Changing Ocean, Marine Ecosystems, and Dependent Communities

As AR5 concluded, the ocean is warming as a direct result of anthropogenic changes to the radiative properties of the atmosphere and the heat budget of the Earth (very likely) (Bindoff et al., 2013²²). Over the past few decades our ocean observing system has measured an increase in ocean temperature (Figure 5.1). This temperature increase corresponds to an uptake of over 90% of the excess heat accumulated in the Earth system over this period (Bindoff et al., 2013²³; Rhein et al., 2013²⁴). This heat in the ocean also causes it to expand and has contributed about 43% of the observed global mean SLR from 1970–2015 (Section 4.2.2.3.6). 

Since AR5, there have been further improvements in our ability to understand and correct instrumental errors and new estimates also attempt to minimise biases in estimating temperature changes arising from traditional data-void filling strategies (Abraham et al., 2013²⁵; Durack, 2015²⁶; Cheng and Chen, 2017²⁷; Cheng et al., 2017²⁸). New estimates from ocean observations of ocean heat uptake in the top 2000 m between 1993 and 2017 very likely range from 9.2 ± 2.3 ZJ yr-1 to 12.1 ± 3.1 ZJ yr-1 (Johnson et al., 2018²⁹) ⁴. Three recent independent estimates do a better job of accounting for instrumental biases and the sparseness of historical ocean temperature measurements than the older studies assessed in AR5, and provide larger and more consistent estimates of heat uptake rates for the 0-2000 m layer of 5.8 ± 1.0 ZJ yr-1 (Cheng and Chen, 2017³⁰; Cheng et al., 2017³¹; Ishii et al., 2017³²), 6.0 ± 0.8 ZJ yr-1 (updated from Domingues et al. (2008)) and 6.3 ± 1.8 ZJ yr-1 (Cheng and Chen, 2017³³; Cheng et al., 2017³⁴; Ishii et al., 2017³⁵) for the 1971–2010 period assessed by AR5. 

Based on these new published methods and revised atlases we update the estimates for ocean heat uptake (Table 5.1, and SM5.1). For all of the periods assessed in Table 5.1, it is virtually certain that the upper ocean (0–700 m) has warmed. These results are consistent with earlier research into the duration of record needed to detect a significant signal in global ocean heat content (Gleckler et al., 2012³⁶). Critically, the high confidence and high agreement in the ocean temperature data means we can detect discernable rates of increase in ocean heat uptake (Gleckler et al., 2012³⁷; Cheng et al., 2019³⁸). The rate of heat uptake in the upper ocean (0–700 m) is very likely higher in the 1993–2017 (or 2005–2017) period compared with the 1969–1993 period (see Table 5.1). The deeper layer (700–2000 m) heat uptake rate is likely to be higher in the 1993–2017 period compared with the 1969–1993 period.

The direct comparison of the observed changes in ocean heat content and the simulated historical changes is undertaken to detect climate change, to attribute the causes of climate change to the forcings in the system, and to evaluate the performance of ESMs. Attribution studies also reject competing hypotheses to explain the global ocean changes such as natural forcing from solar variability or volcanic eruptions (see Section 1.3) (Bindoff et al., 2013³⁹). Detection and attribution studies have since been used to detect changes in the rate of ocean heat uptake and to attribute these changes to human activity (Gleckler et al., 2016⁴⁰). 

Updated observationally-based estimates of ocean heat uptake are consistent with simulations of equivalent time-periods from an ensemble of CMIP5 ESMs (Table 5.1 and the inset panel in Figure 5.1) (high confidence), once the limitations of the historical ocean observing network and the internally generated variability with a single realisation of the real world are taken into account (see Section 5.2.2.2). Following the CMIP5 protocol, the ESMs are radiatively forced with observationally derived estimates of greenhouse gas concentrations and aerosols, including natural forcing variations from volcanic eruptions and solar forcing, through 2005; after 2006 each of the ESMs uses either the RCP2.6 or RCP8.5 emissions scenarios. The very likely ranges of the observed trends of heat uptake for the four periods and two layers all fall within the very likely range of simulated heat uptake from the ESM ensemble (Table 5.1). The difference between observations and average of the simulations in the upper ocean is an overestimate of heat uptake by about 20% and for the deeper layer there an underestimate by a similar amount, but this difference is still well within the very likely range from the ensemble of simulations. The overall consistency between observationally-based estimates and ESM simulations of the historical period gives greater confidence in the projections; it is very likely that historical simulations agree with observations of the global ocean heat uptake (Table 5.1).

While the collection of the worlds’ ESMs have been criticised for having an ensemble mean that does not exhibit the observed ‘hiatus’ or ‘slowdown’ of global mean surface temperature increase in the early 21st century (Meehl et al., 2011⁵⁷; Trenberth et al., 2016⁵⁸) , it is increasingly clear that this is at least in part due to the redistribution of heat within the climate system from the surface into the interior ocean and between ocean basins. Individual realisations of ESMs do show decades with slow increases in mean surface temperature change comparable to what was observed, even though these cases exhibit continued interior ocean heat uptake, and every ensemble member exhibits surface warming closer to the ensemble-mean over multi-decadal timescales (Meehl et al., 2011⁵⁹; England et al., 2015⁶⁰; Knutson et al., 2016⁶¹)

The ensemble average of the CMIP5 ESMs projects widespread ocean warming over the coming century, concentrated in the upper ocean (Figures 5.2c and 5.3) (Kuhlbrodt and Gregory, 2012). The anthropogenic heat will penetrate into the ocean following well-established circulation pathways (Jones et al., 2016a⁶³). The greatest vertically integrated heat uptake occurs where there is already the formation of interior waters, such as Antarctic Intermediate Water along the Antarctic Circumpolar Current (Frölicher et al., 2015⁶⁴) or NADW precursors in the Nordic Seas (Figure 5.2c), but all water-masses ⁵ that are subducted over decades are 

expected to experience significant warming (see Figure 5.3). The warming in the subtropical gyres penetrates deeper into the ocean than other gyres (roughly 15°N–45°N and 15°S–45°S in Figure 5.3), following the wind-driven bowing down of the density surfaces (the solid lines in Figure 5.3) in these gyres (Terada and Minobe, 2018⁶⁵). The greater warming at 700-2000 m in the Atlantic than the Pacific or Indian Oceans (Figure 5.3) reflects the strong southward transport of recently formed NADW at these depths by the AMOC. Two areas that commonly exhibit substantially reduced near-surface warming over the course of the 21st century are the northern north Atlantic, where a slowing AMOC (see Section 6.7.1.1) reduces the northward heat transport and brings the surface temperatures closer to what is found in other ocean basins at these latitudes (Collins et al., 2013⁶⁶), and the southern side of the Southern Ocean, where water upwells that has been submerged for so long that it has not yet experienced significant anthropogenic climate change (Armour et al., 2016⁶⁷). Most of these projected warming patterns are broadly consistent across the current and previous generations of climate models (Mitchell et al., 1995⁶⁸; Collins et al., 2014⁶⁹) as well as observations and theoretical understanding. These multiple lines of evidence give high confidence that the projections describe the changes in the real world (high agreement, robust evidence).

The near surface salinity of the ocean is both observed and projected to evolve in ways that reflect the increased intensity of the Earth’s hydrologic cycle (Durack, 2015⁷⁰) and the increasing near-surface ocean stratification (Zika et al., 2018⁷¹). As described in WGI AR5, the ocean surface in areas that currently have net evaporation are expected to become saltier, while areas with net precipitation are expected to get fresher (Rhein et al., 2013⁷²), as the patterns of precipitation and evaporation are generally expected to be amplified (Held and Soden, 2006⁷³). At longer time-scales of decades, the larger scale changes in the ocean circulation and basin integrated freshwater imbalances emerge in the near-surface salinity changes, as shown in Figure 5.3b, with an increasingly salty tropical and subtropical Atlantic and Mediterranean contrasting with a freshening Pacific and polar Arctic emerging as robust signals across the suite of ESMs (Collins et al., 2013⁷⁴). The freshening of the high latitudes in the north Atlantic and Arctic basin is consistent with the widely expected weakening of the AMOC (also discussed in Section 6.7), hydrological cycle changes and a decline in the volume of sea ice (discussed in Section 3.2.2).

Projected salinity changes in the subsurface ocean reflect changes in the rates of formation of water masses or their newly formed properties (Purich et al., 2018⁷⁵). Thus, projected freshening of the Southern Ocean surface leads to a freshening of the Antarctic Intermediate Water that is subducted there, flowing northward from the Southern Ocean as a relatively fresh water-mass at depths of 500–1500 m (Figure 5.3b). Increased surface salinity in the Atlantic subtropical gyres are pumped into the interior by the winds, leading to an increased salinity of the interior subtropical gyres, along with contributions from increasingly salty Mediterranean water (Jordà et al., 2017⁷⁶). Conversely, freshwater capping of the northwestern north Atlantic is projected to inhibit deep convection in the Labrador Sea and the consequent production of Labrador Sea Water in some models (Collins et al., 2013⁷⁷), and contributes to the increased salinity of the north Atlantic between 1000–2000 m depths (Figure 5.3b). 

Identifying the specific patterns of anthropogenic climate changes in oceanic observations is complicated by the presence of basin-scale natural variability with timescales ranging from tidal to multi-decadal, and due to the difficulties associated with maintaining high-precision observing systems spanning the ocean basins and limited observational coverage of the extratropical Southern Hemisphere before 2006 (Rhein et al., 2013⁷⁸). Inferences based on oceanographic observations from the 1970s onward show wide-spread warming of the upper 700 m (Figure 5.2a), in broad agreement with the ensemble of historical CMIP5 ESM simulations (Figure 5.2b). These ESMs indicate that anthropogenic regional warming over the past half-century should be discernable at the 95% confidence level in much of the upper oceans (un-stippled areas in Figure 5.2b). Most of the areas where observational analyses (Figure 5.2a) exhibit long-term cooling are either regions where the internally generated variability is large enough to mask the trends (e.g., the Eastern Tropical Pacific, Northwest Atlantic, and Kurushio extension east of Japan, which are stippled in Figure 5.2b), or where the observational coverage early in the record is limited and different analyses can disagree about trends (e.g., the Southern Ocean and extratropical South Pacific). When internal variability is taken into account, the broad consistency in the magnitude and regional distribution of observed and simulated 50-year trends gives confidence to the ESM projections of longer-term oceanic changes described previously.

