matter does not exist, there is some evidence supporting the hypothesis that globa warming might lead to fewer but more intense tropical cyclones globally (Knutson et al. 2010). Evidence exists that the observed expansion of the tropics since approximatel 1979 is accompanied by a pronounced poleward migration of the latitude at which th maximum intensities of storms occur at a rate of 1° of latitude per decade (Kossin et al. 2014; Hartmann et al., 2013; Seidel et al., 2008). If this trend is confirmed, it woul increase the frequency of events in coastal areas that are not exposed regularly to th dangers caused by cyclones. Hurricane Sandy in 2012 may be an example of thi (Woodruff et al., 2013).
3. Water flux and salinity
3.1 Regional patterns of salinity, and changes in salinity” and freshwater content
The ocean plays a pivotal role in the global water cycle: about 85 per cent of th evaporation and 77 per cent of the precipitation occur over the ocean (Schmitt, 2008) The horizontal salinity distribution of the upper ocean largely reflects this exchange o freshwater: high surface salinity is generally found in regions where evaporatio exceeds precipitation, and low salinity is found in regions of excess precipitation an runoff. Ocean circulation also affects the regional distribution of surface salinity.
The Earth’s water cycle involves evaporation and precipitation of moisture at the Earth’ surface. Changes in the atmosphere’s water vapour content provide strong evidenc that the water cycle is already responding to a warming climate. Further evidenc comes from changes in the distribution of ocean salinity (Rhein et al. 2013; FAQ. 3.2) Diagnosis and understanding of ocean salinity trends are also important, becaus salinity changes, like temperature changes, affect circulation and stratification, an therefore the ocean’s capacity to store heat and carbon as well as to change biologica productivity.
Seawater contains both salt and fresh water, and its salinity is a function of the weigh of dissolved salts it contains. Because the total amount of salt does not change over
2 ‘Salinity’ refers to the weight of dissolved salts in a kilogram of seawater. Because the total amount o salt in the ocean does not change, the salinity of seawater can be changed only by addition or removal o fresh water.
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human time scales, seawater’s salinity can only be altered—over days or centuries—b the addition or removal of fresh water.
The water cycle is expected to intensify in a warmer climate. Observations since th 1970s show increases in surface and lower atmospheric water vapour (Figure 4a), at  rate consistent with observed warming. Moreover, evaporation and precipitation ar projected to intensify in a warmer climate. Recorded changes in ocean salinity in th last 50 years support that projection (Rhein et al. 2013; FAQ. 3.2).
The atmosphere connects the ocean’s regions of net fresh water loss to those of fres water gain by moving evaporated water vapour from one place to another. Th distribution of salinity at the ocean surface largely reflects the spatial pattern o evaporation minus precipitation (Figure 4b), runoff from land, and sea ice processes There is some shifting of the patterns relative to each other, because of the ocean’ currents. Ocean salinity acts as a sensitive and effective rain gauge over the ocean. I naturally reflects and smoothes out the difference between water gained by the ocea from precipitation, and water lost by the ocean through evaporation, both of which ar very patchy and episodic (Rhein et al. 2013; FAQ. 3.2). Data from the past 50 years sho widespread salinity changes in the upper ocean, which are indicative of systemati changes in precipitation and runoff minus evaporation.
(Figure 4b). Subtropical waters are highly saline, because evaporation exceeds rainfall whereas seawater at high latitudes and in the tropics—where more rain falls tha evaporates—is less so. The Atlantic, the saltiest ocean basin, loses more freshwate through evaporation than it gains from precipitation, while the Pacific is nearly neutral i.e., precipitation gain nearly balances evaporation loss, and the Southern Ocean i dominated by precipitation. (Figure 4b; Rhein et al. 2013; FAQ. 3.2). Changes in surfac salinity and in the upper ocean have reinforced the mean salinity pattern (4c). Th evaporation-dominated subtropical regions have become saltier, while th precipitation-dominated subpolar and tropical regions have become fresher. Whe changes over the top 500 m are considered, the evaporation-dominated Atlantic ha become saltier, while the nearly neutral Pacific and precipitation-dominated Souther Ocean have become fresher (Figure 4d; Rhein et al. 2013; FAQ. 3.2).