Detailed regional patterns of trends in temperature and heat content at depths of 0–2000 m during the early 21st century are consistent in various analysis, owing to the improved observing network (Roemmich et al., 2015⁸¹; Desbruyères et al., 2016a⁸²) (Figure 5.2d). At depths of 700–2000 m, observations in all of the ocean basins show broadly warming trends in the well-observed Argo era (2006 to present), with particularly significant warming patterns in the Southern Hemisphere extratropics around 40oS and the subpolar north Atlantic (Figure 5.3a). These observed changes support the notion that deep ocean heat content has been continuously increasing. As a result, regional climate change signatures emerge from confounding natural variability sooner in the 700–2000 m depth range than in upper 700 m of the ocean, where interannual modes of variability have a larger influence on the circulation (for a more complete discussion see Johnson et al. (2018). Despite regional patches of cooling water in the upper 700 m (Figure 5.2d), every one of the world’s ocean basins volume averaged over depths of 0–2000 m has experienced significant warming over the last decade (Figure 5.3, and also Desbruyères et al. (2016a)⁸²). The greatest warming of the top 2000 m has been in the Southern Ocean (Roemmich et al., 2015⁸³; Trenberth et al., 2016⁸⁴), the tropical and subtropical Pacific Ocean (Roemmich et al., 2015⁸⁵), and the tropical and subtropical Atlantic Ocean (Cheng and Chen, 2017⁸⁶). The Southern Hemisphere extratropical oceans accounted for 67–98% of the total ocean heat increase in the uppermost 2000 m for the period of 2006–2013 (Roemmich et al., 2015⁸⁷). Shi et al. (2018)⁸⁸ suggest that the dominant ocean heat uptake by the Southern Hemisphere in the early 21st century is expected to become more balanced between the hemispheres as the asymmetric cooling by aerosols decreases.

Large-scale patterns of natural variability at interannual to decadal time scales can mask the long-term warming trend in the upper 700 m, particularly in the tropical Pacific and Indian Oceans (England et al., 2014⁸⁹; Liu et al., 2016⁹⁰) and in the north Atlantic (Buckley and Marshall, 2015⁹¹). The most significant upper 700 m warming between five-year averages centered on 2007–2015 occurred in a large extratropical band of the Southern Hemisphere between 30ºS–60ºS, and in the tropical Indian Ocean, the eastern North Pacific and western subtropical north Atlantic (Figure 5.2d). Warming of the southern hemisphere subtropical gyres is driven, in part, by an intensification of Southern Ocean winds in recent decades, facilitating the penetration of heat to deeper depths (Gao et al., 2018⁹²). Marginal seas, such as the Mediterranean and Red seas have also exhibited notable warming. Conversely, over this timeframe there were also regions of cooling in the upper 700 m, notably in the north Atlantic around 40oN–60oN and in the western tropical Pacific (Figure 5.2d). Recent relatively cold and fresh surface and subsurface conditions in the north Atlantic have been attributed to anomalous atmospheric forcing (Josey et al., 2018⁹³) or weakened transport by the north Atlantic Current and AMOC (Smeed et al., 2018⁹⁴), and in turn may have contributed to an intensification of deep convection in the Labrador Sea since 2012 (Yashayaev and Loder, 2017⁹⁵). All these observed decadal changes can be related to internal decadal variability (Robson et al., 2014⁹⁶; Yeager et al., 2015⁹⁷) even though they resemble expected longer-term anthropogenically forced trends. Substantial decadal-scale warming and cooling trends in the tropical Pacific and Indian oceans can arise from natural El Niño-Southern Oscillation (ENSO) and Indian Ocean Dipole variability (Han et al., 2014⁹⁸). Large ensembles of freely running CMIP5 ESM simulations also show that internal variability can dominate the regional manifestation of the anthropogenic climate signal on decadal timescales (Kay et al., 2014⁹⁹). This is illustrated by the differing warming trends in Figure 5.2e and 5.2f from two identical ESMs that differ only in the weather in their 1850 initial conditions, averaged over the whole 21st century, by contrast, the ensemble of CMIP5 models project statistically significant anthropogenic regional upper 700 m heat content trends almost everywhere (Figure 5.2c).

There are well documented changes in observed ocean temperatures and salinities (Abraham et al., 2013¹⁰⁰; Ishii et al., 2017¹⁰¹). However, attributing these changes in the state of the ocean to anthropogenic causes can be challenging due to the presence of internally generated variability, which can swamp the underlying climate change signal in short records and on regional scales. As can be seen in Figure 5.2, the observed long-term trends (Figure 5.2a) exhibit a striking similarity to the CMIP5 ensemble mean in areas where the models suggest that anthropogenic changes should be statistically significant (Figure 5.2b). However, the trends in the shorter well-observed period covering 2005–2017 (Figure 5.2d) exhibits strong trends from internal variability, as illustrated by the differences of two ensemble members of the same ESM with the same forcing but initialised with different weather (Figure 5.2e and 5.2f). Detection and Attribution studies take the internal variability into account and separate the underlying climate signals with the same spatio-temporal sampling as the observations, and apply a range of statistical tests to determine the coherence of the observations with the co-sampled observations (Bindoff et al., 2013¹⁰²; AR5 WG1 Box 10.1). 

Since AR5, the use of different and updated oceanographic data sets and increase in the number of ensembles of the CMIP5 simulations (Kay et al., 2014¹⁰³) has improved the overall detection and attribution of human influence. Together these measures increase the coherence of the simulations and reduce noise. For example, an isotherm approach used to reduce the noise from the displacement of isotherms in the upper water column allowing detection in each of the mid-latitude ocean basins was achieved on 60-year time series (Weller et al., 2016¹⁰⁴). Using all the available ocean temperature and salinity profiles from the Southern Ocean, Swart et al. (2018) show that the warming and freshening patterns were consistent primarily with increased human induced greenhouse gases and secondarily from ozone depletion in the stratosphere, but inconsistent with internal variability. Together the evidence from the AR5, and the discussion above with the new evidence on regional scales across the global oceans, we conclude that the observed long-term upper ocean temperature changes are very likely to have a substantial contribution from anthropogenic forcing.

The wind-driven ocean circulation at the end of the 21st century is expected to be qualitatively similar to that in the present day, even as important buoyancy-loss driven overturning circulations are expected to weaken. ESM projections suggest that some major ocean current transports will exhibit a modest increase (such as the Kuroshio Extension (Terada and Minobe, 2018) or a small decrease such as for the Indonesian Throughflow (Sen Gupta et al., 2016); many predominantly wind-driven current-system transports are expected to exhibit smaller than 20% changes by 2100 with RCP8.5 forcing. Climate-change induced changes of the circulation in other mid-latitude basins may be difficult to detect or reliably project because of significant natural variability at inter-annual (e.g., El Niño) to decadal (e.g., the Pacific Decadal Oscillation) timescales. The Antarctic Circumpolar Current is projected to be subject to strengthening westerly winds and substantially reduced rates of Antarctic Bottom Water (AABW) formation, as assessed in the Cross-Chapter Box 7 in Chapter 3. The heat transported by the buoyancy-loss driven AMOC, in particular, contributes to the relatively clement climate of northern Europe and the north Atlantic Basin as a whole, although the wind-driven ocean gyres also contribute to the meridional ocean heat transport (see the review by Buckley and Marshall (2015). As a result, there is a concern that significant changes in ocean circulation could lead to localised climate changes that are much larger than the global mean. Projected and observed changes in the AMOC and the rates of formation of deep water-masses in the north Atlantic are discussed in Chapter 6.7.1, along with the possibility of abrupt or enduring changes resulting from forcing by Greenlandic meltwater. A significant reduction in AMOC would, in turn, modestly weaken the Gulf Stream transport, which also has a substantial wind driven component (Frajka-Williams et al., 2016¹⁰⁵). Most aspects of the large-scale wind-driven ocean circulation are very likely to be qualitatively similar to the circulation in the present day, with only modest changes in transports and current location. 

The global ocean below 2000 m has warmed significantly between the 1980s and 2010s (Figure 5.4), contributing to ocean heat uptake and through thermal expansion to SLR (Purkey and Johnson, 2010¹⁰⁶; Desbruyères et al., 2016b¹⁰⁷). The observed deep warming rate varies regionally and by depth reflecting differences in the waters influencing particular regions. The deep and abyssal north Atlantic, fed by North Atlantic Deep Water (NADW), has reversed from warming to cooling over the past decade, possibly associated with the North Atlantic Oscillation (NAO) (e.g., Yashayaev, 2007; Desbruyères et al., 201¹⁰⁸4) or longer-term weakening in north Atlantic overturning circulation (Caesar et al., 2018¹⁰⁹; Thornalley et al., 2018¹¹⁰). The strongest warming is observed in regions of the deep ocean AABW (Purkey et al., 2014¹¹¹). Regions of the ocean fed by AABW from the Weddell Sea have exhibited a possible slowdown in local AABW warming rates (Lyman and Johnson, 2014¹¹²), while the Pacific, fed by AABW from the shelves along the Ross and Adelie Coast, has continued to warm at an accelerating rate between 1990 and 2018 (Desbruyères et al., 2016b¹¹³). 

To date, assessment of deep ocean (below 2000 m) heat content has mostly been from ship-based data collected along decadal repeats of oceanographic transects (Figure 5.4b) (Talley et al., 2016¹¹⁴). While relatively sparse in space and time compared to the upper ocean, these transects were positioned to optimise sampling of most deep ocean basins and provide the highest quality of salinity, temperature and pressure data. Argo floats capable of sampling to 6000 m have just started to populate select deep ocean basins; this Deep Argo data has just started providing regional deep ocean warming estimates (Johnson et al., 2019¹¹⁵). Decadal monitoring by the full global Deep Argo array (Johnson et al., 2015¹¹⁶), complemented by indirect estimates from space (Llovel et al., 2014¹¹⁷; Von Schuckmann et al., 2014), will strongly reduce the currently large uncertainties of deep ocean heat content change estimates in the future.

The spatial and temporal sparseness of observations below 4000 m, along with significant differences between various ESMs, limits our understanding of the exact mechanisms driving the abyssal ocean variability. However, ESMs consistently predict an anthropogenic climate-change induced long-term abyssal warming trend originating in the Southern Ocean due to a reduction in the formation rates of cold AABW (Heuzé et al., 2015¹¹⁸). Although the abyssal modes of natural variability are not as pronounced as closer to the surface, deep ocean heat content can vary on relatively short time scales through the communication of topographic and planetary waves driven by changes in the rate of deep water formation at high latitudes (Kawase, 1987¹¹⁹; Masuda et al., 2010; Spence et al., 2017). AABW has shown variability in properties and production rates over the past half century (Purkey and Johnson, 2013¹²¹; Menezes et al., 2017¹²²). A slowdown in AABW formation rates may arise from freshening of shelf waters, changes in local winds driving cross shelf mixing, or larger scale dynamics controlling the spin up or down of Southern Ocean gyres influencing the density of outflowing waters over deep sills. Large-scale circulation changes can also alter the properties of the ambient water that is entrained as dense water descends along the Antarctic continental slopes (Spence et al., 2017¹²³). Evolving AABW properties may also reflect changes in deep Southern Ocean convection. The Weddell Polynya is a large opening in the wintertime ice of the Weddell Sea that is kept ice-free despite intense heat loss to the atmosphere by convective mixing bringing up warm and salty water from the deep ocean. (See Box 3.2 for a more extensive discussion of polynyas and the Weddell Polynya in particular). The Weddell Polynya was present in three of the first years of infrared satellite observations of wintertime sea ice concentrations in the mid-1970s, but it has been closed since 1976, only to reopen in 2016 and 2017. The prominent Weddell Polynya in the mid-1970s greatly increased the volume of the coldest waters in the deep Weddell Sea. Weddell Polynyas are documented to drive abyssal cold and salty signals and can spread thermal signals as waves further and faster than could be explained by slow advective signals (Martin et al., 2015¹²⁴; Zanowski and Hallberg, 2017¹²⁵); these waves do not directly heat individual water parcels, but instead warm the ocean where they cause the coldest deep layers to spread laterally and thin. However, recovery from the large Weddell polynya of the early 1970s can only explain about 20% of the observed abyssal warming trend (Zanowski et al., 2015¹²⁶).

The ocean’s properties are changing most rapidly in the near surface waters that are more immediately exposed to atmospheric forcing. As a result of the surface-intensified warming, the upper few hundred meters of the ocean are becoming more stably stratified (Helm et al., 2011¹²⁶; Talley et al., 2016¹²⁸). The combination of surface intensified warming and near-surface freshening at high latitudes leading to a projection of more intense near-surface stratification (the downward-increasing vertical gradient of density) across all ocean basins (Figures 5.3 and 5.5) is a robust result with a high agreement across successive generations of coupled climate models (Capotondi et al., 2012¹²⁹; Bopp et al., 2013¹³⁰). Based on the projected changes from individual models between 1986–2005 and 2081–2100, the mean stratification of the upper 200 m averaged between 60°S–60°N, normalised by the ensemble mean value from 1986–2005 will very likely increase by between 1.0–9.3% (with 95% confidence and a CMIP5 median change of 2.6%) for RCP2.6, and by between 12.2–30.0% (median value 21.2%) for RCP8.5. Inferences from oceanic observations (Good et al., 2013¹³¹) suggest that the 20-year mean stratification averaged between 60°S–60°N and over the top 200 m very likely increased by between 2.18–2.42% from 1971–1990 to 1998–2017. By contrast, the bottom intensified warming in the abyss (see Figure 5.4) which is consistent with a slowing in the rate of AABW formation, is also associated with a reduction in the abyssal stratification of the ocean (Lyman and Johnson, 2014¹³²; Desbruyères et al., 2016b¹³³). Both of these changes have consequences for the evolving turbulence and ocean water-mass structure. Based on observational evidence, theoretical understanding and robust ESM projections, it is very likely that stratification in the upper few hundred meters of the ocean below the mixed layer will increase significantly in the 21st century over most ocean basins as a result of climate change, and abyssal stratification will likely decrease.

Many dynamical consequences of increased stratification are understood with very high confidence (see, for instance, Gill (1982) and Vallis (2017)). For the same turbulent kinetic energy dissipation, locally increased stratification reduces the turbulent vertical diffusivity of heat, salinity, oxygen and nutrients (see Section 5.2.2.2.4). Increased stratification in the tropics and subtropical gyres will likely lead to a net reduction in the vertical diffusivities of nutrients and other gases within the main thermocline, reducing the flux of nutrients into the euphotic zone and increasing the gradient in oxygen concentrations between the near surface ocean and the interior. Increasing upper ocean stratification (Figure 5.5) acts to restrict the depth of the ocean’s surface mixed layer. Increasing stratification increases the buoyancy frequency and the lateral propagation speed of internal gravity waves and boundary waves by about half the percentage change of the stratification itself. Increasing stratification increases both the length of the internal deformation radius (a typical length scale in baroclinic eddy dynamics) and the horizontal scales of internal tides (see Section 5.2.2.2.3) proportionately with the changes in the internal gravity wave speeds. An increase in stratification will increase the lateral propagation of internal Rossby waves (which set up the basin-scale ocean density structure) proportionately. For the same forcing, increasing stratification reduces the geostrophically balanced slope of density surfaces, and hence the vertical extent of basin-scale wind-driven gyres or coastal upwelling circulations. The flattening of density surfaces by increased stratification inhibits advective exchange between the surface and interior ocean (Wang et al., 2015a¹³⁴), with consequences for the uptake of anthropogenic carbon (Section 5.2.2.3), the evolving oxygen distribution (Section 5.2.2.4) and the supply of nutrients to support primary production (Section 5.2.2.5).

Coastal systems are subject to the same large-scale warming trends as the open ocean, but the local response may be dominated by a complex of localised changes in factors such as circulation, mixing, river plumes or the seasonal upwelling of cold water. Using ESMs to project how these factors will interact often requires much finer resolution than is currently affordable in global models, however regional high-resolution models can be effective, especially in marginal seas like the Mediterranean with restricted interactions with the open ocean and that respond primarily to local forcing (Adloff et al., 2015¹³⁵). High resolution regional models have also been used to project robust localised ocean climate changes in wide shelf seas with more extensive interactions with the open ocean, like those in northwestern Europe (Tinker et al., 2016¹³⁶). The technical difficulties of using nested regional models are much greater in coasts adjacent to energetic large-scale currents like the Gulf Stream, Kuroshio, and Agulhas, and projecting detailed coastal climate change such places may require the use of expensive high resolution global models (Saba et al., 2016¹³⁷). These physical coastal changes have consequences that cascade through ecosystems to people, as is illustrated in detail for eastern boundary upwelling systems in Box 5.2.

Both human structures and ecological systems in the coastal zone are directly impacted by tidal amplitudes, which contribute to high-water levels and the tidal flushing rates of estuaries, embayments, marshes and mangroves. The tides are the response of a forced-damped-resonance system (Arbic et al., 2009¹³⁸). The M2 tide is the dominant tidal constituent in most places, with a period of half a lunar day, or 12 hours, 25 minutes; the M2 tides are created by the differential motion of the solid Earth and oceans in response to the gravitational attraction of the moon (Newton, 1687¹³⁹; Laplace, 1799¹⁴⁰). The astronomical forcing evolves only slowly, however the tidal damping and basin resonance at tidal frequencies can change in response to changes in sea level, stratification and coastal conditions (Müller, 2012¹⁴¹; Schindelegger et al., 2018¹⁴²). Several recent studies have analysed historical coastal tide gauge data and found amplitude trends of order 1–4% per century (Ray, 2009; Woodworth, 2010; Müller et al., 2011). In some locations, the changes in the tides have been of comparable importance to changes in mean sea level for explaining changes in high water levels (Jay, 2009). For many individual tide gauges, the trends in tidal amplitude are strongly positively or negatively correlated with local time-mean sea level trends (Devlin et al., 2017¹⁴⁵). Another source of secular tidal changes, changes in oceanic stratification, modifies the rate of energy conversion from the barotropic tides to the internal tides (Jayne and St. Laurent, 2001¹⁴⁶), the vertical profile of turbulent viscosity on shelves (Müller, 2012¹⁴⁷), and the propagation speed of the internal tides (Zhao, 2016¹⁴⁸). For example, Colosi and Munk (2006) found an increase in the amplitude of the principal lunar semidiurnal tide M2 in Honolulu of about 1 cm over the past 100 years, which they attributed primarily to changes in oceanic stratification bringing about local changes in relative phases of the internal and external M2 tides, increasing constructive interference. Both sea level and stratification are expected to exhibit robust secular positive trends in the coming century due to climate change, at rates that are significantly larger than historical trends, and people may choose to replace natural beaches and marshes with sea-walls in response to rising sea levels. As a result, it is very likely that the majority of coastal regions will experience statistically significant changes in tidal amplitudes over the course of the 21st century.

Because coastal tides are near resonance in many locations, small changes in sea level and bay shape can change the local tides significantly. For example, the insertion of tidal power plants can have a significant impact on the local tides (Ward et al., 2012¹⁴⁹). Various observational and modeling studies demonstrate that SLR has spatial heterogeneous impacts on the tides, with some locations experiencing decreased tidal amplitudes and others experiencing increased tidal amplitudes (Pickering et al., 2012¹⁵⁰; Devlin et al., 2017¹⁵¹; Pickering et al., 2017¹⁵²). Projections of tidal changes indicate that the patterns and even the sign of changes in tidal amplitudes depend on whether the coastlines are allowed to recede with rising sea levels or are held in place (Pickering et al., 2017¹⁵³; Schindelegger et al., 2018¹⁵⁴). Pelling et al. (2013)¹⁵⁵ and Hwang et al. (2014)¹⁵⁶ demonstrate that the rapid coastline changes in China’s Bohai Sea have already altered the tides in that region and throughout the Yellow Sea (Hwang et al., 2014¹⁵⁷). Pelling and Green (2014) examine the impact of flood defenses as well as SLR on tides on the European Shelf. Such tidal changes have implications for designing flood defenses, for tidal renewable energy, for tidal flushing timescales of estuaries and embayments, and for navigational dredging requirements (Pickering et al., 2012¹⁵⁸) (Section 5.4.2). The sign and amplitude of local changes to tides are very likely to be impacted by both human coastal adaptation measures and climate drivers (listed above).

ESMs are able to capture the dynamics of the climate system, but all numerical models have approximations and biases. The most commonly used type of ocean component in ESMs is known to exhibit numerically induced vertical mixing that can be a significant fraction of the physical mixing (Ilıcak et al., 2012¹⁵⁹; Megann, 2018¹⁶⁰). Because so many ocean models exhibit the same sign of bias, there is a systematic warming of the lower-main thermocline that is not cancelled out when taking the average over the ensemble of all the models in CMIP5. These biases are widely known within the ocean modelling community, and various groups are working to reduce these biases in future ESMs with better ocean model numerics and parameterisations. To correct for model biases, ESM projections are always taken as the difference from a control run without the anomalous forcing. However, some aspects of the ocean response to climate change are nonlinear, and model biases can introduce uncertainties into climate projections. In the case of heat uptake, this is of the order of 10% uncertainty, while for the rate of steric SLR (which depends on the nonlinear equation of state of seawater) the uncertainty in CMIP5 models is of the order of 20% (Hallberg et al., 2012¹⁶¹). 

Mesoscale eddies (geostrophic rotating vortices with spatial scales of 10–100 km that penetrate deeply into the water column, and are often described as the ocean’s weather) play an important role in regulating the changes to the larger scale ocean circulation, especially in the Antarctic Circumpolar current, as is discussed in Cross Chapter Box 7. In addition, sub-mesoscale eddies (rotationally influenced motions with smaller horizontal scales of hundreds of metres to about 10 km and intrinsic timescales of a few days that especially arise in association with fronts in the ocean’s surface properties) are known to be particularly important in the dynamics of the near-surface ocean boundary layer (see the review by Mahadevan (2016)). Sub-mesoscale instabilities are associated with re-stratifying overturning circulations that can limit the thickness of the well-mixed ocean surface boundary layer near fronts (Bachman et al., 2017¹⁶²). Moreover, sub-mesoscale motions generate strong vertical velocities that drive fluxes of nutrients from the interior ocean into the euphotic zone or create pockets of reduced mixing with increased phytoplankton residency time within the euphotic zone (Lévy et al., 2012¹⁶³). Intense mesoscale eddies are known to create favourable conditions for sub-mesoscale instabilities as shown in both observational (Bachman et al., 2017¹⁶⁴) and numerical studies (Brannigan et al., 2017¹⁶⁵). Intensifying Southern Ocean eddy fields will have a significant local impact on biological productivity, ecosystem structure, and carbon uptake, both directly and via sub-mesoscale processes. At typical CMIP5 ESM resolutions, it is only in the tropics that mesoscale eddies are adequately resolved to explicitly model their effects (Hallberg, 2013¹⁶⁶), while sub-mesoscale eddies are not resolved anywhere, so eddy effects need to be parameterised in ESMs. Despite great progress over the past 30 years in parameterising eddy effects, uncertainties in these parameterisations and how eddies will respond to novel conditions continue to contribute to uncertainties in projections of oceanic climate change (medium confidence).