Observed surface salinity changes also suggest a change in the global water cycle ha occurred (Chapter 4). The long-term trends show a strong positive correlation betwee the mean climate of the surface salinity and the temporal changes in surface salinit from 1950 to 2000. This correlation shows an enhancement of the climatological salinit pattern: fresh areas have become fresher and salty areas saltier.
Ocean salinity is also affected by water runoff from the continents, and by the meltin and freezing of sea ice or floating glacial ice. Fresh water added by melting ice on lan will change global-averaged salinity, but changes to date are too small to observe (Rhei et al. 2013; FAQ. 3.2).
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+6 (a) Trend i og (Otal precipitabl water vapou °° (4988-2010 -0. -1. (kg m* per decade)
too () Mea evaporatio o minu precipitatio -100
(cm yr’)
E Se 0.8 (c) Trend i 04 surface salinit 0.0 (1950-2000 -0.4
-08
(PSS78 per decade)
a7 (d) Mea a surface salinity
The boundaries and names shown and the designations used on this map do not imply official endorsement or acceptance by the United Nations.
Figure 4. Changes in sea surface salinity are related to the atmospheric patterns of evaporation minu precipitation (E — P) and trends in total precipitable water: (a) Linear trend (1988-2010) in tota precipitable water (water vapour integrated from the Earth’s surface up through the entire atmosphere (kg m-2 per decade) from satellite observations (Special Sensor Microwave Imager) (after Wentz et al. 2007) (blues: wetter; yellows: drier). (b) The 1979-2005 climatological mean net E —P (cm yr—1) fro meteorological reanalysis (National Centers for Environmental Prediction/National Center fo Atmospheric Research; Kalnay et al., 1996) (reds: net evaporation; blues: net precipitation). (c) Tren (1950-2000) in surface salinity (PSS78 per 50 years) (after Durack and Wiljffels, 2010) (blues freshening yellows-reds saltier). (d) The climatological-mean surface salinity (PSS78) (blues: <35; yellows—reds: >35) From Rhein et al. 2013; FAQ. 3.2. Fig 1.
In conclusion, according to the last IPCC ARS, “It is very likely that regional trends hav enhanced the mean geographical contrasts in sea surface salinity since the 1950s: saline
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surface waters in the evaporation-dominated mid-latitudes have become more saline while relatively fresh surface waters in rainfall-dominated tropical and polar region have become fresher” (Stocker et al., 2013). “The mean contrast between high- an low-salinity regions increased by 0.13 [0.08 to 0.17] from 1950 to 2008. It is very likel that the inter-basin contrast in freshwater content has increased: the Atlantic ha become saltier and the Pacific and Southern Oceans have freshened. Although simila conclusions were reached in AR4, recent studies based on expanded data sets and ne analysis approaches provide high confidence in this assessment” (Stocker et al., 2013) “The spatial patterns of the salinity trends, mean salinity and the mean distribution o evaporation minus precipitation are all similar. These similarities provide indirec evidence that the pattern of evaporation minus precipitation over the oceans has bee enhanced since the 1950s (medium confidence)” Stocker et al., (2013). “Uncertainties i currently available surface fluxes prevent the flux products from being reliably used t identify trends in the regional or global distribution of evaporation or precipitation ove the oceans on the time scale of the observed salinity changes since the 1950s” (Stocke et al., 2013).
4. Carbon dioxide flux and ocean acidification
4.1 Carbon dioxide emissions from anthropogenic activities
Since the start of Industrial Revolution, human activities have been releasing larg amounts of carbon dioxide into the atmosphere. As a result, atmospheric CO. ha increased from a glacial to interglacial cycle of 180-280 ppm to about 395 ppm in 201 (Dlugokencky and Tans, 2014). Until around 1920, the primary source of carbon dioxid to the atmosphere was from deforestation and other land-use change activities (Ciais e al., 2013). Since the end of World War II, anthropogenic emissions of CO2 have bee increasing steadily. Data from 2004 to 2013 show that human activities (fossil fue combustion and cement production) are now responsible for about 91 per cent of th total CO2 emissions (Le Quéré et al. 2014).