Ocean turbulent mixing is a key process regulating the ocean circulation and climate. Turbulent mixing is important for the uptake and redistribution of heat, carbon, nutrients, oxygen and other tracers (properties that are carried along with the flow of water) in the ocean (Schmittner et al., 2009¹⁶⁷; MacKinnon et al., 2017¹⁶⁸). Both observations and theory indicate that turbulent mixing in the ocean is not constant in space or time. Global estimates of both the turbulent kinetic energy dissipation rate and the vertical diffusivity, two measures of ocean turbulence, vary over several orders of magnitude throughout the ocean (Figure 5.6) (Polzin et al., 1997¹⁷⁰; Waterman et al., 2012¹⁷¹; Whalen et al., 2012¹⁷²; Alford et al., 2013¹⁷³; Hummels et al., 2013¹⁷⁴; Sheen et al., 2013¹⁷⁵; Waterhouse et al., 2014¹⁷⁶; Kunze, 2017¹⁷⁷). For a given energy dissipation rate, the turbulent diffusivities of heat, salinity, nutrients and other tracers tend to be smaller with stronger stratification. This dependency on stratification helps explain why the observationally inferred diffusivity in the heavily stratified main thermocline (250–1000 m depth) is of similar magnitude to those deeper in the water column, while the turbulent energy density and dissipation rate are much stronger at the shallower depths (Whalen et al., 2012¹⁷⁸). Oceanic turbulence also fluctuates in time, is modulated by tidal cycles (Klymak et al., 2008¹⁷⁹), the mesoscale eddy field and seasonal changes (Whalen et al., 2018¹⁸⁰). In the mixed layer and directly below, turbulence changes according to local conditions, such as the winds, heating rates and local stratification (Sloyan et al., 2010¹⁸¹; Moum et al., 2013¹⁸²; D’Asaro, 2014¹⁸³; Tanaka et al., 2015¹⁸⁴) at diurnal to seasonal and longer timescales. These variations in near-surface turbulence need to be taken into account for ESMs to reproduce more accurately the observed seasonal cycle of surface properties and spatial structure of the depth of the thermally well-mixed near surface layer of the ocean. The spatial and temporal patterns of ocean turbulence help shape ocean tracer distributions (heat, dissolved greenhouse gases and nutrients) and how they will evolve in a changing climate (high confidence). 

Ocean turbulent mixing requires energy sources, many of which are expected to change with a changing climate. Surface wind and buoyancy forcing, the mean and eddying larger-scale ocean circulation itself, and the barotropic tides are all thought to be significant sources of the energy that drives mixing (Wunsch and Ferrari, 2004¹⁸⁵). Often this energy first passes through the ocean’s pervasive field of internal gravity waves that propagate and refract through the varying ocean circulation, often breaking into turbulent mixing far from their sources (Eden and Olbers, 2014¹⁸⁶; Alford et al., 2016¹⁸⁷; Melet et al., 2016¹⁸⁸; Meyer et al., 2016¹⁸⁹; Zhao et al., 2016b¹⁹⁰). The energy contributing to the internal waves from the winds and the subsequent turbulence will be altered by changes in tropical storm activity or sea ice coverage. For example, the increasing extent of ice-free Arctic Ocean has already been observed to lead to increased wind-driven internal waves (Dosser and Rainville, 2016¹⁹¹). The Southern Annular Mode is expected to intensify as a result of climate change (Young et al., 2011¹⁹²; Jones et al., 2016b¹⁹³), bringing with it stronger winds, and more wind-energy input over most of the Southern Ocean and a more intense mesoscale eddy field (Hogg et al., 2015¹⁹⁴). Changes in the near-bottom stratification will alter the rate that the barotropic tides generate internal waves, thereby altering the strength and distribution of the tidally generated mixing. Some of the parameterisations of interior ocean mixing used in CMIP5 ESMs take some changing turbulent energy sources into account (Jayne and St. Laurent, 2001¹⁹⁵), and more comprehensive mixing treatments are being developed for use in future generations of ESMs (Eden and Olbers, 2014¹⁹⁶). However, not all of the physical processes leading to the rich structure of mixing shown in Figure 5.6 are well understood or included in ESMs; the prospect of significant changes in the patterns and intensity of ocean turbulent mixing is a potential source of uncertainty (probably at the 10% level) in projections of physical and ecological changes in the ocean, including heat uptake, stratification changes, steric SLR, deoxygenisation and nutrient fluxes (medium confidence).

The analyses of the steadily growing number of surface ocean CO2 observations (now more than 20 million observations, SOCATv6 (www.socat.info/index.php/2018/06/19/v6-release) demonstrate that the net ocean uptake of CO2 from the atmosphere has increased from around 1.2 ± 0.5 Pg C yr-1 in the early 1980s to 2.0 ± 0.5 Pg C yr-1 in the years 2010–2015 (Rödenbeck et al., 2014; Landschützer et al., 2016). Once new estimates of the outgassing flux stemming from river derived carbon of 0.8 Pg C yr-1 (Resplandy et al. 2018) are accounted for, these new observations imply that the rate of global ocean uptake of anthropogenic CO2 increased from 2.0 ± 0.5 Pg C yr-1 to 2.8 ± 0.5 Pg C yr-1 between the early 1980s and 2010–2015 (Rödenbeck et al., 2014; Landschützer et al., 2016; Le Quéré et al., 2018). This increase is supported by the current generation of ocean carbon cycle models (Le Quéré et al., 2018), and commensurate with the increase in atmospheric CO2. 

The continuing efforts to re-measure dissolved inorganic carbon (DIC) along many of the repeat hydrographic lines that were occupied during the 1980s and 1990 (Talley et al., 2016), alongside the preparation of a global quality controlled database of ocean interior observations (Olsen et al., 2016a), have led to progress since AR5 regarding to the oceanic interior storage of anthropogenic CO2. Several studies analysed the changes in the amount of anthropogenic CO2 that have accumulated between different occupations in the different ocean basins (Wanninkhof et al., 2010; Pérez et al., 2013; Woosley et al., 2016; Carter et al., 2017), confirming that the anthropogenic CO2 taken up from the atmosphere is transported to depth, where most of it is stored. Using a newly developed reconstruction method, Gruber et al. (2019) extended these results to the globe. They find that between 1994 and 2007, across two standard deviations, that the global ocean has accumulated an additional 30-38 Pg C of anthropogenic CO2, which is equivalent to an air-sea CO2 flux of between 2.3–2.9 Pg C yr-1 (coherent with surface ocean CO2 observations), bringing the total inventory for the year 2007 to 150 ± 20 Pg C. Extrapolating this estimate to the year 2010 gives an inventory of 158 ± 18 Pg C, which is statistically indistinguishable from the ‘best’ estimate provided by Khatiwala et al. (2013) of 155 ± 31 Pg C and more recently also found from a steady-state ocean model (DeVries, 2014) for this reference year. If the inventory-based estimates are adjusted for the loss of natural carbon, a very likely total increase in storage between 1994 and 2007 of 24–34 Pg C, or around 25% of total emissions, is found (Gruber, 2019).

Thus, there is very high confidence from surface ocean and ocean interior carbon data that the strength of the ocean sink for anthropogenic carbon has increased in the last two decades in response to the growth of atmospheric CO2. Multiple lines of evidence indicate that it is very likely that the ocean has taken up 20–30% of the global emissions of CO2 from the burning of fossil fuels, cement production, and land-use change since the mid 1980s. The consistency between independent surface ocean observations and the ocean interior data-based reconstructions supports the assessment of very high confidence and provides robust evidence that fraction of emissions taken up by the ocean has not changed in a statistically significant manner in the last few decades and remains consistent with AR5.

Alongside a globally integrated perspective, these new surface ocean observations also reveal a substantial degree of variability at interannual and decadal scales (Rödenbeck et al., 2015; Landschützer et al., 2016; Le Quéré et al., 2018). Most notable are the air-sea CO2 flux variations in the tropics linked to ENSO variations (Rödenbeck et al., 2015; Landschützer et al., 2016), as well as the strong decadal variations in the high latitudes, especially the Southern Ocean (Landschützer et al., 2015; Munro et al., 2015; Ritter et al., 2017), discussed further in Chapter 3 (Section 3.2.1.2.4). Fluctuations in the Southern Ocean CO2 flux are important as they impart a substantial imprint also on the global uptake fluxes. For instance, reduced Southern Ocean uptake in the 1990–2000 period coincided with an exceptionally weak global net uptake of only about 0.8 ± 0.5 Pg C yr −1.

Thus, there is growing evidence from multiple datasets that the ocean carbon sink exhibits decadal variability at regional scales that significantly alter the globally integrated sink (medium confidence).

Detailed analyses of the spatial structure of the change in storage of anthropogenic CO2 confirm the variable nature of the ocean carbon sink suggested by the surface observations (Pérez et al., 2013), which are most likely a consequence of changes in ocean circulation (DeVries and Weber, 2017). The increase in anthropogenic CO2 between 1994 and 2007 occurs throughout the upper 1000 m, but with very different penetration depths, reflecting largely differences in the efficiency, with which the anthropogenic CO2 is transported from the surface to depth (Gruber et al., 2019) (Figure 5.7). This spatial distribution of how the amount of anthropogenic CO2 has changed between 1994 and 2007 is similar to the distribution of anthropogenic CO2 reconstructed for 1994 (Sabine et al., 2004), although the imprint of regional variations in ocean circulation and transport are discernible (Gruber, 2019).

Analyses of direct measurements of ocean chemistry from time series stations and merged shipboard studies show consistent decreases in surface-ocean pH over the past few decades. Reductions range between 0.013–0.03 pH units decade-1 over records that span up to 25 years (Table SM5.3). Focusing on the individual time series locations with records longer than 15 years, there is an overall decline of 0.017–0.027 (across 99% confidence intervals). Trends calculated from repeat measurements on ocean surveys show a consistent value of around –0.02 pH units decade-1 for diverse oceanic regions (Table SM5.3), with greater subsurface than surface trends reported in the subtropical oceans (Dore et al., 2009). At larger spatial scales, surface-ocean pH trends are assessed using shipboard observations of the fugacity of CO2 and estimates of ocean alkalinity (Takahashi et al., 2014; Lauvset et al., 2015). Between 1991–2011, mean surface-ocean pH has declined by 0.018 ± 0.004 units decade–1 in 70% of ocean biomes, with the largest declines in the Indian Ocean (–0.027 units decade–1), eastern Equatorial Pacific (–0.026 units decade–1) and the South Pacific subtropical (–0.022 units decade–1) biomes (Lauvset et al., 2015). Due to the close link between carbonate ion concentrations and pH, mean trends in the stability of mineral forms of aragonite and calcite (known as the ‘saturation state’) that are important for organisms such as coccolithophorids, pteropods and corals follow those of pH, with high-latitude regions most vulnerable to under-saturation due to naturally lower mean values.