CO, emissions from fossil fuel consumption can be estimated from the energy data tha are available from the United Nations Statistics Division and the BP Annual Energ Review. Data in 2013 suggests that about 43 per cent of the anthropogenic CO emissions were produced from coal, 33 per cent from oil and 18 per cent from gas, an 6 per cent from cement production (Figure 5).
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Growth rates
Data: CDIAC/GCP 2012-201 16- Coal 3.0   8 12- Oil 1.4  © to  8  a 6 Gas 1.4   oO 2- Cement 4.7%
1960 1970 1980 1990 2000 2010
Figure 5. CO emissions from different sources from 1958 to 2013 (Le Quéré et al. 2014).
Coal is an important and, recently, growing proportion of CO2 emissions from fossil fue combustion. From 2012 to 2013, CO, emissions from coal increased 3.0 per cent compared to the increase rate of 1.4 per cent for oil and gas (Le Quéré et al. 2014). Coa accounted for about 60 per cent of the CO2 emission growth in the same period. This i largely because many large economies of the world have recently resorted to using coa as an energy source for a wide variety of industrial processes, instead of oil, gas an other energy sources.
4.2 The ocean as a sink for atmospheric CO2
The global oceans serve as a major sink of atmospheric CO2. The oceans take up carbo dioxide through mainly two processes: physical air-sea flux of atmospheric CO. at th ocean surface, the so called “solubility pump” and through the active biological uptak of CO, into the biomass and skeletons of plankters the so-called “biological pump” Colder water can take up CO2 more than warm water, and if this cold, denser wate sinks to form intermediate, deep, or bottom water, there is transport of carbon awa from the surface ocean and thus from the atmosphere into the ocean interior. Thi "solubility pump" helps to keep the surface waters of the ocean on average lower in C than the deep water, a condition that promotes the flux of the gas from the atmospher into the ocean.
Phytoplankton take up CO2 from the water in the process of photosynthesis, some o which sinks to the bottom in the form of particles or is mixed into the deeper waters a dissolved organic or inorganic carbon. Part of this carbon is permanently buried in th sediments and other part enters into the slower circulation of the deep ocean. Thi "biological pump" serves to maintain the gradient in CO, concentration between th surface and deep waters.
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Depth
B 6 5 500 5 4 ) 4 £1000 3 & 3 2 1500 2 1 2000 S 1 c   a Anthropogeni CO S 1000  & 1000 (umo! kg" a
2 Latitude
0
The boundaries and names shown and the designations used on this map do not imply official endorsement or acceptance by the United Nations.
Figure 6. Anthropogenic CO, distributions along representative meridional sections in the Atlantic, Pacific and Indian oceans for the mid-1990s (Sabine et al. 2004).
Because the ocean mixes slowly, about half of the anthropogenic CO} (Cant) stored in th ocean is found in the upper 10 per cent of the ocean (Figure 6.). On average, th penetration depth is about 1000 meters and about 50 per cent of the anthropogeni CO, in the ocean is shallower than 400 meters.
Globally, the ocean shows large spatial variations in terms of its role as a sink o atmospheric CO, (Takahashi et al. 2009). Over the past 200 years the oceans hav absorbed 525 billion tons of CO2 from the atmosphere, or nearly half of the fossil fue emissions over the period (Feely et al. 2009). The oceanic sink of atmospheric CO, ha increased from 4.0 + 1.8 GtCO, (1 GtCO, = 10° tons of carbon dioxide) per year in th 1960s to 9.5 + 1.8 GtCO2 per year during 2004-2013. During the same period, th estimated annual atmospheric CO, captured by the ocean was 2.6 +0.5 Gt of CO compared with around 1.9 Gt of CO, during the sixties (Le Queré et al., 2014). However due to the decreased buffering capacity, caused by this CO2 uptake, the proportion o anthropogenic carbon dioxide that goes into the ocean has been decreasing.