It is virtually certain that ocean pH is declining, and the very likely range of this decline is 0.017–0.027 pH units per decade for the 8 locations where individual time series observations longer than 15 years exist. This trend is lowering the chemical stability of mineral forms of calcium carbonate and can be attributed to rising atmospheric CO2 levels. 

CMIP5 models are in good agreement with historical observations of declining surface-ocean pH (Figure 5.8a). Models project global surface-ocean declines between 2006–2015 and 2081–2100 of 0.287–0.291 and 0.036–0.042 pH units (both across 99% confidence intervals) for the RCP2.6 and RCP8.5 scenarios, respectively, with higher reductions in the subsurface of subtropical oceans (Bopp et al., 2013; Gattuso et al., 2015). These changes in pH will be greatest in the Arctic Ocean and the high latitudes of the Atlantic and Pacific Oceans due to their lower buffer capacity and are lowest in contemporary upwelling systems (Figure 5.8b) and will also reduce the stability of calcite minerals (Bopp et al., 2013; Gattuso et al., 2015). The area of the surface ocean (0–10 m) characterised by undersaturated conditions in CMIP5 models by 2081–2100 reduces from a very likely range of 6.4–9.5 x 1012 m2 or 5.5–7.3 x 1013 m2 under RCP8.5 (as much as 16–20% of ocean surface area for aragonite), to just 0.01–0.2  x 1012 m2 or 0.01–0.13 x 1013 m2 under RCP2.6 for either calcite or aragonite minerals, respectively. Under RCP8.5, hotspots for undersaturated waters for calcite remain restricted to the Arctic Ocean, while for aragonite, much of the Southern Ocean and the North Pacific and Northwestern Atlantic Oceans are also projected to become undersaturated (Orr et al., 2005; Hauri et al., 2015; Sasse et al., 2015). These results arise from the very well understood reductions in carbonate ion concentrations at lower pH, the vulnerability of regions with naturally low mean values, and the greater overall sensitivity of aragonite solubility. Regional models, with higher resolution that ESMs, also project year-round corrosive conditions for aragonite in some eastern boundary upwelling systems (Franco et al., 2018a). In the ocean interior, the decline in pH and calcium carbonate saturation state is more uncertain across models (Steiner et al., 2014) as it is modulated by changes to ocean overturning and water mass subduction (Resplandy et al., 2013; Chen et al., 2017). Projected benthic changes in pH over the next century are highly localised and are linked to transport of surface anomalies to depth, with over 20% of the north Atlantic sea floor deeper than 500 m projected to experience pH reductions greater than 0.2 units by 2100 under the RCP8.5 scenario (Gehlen et al., 2014). Changes in pH in the abyssal ocean (>3000 m deep) are greatest in the Atlantic and Arctic Oceans, with lesser impact in the Southern and Pacific Oceans by 2100, mainly due to the circulation timescales (Sweetman et al., 2017). 

Overall, it is virtually certain that the future surface open ocean will experience pH drops of either 0.036–0.042 (RCP2.6) or 0.287–0.291 (RCP8.5) pH units by 2081–2100, relative to 2006–2105. These pH changes are very likely to cause 16–20% of the surface ocean, specifically the Arctic and Southern Oceans, as well as the northern Pacific and northwestern Atlantic Oceans, to experience year-round corrosive conditions for aragonite by 2081–2100. It is virtually certain these impacts will be avoided under the RCP2.6 scenario. There is medium confidence, due to the potential for parallel changes in ocean circulation, that the Arctic and north Atlantic seafloors will experience the largest pH changes over the next century.

Although ocean acidification results in long-term trends in mean ocean chemistry, it can also influence seasonal cycles. Observation-based products indicate that the seasonal cycle of global surface-ocean pCO2 increased in amplitude by 2.2 ± 0.4 μatm between 1982 and 2014 (Landschützer et al., 2018). CMIP5 models and data-based products similarly project consistent future increases in the seasonal cycle of surface-ocean pCO2 under the RCP8.5 emissions scenario, with enhanced amplification in high-latitude waters (McNeil and Sasse, 2016). The amplitude of the seasonal cycle of global surface-ocean free acidity ([H+]) is projected to increase by 71–91% (across 90% confidence intervals) over the 21st century under RCP8.5, also with greater amplification in the high-latitudes (Kwiatkowski and Orr, 2018). Conversely, models project a 12–20% reduction (across 90% confidence intervals) in the seasonal amplitude of surface-ocean pH, as changes in pH represent relative changes in [H+] due to their logarithmic relationship, and there are typically greater projected increases in annual mean state [H+] than the seasonal amplitude of [H+]. Models also project a 4–14% (across 90% confidence intervals) reduction in the seasonal amplitude of global mean surface-ocean aragonite saturation state under RCP8.5, with a slight amplification in the subtropics being outweighed by dampening elsewhere. The contrasting changes in the seasonal amplitudes of ocean carbonate chemistry variables derive from different sensitivities to atmospheric CO2 and climate change and to diverging trends in the seasonal cycles of DIC, alkalinity and temperature. Model skill at simulating the seasonal cycles of carbonate chemistry is moderate, with persistent biases in the Southern Ocean, particularly for pCO2, [H+] and pH (Kwiatkowski and Orr, 2018; Mongwe et al., 2018).

Overall, we assess that alongside the strong mean state changes, it is very likely that the amplitude of the seasonal cycle in free acidity will increase by 71–91%, while it is very likely that the seasonal cycles of pH and aragonite saturation will decrease by 12–20% and 4–14%, respectively. 

WGII AR5 concluded that the vulnerability of most organisms to warming is set by their physiology, which defines their limited temperature ranges and thermal sensitivity (Pörtner et al., 2014³⁸⁵). Although different hypotheses have been proposed since AR5 to explain the mechanism linking temperature sensitivity of marine organisms and their physiological tolerances (Schulte, 2015³⁸⁶; Pörtner et al., 2017³⁸⁷; Somero et al., 2017³⁸⁸), evidence from physiological experiments and observations from paleo- and contemporary periods continue to support the conclusion from AR5 on the impacts of temperature change beyond thermal tolerance ranges on biological functions such as metabolism, growth and reproduction (Payne et al., 2016³⁸⁹; Pörtner and Gutt, 2016³⁹⁰; Gunderson et al., 2017³⁹¹), contributing to changes in biogeography and community structure (Beaugrand et al., 2015³⁹²; Stuart-Smith et al., 2015³⁹³) (high agreement, high confidence). Comparison of biota across land and ocean suggests that marine species are generally inhabiting environment that is closer to their upper temperature limits, explaining the substantially higher rate of local extirpation related to warming relative to those on land (Pinsky et al., 2019³⁹⁴). Hypoxia and acidification can also limit the temperature ranges of organisms and exacerbate their sensitivity to warming (Mackenzie et al., 2014³⁹⁵; Rosas-Navarro et al., 2016³⁹⁶; Pörtner et al., 2017³⁹⁷), although interactions vary strongly between species and biological processes (Gobler and Baumann, 2016³⁹⁸; Lefevre, 2016³⁹⁹).

Shifts in distribution of marine species from phytoplankton to marine mammals continued to be observed since AR5 across all ocean regions (Poloczanska et al., 2016⁴⁰⁰). Recent evidence continues to support that a large proportion of records of observed range shifts in the epipelagic ecosystem (Poloczanska et al., 2016⁴⁰¹) are correlated with ocean temperature, with an estimated average shift in distribution (including range centroids, northward and southward boundaries) from these records of 51.5 ± 33.3 km per decade since the 1950s (Figure 5.13). Such rate of shift is significantly faster than those records for organisms in the seafloor; the latter has an average rate of distribution shift of 29.0 ± 15.5 km per decade (44% of the records for seafloor species with range shifts that are consistent with expectation from the observed temperature changes)(very likely) (Figure 5.13). Comparison of global seafloor-derived planktonic foraminifera from pre-industrial age with recent (from year 1978) communities show that the recent assemblages differ from their pre-industrial with increasing dominance of warmer or cooler species that are mostly consistent with temperature changes (Jonkers et al., 2019⁴⁰²). Rate of observed responses also varies between and within animal groups among ocean regions, with zooplankton and fishes having faster recorded range shifts (Pinsky et al., 2013⁴⁰³; Asch, 2015⁴⁰⁴; Jones and Cheung, 2015⁴⁰⁵; Poloczanska et al., 2016⁴⁰⁶). For example, analysis of the Continuous Plankton Recorder (CPR) data-series from the north Atlantic in the last decades shows that the range of dinoflagellates tended to closely track the velocity of climate change (the rate of isotherm movement). In contrast, the distribution range of diatoms shifted much more slowly (Chivers et al., 2017⁴⁰⁷) and its distribution seems to be primary influenced by multi-decadal variability rather than from secular temperature trends. The CPR surveys have also provided evidence that some calanoid copepods are expanding poleward in the Northeast Atlantic, at a rate up to 232 km per decade (Beaugrand, 2009⁴⁰⁸; Chivers et al., 2017⁴⁰⁹), although different calanoid species respond differently in the rate and direction of shifts (Philippart et al., 2003⁴¹⁰; Edwards and Richardson, 2004⁴¹¹; Asch, 2015⁴¹²; Crespo et al., 2017⁴¹³). Overall, the observed changes in biogeography are consistent with expected responses to changes in ocean temperature for the majority of marine biota (high confidence). This is also consistent with theories and experimental evidence that scale from individual organisms’ physiological responses to community level effects (high confidence). Sensitivity of organisms’ biogeography varies between taxonomic groups (high confidence). 