Estimates of the global inventory of anthropogenic carbon, Cant (including marginal seas have a mean value of 118 PgC and a range of 93 to 137 PgC in 1994 and a mean of 16 PgC and range of 134 to 186 PgC in 2010 (Rhein et al 2013). When combined with mode results Khatiwala et al. (2013) arrive at a “best” estimate of the global ocean inventor (including marginal seas) of anthropogenic carbon from 1750 to 2010 of 155 PgC with a uncertainty of +20 per cent (Rhein et al 2013).
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The storage rate of anthropogenic CO, is assessed by calculating the change in Can concentrations between two time periods. Regional observations of the storage rate ar in general agreement with that expected from the increase in atmospheric CO concentrations and with the tracer-based estimates. However, there are significan spatial and temporal variations in the degree to which the inventory of Cant track changes in the atmosphere (Figure 7, Rhein et al 2013).
Indian Ocean
Atlantic Ocean (mol m? y" 24
Pacific Ocean (mol m? y* - 0.8
(mol m? y* 0.9
WS oth 25°W 92 25 ois %e
Le gsGw 7557.2 POS OR ye
The boundaries and names shown and the designations used on this map do not imply official endorsement or acceptance by the United Nations.
Figure 7. Maps of storage rate distribution of C,,¢ in (mol m? yr?) averaged over 1980-2005 for the thre ocean basins (left to right: Atlantic, Pacific and Indian Ocean). From Khatiwala et al 2009, a slightl different colour scale is used for each basin.
Comprehensive evaluation of available data shows that in the context of the globa carbon cycle, it is only the ocean that has acted as a net sink of carbon from th atmosphere. The land was a source early in the industrial age, and since about 1950 ha trended toward a sink, but it is not yet clearly a net sink. (Ciais et al. 2013 and Khatiwal et al. 2009, Khatiwala et al. 2013). Latest data from 2004 to 2013 show that the globa oceans take up about one-fourth (26 per cent, Le Quéré, 2014) of the total annua anthropogenic emissions of CO2. This is a very important physical and ecological servic that the ocean has performed in the past and performs today, that underpins al strategies to mitigate the negative impacts of global warming.
4.3 Ocean acidification
As already seen in the previous section, the global oceans serve as an important sink o atmospheric CO, effectively slowing down global climate change. However, this benefi comes with a steep bio-ecological cost. When CO) reacts with water, it forms carboni acid, which then dissociates and produces hydrogen ions. The extra hydrogen ion consume carbonate ions (CO;”) to form bicarbonate (HCO3). In this process, the pH an concentrations of carbonate ions (CO3”) are decreasing. As a result, the carbonat mineral saturation states are also decreasing. Due to the increasing acidity, this proces is commonly referred to as “ocean acidification (OA)”. According to the IPCC AR 4 and 5 “Ocean acidification refers to a reduction in pH of the ocean over an extended period typically decades or longer, caused primarily by the uptake of carbon dioxide (CO) fro the atmosphere.” (...)” Anthropogenic ocean acidification refers to the component of p reduction that is caused by human activity” (Rhein et al. 2013).
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Although the average oceanic pH can vary on interglacial time scales, the changes ar usually on the order of “0.002 units per 100 years; however, the current observed rat of change is “0.1 units per 100 years, or roughly 50 times faster. Regional factors, suc as coastal upwelling, changes in riverine and glacial discharge rates, and sea-ice los have created “OA hotspots” where changes are occurring at even faster rates. Althoug OA is a global phenomenon that will likely have far-reaching implications for man marine organisms, some areas will be affected sooner and to a greater degree.
Recent observations show that one such area in particular is the cold, highly productiv region of the sub-arctic Pacific and western Arctic Ocean, where unique biogeochemica processes create an environment that is both sensitive and particularly susceptible t accelerated reductions in pH and carbonate mineral concentrations. The O phenomenon can cause waters to become undersaturated in carbonate minerals an thereby affect extensive and diverse populations of marine calcifiers.