The rate and direction of observed range shifts are shaped by the interaction between climatic and non-climatic factors (Poloczanska et al., 2013⁴¹⁴; Sydeman et al., 2015⁴¹⁵; Poloczanska et al., 2016⁴¹⁶), such as local temperature and oxygen gradients in the habitat across depth (Cheung et al., 2013⁴¹⁷; Deutsch et al., 2015⁴¹⁸), latitude and longitude (Burrows et al., 2014⁴¹⁹; Barton et al., 2016⁴²⁰), ocean currents (Sunday et al., 2015⁴²¹; Barton et al., 2016⁴²²; García Molinos et al., 2017⁴²³), bathymetry in all or part of their life stages (for organisms living on or close to the seafloor) (Pinsky et al., 2013⁴²⁴; Kleisner et al., 2015⁴²⁵), geographical barriers (Pinsky et al., 2013⁴²⁶; Burrows et al., 2014⁴²⁷), availability of food and critical habitat (Sydeman et al., 2015⁴²⁸), fishing and other non-climatic human impacts (Engelhard et al., 2014⁴²⁹; Hoegh-Guldberg et al., 2014⁴³⁰). Moreover, observed range shifts in respond to climate change in some regions such as the north Atlantic are strongly influenced by warming due to multi-decadal variability (Edwards et al., 2013⁴³¹; Harris et al., 2014⁴³²), suggesting that there is a longer time-of-emergence of range shifts from natural variability and a need for longer biological time series for robust attribution. The rate of shifts in biogeography of organism is influenced by multiple climatic and non-climatic factors (high confidence) that can result in non-synchronous shifts in community composition (high confidence). There is general under-representation of biogeographical records in low latitudes (Dornelas et al., 2018⁴³³), rendering detection and attribution of shifts in biogeography in these regions having medium confidence. The variation in responses of marine biota to range shifts can cause spatial restructuring of the pelagic ecosystem with consequences for organisms at higher trophic levels (Chivers et al., 2017⁴³⁴; Pecl et al., 2017⁴³⁵) (high confidence). Marine ectotherms have demonstrated some capacity for physiological adjustment and evolutionary adaptation that lowers their sensitivity to warming and decrease in oxygen (Pörtner et al., 2014⁴³⁶; Cavallo et al., 2015⁴³⁷) (low confidence). However, historical responses in abundance and ranges of marine species to ocean warming suggest that adaptation not always suffices to mitigate projected impacts (WGII AR5 Chapter 6) (high confidence). 

Marine reptiles, seabirds and mammals breathe air, instead of obtaining oxygen from water, and many of them spend some of their life cycle on land, being their abundance and distribution still affected by temperature (Pörtner et al., 2014⁴³⁸). Long term population changes and shifts in distribution associated with climate change have been observed for temperate species of seabirds and marine mammals (Henderson et al., 2014⁴³⁹; Hiscock and Chilvers, 2014⁴⁴⁰; Ramp et al., 2015⁴⁴¹) (high confidence). For example, Laysan, Phoebastria immutabilis, and Wandering, Diomedea exulans, albatross have responded positively to climate change as they have been able to take advantage of the increased intensity of winds. This has allowed them to forage farther and faster, making their foraging trips shorter, increasing their foraging efficiency and breeding success (Descamps et al., 2015⁴⁴²; Thorne et al., 2016⁴⁴³). For reptiles, like sea turtles and snakes, temperature directly affects important life history traits including hatchling size, sex, viability and performance (high confidence) (Hays et al., 2003⁴⁴⁴; Pike, 2014⁴⁴⁵; Dudley et al., 2016⁴⁴⁶; Santos et al., 2017⁴⁴⁷). This is particularly important for marine turtles as changing temperatures will affect the hatchling sex ratio because sex is determined by nest site temperature (high confidence) (Hatfield et al., 2012⁴⁴⁸; Santidrián Tomillo et al., 2014⁴⁴⁹; Patricio et al., 2017⁴⁵⁰). Loss of breeding substrate, including mostly coastal habitats such as sandy beaches (Section 5.3.3), can reduce the available nesting or pupping habitat for land breeding marine turtles, lizards, seabirds and pinnipeds (Fish et al., 2005⁴⁵¹; Fuentes et al., 2010⁴⁵²; Funayama et al., 2013⁴⁵³; Reece et al., 2013⁴⁵⁴; Katselidis et al., 2014⁴⁵⁵; Patino-Martinez et al., 2014⁴⁵⁶; Pike et al., 2015; Reynolds et al., 2015; Marshall et al., 2017) (high confidence). Climatic hazards such as SLR contributes to the loss of these coastal habitats (see Section 5.3 and Chapter 3). Changes in ocean temperature will also indirectly impact marine mammals, seabirds and reptiles by changing the abundance and distribution of their prey (Polovina, 2005⁴⁷⁹; Polovina et al., 2011⁴⁸⁰; Doney et al., 2012⁴⁸¹; Sydeman et al., 2015⁴⁸²; Briscoe et al., 2017⁴⁸³; Woodworth-Jefcoats et al., 2017⁴⁸⁴) (high confidence). The distributions of some of these large animals is determined by the occurrence and persistence of oceanic bridges and barriers that are related to climate driven processes (Ascani et al., 2016⁴⁸⁵; McKeon et al., 2016⁴⁸⁶). For example, the decline of Arctic sea ice is affecting the range and migration patterns of some species and is allowing the exchange of species previously restricted to either the Pacific or Atlantic oceans (Alter et al., 2015⁴⁸⁷; George et al., 2015⁴⁸⁸; Laidre et al., 2015⁴⁸⁹; MacIntyre et al., 2015⁴⁹⁰; McKeon et al., 2016⁴⁹¹; Breed et al., 2017⁴⁹²; Hauser et al., 2017⁴⁹³) (Chapter 3). Also, the range expansion of some of these predatory megafauna can affect species endemic to the habitat; for example, while the decrease in summer sea ice in the Arctic may favour the expansion of killer whales (Orcinus orca), their occurrence can result in narwhale (Monodon monoceros) to avoid the use of key habitats to reduce the risk of killer whales’ predation (Bost et al., 2009⁴⁵⁹; Sydeman et al., 2015⁴⁶⁰; Breed et al., 2017⁴⁶¹) (see Chapter 3; section 3.2.1.4). In addition, marine mammals, seabirds and sea turtles present habitat requirements associated with bathymetric and mesoscale features that facilitate the aggregation of their prey (Bost et al., 2015⁴⁶²; Kavanaugh et al., 2015⁴⁶³; Hindell et al., 2016⁴⁶⁴; Hunt et al., 2016⁴⁶⁵; Santora et al., 2017⁴⁶⁶). The persistence and location of these features are linked to variations in climate (Crocker et al., 2006⁴⁶⁷; Baez et al., 2011⁴⁶⁸; Dugger et al., 2014⁴⁶⁹; Abrahms et al., 2017⁴⁷⁰; Youngflesh et al., 2017⁴⁷¹) and to foraging success, juvenile recruitment, breeding phenology, growth rates and population stability (Costa et al., 2010⁴⁷²; Ancona and Drummond, 2013⁴⁷³; Ducklow et al., 2013⁴⁷⁴; Chambers et al., 2014⁴⁷⁵; Descamps et al., 2015⁴⁷⁶; Abadi et al., 2017⁴⁷⁷; Bjorndal et al., 2017⁴⁷⁸; Fluhr et al., 2017; Youngflesh et al., 2017) (high confidence). Overall, recent evidence further support that impacts of climate change on some marine reptiles, mammals and birds have been observed in recent decades (high confidence) and that the direction of impacts vary between species, population and geographic locations (Trivelpiece et al., 2011; Hazen et al., 2013; Clucas et al., 2014; Constable et al., 2014; George et al., 2015) (high confidence). 

Warming has contributed also to observed changes in phenology (timing of repeated seasonal activities) of marine organisms (Gittings et al., 2018⁵⁰⁵), although observations are biased towards the northeast Atlantic (Poloczanska et al., 2016⁵⁰⁶; Thackeray et al., 2016). Shifts in the timing of interacting species have occurred in the last decades, eventually leading to uncoupling between prey and predators, with cascading community and ecosystem consequences (Kharouba et al., 2018; Neuheimer et al., 2018⁵⁰⁷). Timing of spring phenology of marine organisms is shifting to earlier in the year under warming, at an average rate of 4.4 ± 1.1 days per decade (Poloczanska et al., 2013⁵⁰⁸), although it is variable among taxonomic groups and among ocean regions (Lindley and Kirby, 2010⁵⁰⁹). This is consistent with the expectations based on the close relationship between temperature and these biological events, supporting evidence from AR5 (Bruge et al., 2016⁵¹⁰; Poloczanska et al., 2016⁵¹¹). Thus, the growing amount of literature and new studies since AR5 WGII and SR15 further support that phenology of marine ectotherms in the epipelagic systems are related to ocean warming (high confidence) and that the timing of biological events has shifted earlier (high confidence).

Observed impacts of multiple climatic hazards

WGII AR5 concludes that multiple climatic hazards from ocean acidification, hypoxia and decrease in nutrient and food supplies pose risks to marine ecosystems, and the risk can be elevated when combined with warming (Riebesell and Gattuso, 2014⁵¹²; Gattuso et al., 2015⁵¹³). In a recent meta-analysis of 632 published experiments, primary production by temperate non-calcifying plankton increases with elevated temperature and CO2, whereas tropical plankton decreases productivity because of acidification (Nagelkerken and Connell, 2015⁵¹⁴). Also, temperature increases consumption and metabolic rates of herbivores but not secondary production; the latter decreases with acidification in calcifying and non-calcifying species. These effects together create a mismatch with carnivores whose metabolic and foraging costs increase with temperature (Nagelkerken and Connell, 2015⁵¹⁵). Warming may also exacerbate the effects of ocean acidification on the rate of photosynthesis in phytoplankton (Lefevre, 2016⁵¹⁶). There is some, but limited, reports of observed impacts on calcified pelagic organisms that are attributed to secular trend in ocean acidification and warming (Harvey et al., 2013⁵¹⁷; Kroeker et al., 2013⁵¹⁸; Nagelkerken et al., 2015⁵¹⁹; Boyd et al., 2016⁵²⁰). For example, Rivero-Calle et al. (2015) reported, using CPR archives, that stocks of coccolithophores (a group of phytoplankton that forms calcium carbonate plateles) have increased by 2% to over 20% in the north Atlantic over the last five decades, and that this increase is linked to synergistic effects of increasing anthropogenic CO2 and rising temperatures, as supported by their statistical analysis and a number of experimental studies. Most of the available evidence supports that ocean acidification and hypoxia can act additively or synergistically between each other and with temperature across different groups of biota (Figure 5.13). Limitation of nutrient and food availability and predation pressures can further increase the sensitivity of organismal groups to climate change in specific ecosystems (Riebesell et al., 2017⁴⁹⁴). Climate change also affects organisms indirectly through the impacts on competitiveness between organisms that favour those that are more adaptive to the changing environmental conditions (Alguero-Muniz et al., 2017⁴⁹⁵) and changes in trophic interactions (Seebacher et al., 2014⁴⁹⁶). Overall, direct in situ observations and laboratory experiments show that there are significant responses to the multiple stressors of warming, ocean acidification and low oxygen on phytoplankton, zooplankton and fishes and that these responses can be additive or synergistic (high confidence, Figure 5.13). 

WGII AR5 and SR15 conclude that projected ocean warming will continue to cause poleward shifts in the distribution and biomass of pelagic species, paralleled by altered seasonal timing of their activities, species abundance, migration pattern and reduction in body size in the 21st century under scenarios of increasing greenhouse gas emission (Pörtner et al., 2014⁴⁹⁷; Hoegh-Guldberg et al., 2018⁴⁹⁸). Simultaneously, projected expansion of OMZ and ocean acidification could lead to shifts in community composition toward hypoxia-tolerant and non-calcified organisms, respectively. However, these projected biological changes in the ocean raise questions about how individuals, communities and food webs will respond to the multiple impacts from climatic and non-climatic stressors in the future, and the feedbacks of the effects of their ecological impacts on modifying the physical and biogeochemical conditions of the ocean (Schaum et al., 2013⁴⁹⁹; Boyd et al., 2016⁵⁰⁰; O’Brien et al., 2016⁵⁰¹; Moore, 2018⁵⁰²). This section focuses on addressing these questions in order to assess the future risk of impacts of climate change on the epipelagic ecosystem.

Future projections on phytoplankton distribution, community structure and biomass

While analysis of outputs from CMIP5 ESMs project that global average NPP and biomass of phytoplankton community will decrease in the 21st century under RCP2.6 and RCP8.5 (see Section 5.2.2.6). However, the future risk of impacts of epipelagic ecosystem can also depend on changes in community structure of phytoplankton species. Barton et al. (2016) projected the biogeography of 87 taxa of phytoplankton (diatoms and dinoflagellates) in the north Atlantic to 2051–2100 relative to the past (1951–2000) with scenarios of changes in temperature and other ocean conditions such as salinity, density and nutrients under RCP8.5. The study found that 74% of the studied taxa exhibit a poleward shift at a median rate of 12.9 km per decade, but 90% of the taxa shift eastward at a median rate of 42.7 km per decade. Such changes may affect food webs and biogeochemical cycles, and with consequence to the productivity of living marine resources (Stock et al., 2014⁵⁰³; Barton et al., 2016⁵⁰⁴).

Outputs from CMIP5 ESMs suggest that projected warming and reduction in nutrient availability in low latitudes, as a result of increasing stratification of the ocean under climate change, will increase the dominance of small-sized phytoplankton, growing more efficiently than larger taxa at low nutrient levels (Dutkiewicz et al., 2013b⁵²²). Dominant groups in subtropical oceans, like the picoplanktonic cyanobacteria Synechococcus and Prochlorococcus, are projected to expand their range of distribution towards higher latitudes and increase their abundances by 14–29%, respectively, under a future warmer ocean (Flombaum et al., 2013⁵²³), although synergistic effects of warming and CO2 on photosynthetic rates could lead to a dominance of Synechococcus over Prochlorococcus (Fu et al., 2007⁵²⁴) (low confidence). Similarly, temperature-driven range shifts towards higher latitudes are also likely for tropical diazotrophic (N2-fixing) cyanobacteria, although they could disappear from parts of their current tropical ranges where future warming may exceed their maximum thermal tolerance limits (Hutchins and Fu, 2017⁵²⁵) (low confidence). Modelling experiments show that the effects of warming on phytoplankton community will be exacerbated by ocean acidification at levels expected in the 21st century for RCP8.5, leading to increasing growth rate responses of some phytoplankton groups, such as diazotrophs and Synechococcus, with predicted increases in biomass up to 10% in tropical and subtropical waters (Dutkiewicz et al., 2015⁵²⁶) (low confidence). Furthermore, warming is projected to interact with decreasing oxygen levels and increases in iron in the nutrient-impoverished subtropical waters, favoring the dominance of the diazotrophic colonial cyanobacteria Trichodesmium (Sohm et al., 2011⁵²⁷; Boyd et al., 2013⁵²⁸; Ward et al., 2013⁵²⁹; Hutchins and Fu, 2017⁵³⁰) (medium confidence). 

Regional differences in the changes in phytoplankton community and their impacts on epipelagic ecosystem are however complex and depends on multiple interactions of co-varying climate change stressors at regional level (Boyd and Hutchins, 2012⁵³¹). Based on global ocean model simulations, Boyd et al. (2015b) show that the interaction between warming, increased CO2 and a decline in phosphate and silicate would benefit coccolithophores against diatoms in the northern north Atlantic, despite decreasing rates of calcification. Evidence, based on long-term experiments of acclimation or adaptation to increasing temperatures in combination with elevated CO2, show that individual growth and carbon fixation rates of coccolithophores at high CO2 are modulated by temperature, light, nutrients and UV radiation, and could increase calcification while the responses are also species-specific (Lohbeck et al., 2012⁵³²; Khanna et al., 2013⁵³³). Calcification of planktonic foraminifera will be however negatively affected by acidification (Roy et al., 2015⁵³⁵), and their populations are predicted to experience the greatest decrease in diversity and abundance in sub-polar and tropical areas, under RCP8.5 (Brussaard et al., 2013⁵³⁵), however environmental controls of calcite production by foraminifera are still poorly understood (low confidence). Boyd et al. (2015b) analysis indicate also that diatoms would benefit from the synergistic effects of increased warming and iron supply in the northern Southern Ocean, as supported by laboratory experiments and field studies with polar diatoms (Rose et al., 2009⁵³⁶) (low confidence). At low-latitude provinces, projected concurrent increases of CO2 and iron, and decreases in both nitrate and phosphate supply, may favour nitrogen fixers, but with ocean regional variability, since iron is thought to limit N2 fixation in the eastern Pacific and phosphorus in the Atlantic Ocean (Gruber, 2019⁵³⁷; Wang et al., 2019⁵³⁸). However, recent experimental work with the diazotrophic colonial Trichodesmium and the unicellular Crocosphaera have shown a broad range of responses from rising CO2, with either increases or decreases in N2 fixation rates, and with mixed evidence on co-limiting processes (Eichner et al., 2014⁵³⁹; Garcia et al., 2014⁵⁴⁰; Gradoville et al., 2014⁵⁴¹; Walworth et al., 2016⁵⁴²; Hong et al., 2017⁵⁴³; Luo et al., 2019⁵44) (low confidence). 

Overall, the response of phytoplankton to the interactive effects of multiple drivers is complex, and presently ESMs do not resolve the full complexity of their physiological responses (Breitberg et al., 2015⁵⁴⁵; Hutchins and Boyd, 2016⁵⁴⁶; O’Brien et al., 2016⁵⁴⁷), precluding a clear assessment of the effects of these regional distinctive multi-stressor patterns (high confidence).

Future projections on zooplankton distribution and biomass

An ensemble of 12 CMIP5 ESMs project average declines of 6.4 ± 0.79% (95% confident limits) and 13.6 ± 1.70%) in zooplankton biomass in the 21st century relative to 1990–1999 historical values under RCP2.6 and RCP8.5 (Kwiatkowski et al., 2019⁵⁴⁸). Also, production of mesozooplankton is projected from a single ESM to decrease by 7.9% between 1951–2000 and 2051–2100 under RCP8.5 (Stock et al., 2014⁵⁴⁹). Such projected decreases in zooplankton biomass and production are partly contributed by climate-induced reduction in phytoplankton production and trophic transfer efficiency particularly in low-latitude ecosystems (Stock et al., 2014⁵⁵⁰) (5.2.2.6). The impacts may be larger than these projections if changes in the relative abundance of carbon, nitrogen and phosphorus are considered by the models (Kwiatkowski et al., 2019⁵⁵¹). The overall projected decrease in zooplankton biomass is characterised by a strong latitudinal differences, with the largest decrease in tropical regions and increase in the polar regions, particularly the Arctic Ocean (Chust et al., 2014⁵⁵²; Stock et al., 2014⁵⁵³; Kwiatkowski et al., 2019⁵⁵⁴) (Chapter 3) (high agreement). However, the projected increase in zooplankton biomass in the polar region may be affected by the seasonality of light cycle at high latitudes that may limit the bloom season at high latitude (Sundby et al., 2016⁵⁵⁵). The projected decrease in zooplankton abundance, particularly in tropical regions, can impact marine organisms higher in the foodweb, including fish populations that are important to fisheries (Woodworth-Jefcoats et al., 2017⁵⁵⁶). Therefore, there is high agreement in model projections that global zooplankton biomass will very likely reduce in the 21st century, with projected decline under RCP8.5 almost doubled that of RCP2.6 (very likely). However, the strong dependence of the projected declines on phytoplankton production (low confidence, 5.2.2.6) and simplification in representation of the zooplankton communities and foodweb render their projections having low confidence. 

Future responses of zooplankton species and communities to climate change are however affected by interactions between multiple climatic drivers. Experiments in laboratory show that acidification could partly counteract some observed effects of increased temperature on zooplankton, although the level and direction of the biological responses vary largely between species (Mayor et al., 2015⁵⁵⁷; Garzke et al., 2016⁵⁵⁸), with results ranging from no effects (Weydmann et al., 2012⁵⁵⁹; McConville et al., 2013⁵⁶⁰; Cripps et al., 2014⁵⁶¹; Alguero-Muniz et al., 2016⁵⁶²; Bailey et al., 2016⁵⁶³), to negative effects (Lischka et al., 2011⁵⁶⁴; Cripps et al., 2014⁵⁶⁵; Alguero-Muniz et al., 2017⁵⁶⁶) or positive effects (Alguero-Muniz et al., 2017⁵⁶⁷; Taucher et al., 2017⁵⁶⁸). These differences in response can affect trophic interactions between zooplankton species; for example, some predatory non-calcifying zooplankton may perform better under warmer and lower pH conditions, leading to increased predation on other zooplankton species (Caron and Hutchins, 2012⁵⁶⁹; Winder et al., 2017⁵⁷⁰). Therefore, the large variation in sensitivity between zooplankton to future conditions of warming and ocean acidification suggests elevated risk on community structure and inter-specific interactions of zooplankton in the 21st century (medium confidence). Consideration of these species-specific responses may further modify the projected changes in zooplankton biomass by ESMs (Boyd et al., 2015a⁵⁷¹).