4.4 The CO2 problem
Based on the most recent data of 2004 to 2013, 35.7 GtCO, (1 GtCO, = 10° tons o carbon dioxide) of anthropogenic CO2 are released into the atmosphere every year (L Quéré et al. 2014). Of this, approximately 32.4 GtCO2 come directly from the burning o fossil fuels and other industrial processes that emit CO. The remaining 3.3 GtCO, ar due to changes in land-use practices, such as deforestation and urbanization. Of thi 35.7 GtCO, of anthropogenically produced CO, emitted annually, approximately 10. GtCO, (or 29 per cent) are incorporated into terrestrial plant matter. Another 15. GtCO, (or 46 per cent) are retained in the atmosphere, which has led to some planetar warming. The remaining 9.5 GtCO> (or 26 per cent) are absorbed by the world’s ocean (Le Quéré et al. 2014).
As the hydrogen ions produced by the increased CO? dissolution take carbonate ions ou of seawater, the rate of calcification of shell-building organisms is affected; they ar confronted with additional physiological challenges to maintain their shells. Althoug alteration of the carbonate equilibrium system in the ocean reducing carbonate io concentration, and saturation states of calcium carbonate minerals will play a rol imposing an additional energy cost to calcifier organisms, such as corals an shellbearing plankton, this is by no means the sole impact of OA.
4.5 What are the impacts of a more acidic ocean?
Throughout the last 25 million years, the average pH of the ocean has remained fairl constant between 8.0 and 8.2. However, in the last three decades, a fast drop ha begun to occur, and if CO emissions are left unchecked, the average pH could fall belo 7.8 by the end of this century (Rhein, et al. 2013).
This is well outside the range of pH change of any other time in recent geologica history. Calcifying organisms in particular, such as corals, crabs, clams, oysters and th tiny free-swimming pteropods that form calcium carbonate shells, could be particularl vulnerable, especially during the larval stage. Many of the processes that cause OA hav © 2016 United Nations 1 

long been recognized, but the ecological implications of the associated chemica changes have only recently been investigated. OA may have important ecological an socioeconomic consequences by impacting directly the physiology of all organisms i the ocean.
The altered environment is imposing an extra energy cost for the acid-base regulation o their internal body milieu. Through biological and evolutionary adaptation this proces might have a huge variation of expression among different types of organisms, a subjec that only recently has become the focus of intense scientific research.
Calcification is an internal process that in its vast majority does not depends directly o seawater carbonate content, since most organism use bicarbonate, that is increasin under acidification scenarios, or CO2 originating in their internal metabolism. It has bee demonstrated in the laboratory and in the field that some calcifiers can compensate an thrive in acidification conditions.
OA is not a simple phenomenon nor will it have a simple unidirectional effect o organisms. The abundance and composition of species may be changed, due to OA wit the potential to affect ecosystem function at all trophic levels, and consequentia changes in ocean chemistry could occur as well. Some species may also be better abl than others to adapt to changing pH levels due to their exposure to environment where pH naturally varies over a wide range. However, at this point, it is still ver uncertain what the ecological and societal consequences will be from any potentia losses of keystone species.
4.6 Socioeconomic impacts of ocean acidification
Some examples of economic disruptions due to OA have been reported. The mos visible case is the harvest failure in the oyster hatcheries along the Pacific Northwes coast of the USA. Hatcheries that supply the majority of the oyster spat to farms nearl went out of business as they unknowingly pumped low pH water, apparently corrosiv to oyster larvae, into their operation. Although intense upwelling that could hav brought low oxygen water to hatcheries might also be a factor in these massiv mortalities, low pH, “corrosive water” tends to recur seasonally in this region Innovations and interactions with scientists allowed these hatcheries to monitor th presence of corrosive incoming waters and adopt preventive measures.
Economic studies have shown that potential losses at local and regional scales may hav negative impacts for communities and national economies that depend on fisheries. Fo example, Cooley and Doney (2009) using data from 2007, found that of the 4 billio dollars in annual domestic sales, Alaska and the New England states likely to be affecte by hotspots of OA, contributed the most at 1.5 billion dollars and 750 million dollars respectively. These numbers clearly show that any disruption in the commercia fisheries in these regions due to OA could have a cascading effect on the local as well a on the national economy.
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