 Future projections on fish distribution, size and biomass

Recent model projections since AR5 and SR15 continue to support global-scale range shifts of marine fishes at rates of tens to hundreds of km per decade in the 21st century, with rate of shifts being substantially higher under RCP8.5 than RCP2.6 (Jones and Cheung, 2015⁵⁷²; Robinson et al., 2015⁵⁷³; Morley et al., 2018⁵⁷⁴). Globally, the general direction of range shifts of epipelagic fishes is poleward (Jones and Cheung, 2015⁵⁷⁵; Robinson et al., 2015⁵⁷⁶), while the projected directions of regional and local range shifts generally follow temperature gradients (Morley et al., 2018⁵⁷⁷). Polewards range shifts are projected to result in decreases in species richness in tropical oceans, and increases in mid to high-latitude regions leading to global-scale species turnover (sum of species local extinction and expansion) (Ben Rais Lasram et al., 2010; Jones and Cheung, 2015⁵⁷⁸; Cheung and Pauly, 2016⁵⁷⁹; Molinos et al., 2016⁵⁸⁰) (medium confidence on trends, low confidence on magnitude because of model uncertainties and limited number of published model simulations). For example, species turnover relative to their present day richness in the tropical oceans (30oN–30oS) is projected to be 14–21% and 37–39% by 2031–2050 and 2081–2100 under RCP8.5 (ranges of mean projections from two sets of simulation for marine fish distributions) (Jones and Cheung, 2015⁵⁸¹; Molinos et al., 2016⁵⁸²). In contrast, high-latitude regions (>60oN–60oS) is projected to have higher rate of species turnover than the tropics (an average of 48% between the two data sets for region >60oN). The high species turnover in the Arctic is explained by species’ range expansion from lower-latitude and the relatively lower present day fish species richness in the Arctic. The projected intensity of species turnover is lower under lower emission scenarios (Jones and Cheung, 2015⁵⁸³; Molinos et al., 2016⁵⁸⁴) (see also Section 5.4.1) (high confidence). Projections from multiple fish species distribution models show hotspots of decrease in species richness in the Indo-Pacific region, and semi-enclosed seas such as the Red Sea and Persian Gulf (Cheung et al., 2013⁵⁸⁵; Burrows et al., 2014⁵⁸⁶; García Molinos et al., 2015⁵⁸⁷; Jones and Cheung, 2015⁵⁸⁸; Wabnitz et al., 2018⁵⁸⁹) (medium evidence, high agreement). In addition, geographic barriers such as land boundaries in the poleward species range edge in semi-enclosed seas or lower oxygen water in deeper waters are projected to limit range shifts, resulting in larger relative decrease in species richness (medium confidence) (Cheung et al., 2013⁵⁹⁰; Burrows et al., 2014⁵⁹¹; García Molinos et al., 2015⁵⁹²; Jones and Cheung, 2015⁵⁹³; Rutterford et al., 2015⁵⁹⁴). 

Warming and decrease in oxygen content is projected to impact growth of fishes, leading to reduction in body size and contraction of suitable environmental conditions (Deutsch et al., 2015⁵⁹⁵; Pauly and Cheung, 2017⁵⁹⁶), with the intensity of impacts being directly related to the level of climate change. The projected reduction in abundance of larger-bodied fishes could reduce predation and exacerbate the increase in dominance of smaller-bodied fishes in the epipelagic ecosystem (Lefort et al., 2015⁵⁹⁷). Fishes exposed to ocean acidification level expected under RCP8.5 showed impairments of sensory ability and alteration of behaviour including olfaction, hearing, vision, homing and predator avoidance (Kroeker et al., 2013⁵⁹⁸; Heuer and Grosell, 2014⁵⁹⁹; Nagelkerken et al., 2015⁶⁰⁰). The combined effects of warming, ocean deoxygenation and acidification in the 21st century are projected to exacerbate the impacts on the body size, growth, reproduction and mortality of fishes, and consequently increases their risk of population decline (medium evidence, high agreement, high confidence). 

An ensemble of global-scale marine ecosystem and fisheries models that are part of the Fisheries and Marine Ecosystems Impact Models Intercomparison Project (FISHMIP) undertook coordinated simulation experiments and projected future changes in marine animals (mainly invertebrate and fish) globally under climate change (Lotze et al., 2018⁶⁰²). These models represent marine biota and ecosystems differently, ranging from population-based to functional traits- and size-based structure and their responses are driven primarily by temperature and NPP, although oxygen, salinity and ocean advection are considered in a subset of models and play a secondary role in affecting the projected changes in biomass (Blanchard et al., 2012⁶⁰³; Fernandes et al., 2013⁶⁰⁴; Carozza et al., 2016⁶⁰⁵; Cheung et al., 2016a⁶⁰⁶). Overall, potential total marine animal biomass is projected to decrease by 4.3 ± 2.0% (95% confident intervals) and 15.0 ± 5.9% under RCP2.6 and RCP8.5, respectively, by 2080–2099 relative to 1986–2005, while the decrease is around 4.9% by 2031-2050 across all RCP2.6 and RCP8.5 (very likely) (Figure 5.14). Accounting for the removal of biomass by fishing exacerbates the decrease in biomass for large-bodied animals which are particularly sensitive to fishing (likely for the direction of changes). Regionally, total animal biomass decreases largely in tropical and mid-latitude oceans (very likely) (Table 5.3, Figure 5.14) (Bryndum-Buchholz et al., 2019⁶⁰⁷). The high uncertainty and the low confidence in the projection in the Arctic Ocean (Chapter 3) is because of the large variations in simulation results for this region between the ESMs and between the FISHMIP models, as well as the insufficient understanding of the oceanographic changes and their biological implications in the Arctic Ocean. In the Southern Ocean, the decrease in consumer biomass is mainly in the southern Indian Ocean while other parts of the Southern Ocean are projected to have an increase in animal biomass by 2100 under RCP8.5, reflecting mainly the projected pattern of changes in NPP from the ESMs (see Section 5.2.2.6).

Future projections on epipelagic components of the biological pump

A wide range of studies, from laboratory experiments, mesocosm enclosures, synthesis of observations to modeling experiments, provide insights into how the multi-faceted components of the ‘biological pump’ (the physical and biologically mediated processes responsible for transporting organic carbon from the upper ocean to depth) are projected to be altered in the coming decades. A synthesis of the individual components reported to both influence the performance of the biological pump, and which are sensitive to changing ocean conditions, is presented in Table 5.4. The table lists the putative controlling of each environmental factor, such as warming, that influences the biological pump, and the reported modification (where available) of each individual factor by changing ocean conditions for both the epipelagic ocean and the deep ocean. Analyses of long-term trends in primary production and particle export production, as well as model simulations, reveal that increasing temperatures, leading to enhanced stratification and nutrient limitation, will have the greatest influence on decreasing the flux of particulate organic carbon (POC) to the deep ocean (Bopp et al., 2013⁶⁰⁸; Boyd et al., 2015a⁶⁰⁹; Fu et al., 2016⁶¹⁰; Laufkötter et al., 2016⁶¹¹). However, different lines of evidence (including observation, modeling and experimental studies) provide low confidence on the mechanistic understanding of how climatic drivers affect different components of the biological pump in the epipelagic ocean, as well as changes in the efficiency and magnitude of carbon export in the deep ocean (see section below and Table 5.4); this renders the projection of future contribution of the biological carbon pump to the export of POC to the deep ocean having low confidence. 

The pelagic realm of the deep ocean represents a key site for remineralisation of organic matter and long-term biological carbon storage and burial in the biosphere (Arístegui et al., 2009), but the observed effects of climate change on deep sea organisms, communities and biological processes are largely unknown (high confidence). Observational and model-based methods provide limited evidence that the transfer efficiency of organic carbon to the sea floor is partly controlled by temperature and oxygen in the mesopelagic zone, affecting microbial metabolism and zooplankton community structure, with highest efficiencies for high-latitude and OMZ) (see Section 5.2.2.4 for more detail on OMZs), while below 1000 m organic carbon transfer is controlled by particle sinking speed (Boyd et al., 2015a; Marsay et al., 2015; DeVries and Weber, 2017). However, there are contrasting results and low confidence on whether transfer efficiencies are highest at low or high latitudes (Boyd et al., 2015a; Marsay et al., 2015; Guidi et al., 2016; DeVries and Weber, 2017; Sweetman et al., 2017). There is also low confidence on the effects of increasing temperatures on POC remineralisation to CO2 versus POC solubilisation to dissolved organic carbon (DOC) by microbial communities and its storage as refractory DOC (i.e., with life times of >16,000 years) (Legendre et al., 2015). 

The global magnitude of the biological pump and how this will be affected by climate change is also uncertain. Model-based studies agree in projecting a global decline in particle gravitational flux to the deep sea floor, but with regional variability in both the total particle export flux and transfer efficiency (DeVries and Weber, 2017; Sweetman et al., 2017) (see Sections 5.2.2 and 5.2.4). However, recent evidence suggest that other physical and biological processes may contribute nearly as much as the gravitational flux to the carbon transport from the surface to the deep ocean (Boyd et al., 2019), with low confidence on the future rate of change in magnitude and direction of these processes. In particular, the ‘active flux’ of organic carbon due to vertical migration of zooplankton and fishes has been reported to account from 10 to 40% of the gravitational sinking flux (Bianchi et al., 2013; Davison et al., 2013; Hudson et al., 2014; Jónasdóttir et al., 2015; Aumont et al., 2018; Gorgues et al., 2019). Predictions based on model studies suggest that mesopelagic zooplankton and fish communities living at deep scattering layers (DSLs) will increase their biomass by 2100, enhancing their trophic efficiency, because of deep-ocean warming (Section 5.2.2.1; Figures 5.2 and 5.3) and shallowing of DSL (Proud et al., 2017) (low confidence). Expansion of OMZs (see Section 5.2.2.4) will also widen the DSL and increase the exposure of mesopelagic organisms to shallower depths (Gilly et al., 2013; Netburn and Anthony Koslow, 2015). In the California Current, the abundance of mesopelagic fishes is closely tied to variations in the OMZ, whose dynamic is linked to the Pacific Decadal Oscillation and ENSO cycles (Koslow et al., 2015). Some large predators, like the Humboldt squid, could indirectly benefit from expanding OMZs due to the aggregation of their primary food source, myctophid fishes (Stewart et al., 2014). However, many non-adapted fish and invertebrates (like diurnal vertical migrators) will have their depth distributions compressed, affecting the carbon transport and trophic efficiency of food webs in the mesopelagic (Stramma et al., 2011; Brown and Thatje, 2014; Rogers, 2015) (low confidence). In OMZ waters, where zooplankton is almost absent, like in the Eastern Tropical North Pacific, the microbial remineralisation efficiency of sinking particles would be reduced, eventually increasing the transfer efficiency of organic matter to the deep ocean and thus biological carbon storage (Cavan et al., 2017) (low confidence; Table 1). However, increases in ocean temperature may also lead to shallower remineralisation of POC in warm tropical regions, counteracting the storage of carbon in the dark ocean (Marsay et al., 2015). Overall, the direct impacts of climate change on the biological pump are not well understood for the deep pelagic organisms and ecosystems (Pörtner et al., 2014), and there is low confidence on the effect of climate change drivers on biological processes in the deep ocean (Table 5.1).

SR15 concluded that climate change will result in an aerial expansion and increased risk of water-borne disease with regional differences (high to very high confidence)(Hoegh-Guldberg et al., 2018). AR5 concluded that warming, excessive nutrient and seawater inundation due to SLR are projected to exacerbate the expansion and threat of cholera (Pörtner et al., 2014) (medium confidence). This assessment focuses on health risks caused by Vibrio bacteria and HABs. Vibrio cholerae (causing cholera) is estimated to be responsible for around 760,000 and 650,000 cases of human illness and death respectively in the world in 2010 (Kirk et al., 2015). An assessment of HABs is given in Box 5.4. 

Vibrio species naturally occur in warm, nutrient-rich and low salinity coastal waters. Since AR5, analysis of the the Continuous Plankton Recorder dataset (Section 5.2.3) has shown a significant increase in Vibrio abundance in the North Sea over the period 1958‒2011 related to sea surface warming (Vezzulli et al., 2016). Other time series data have confirmed a poleward expansion of Vibrio pathogens in mid- to high-latitude regions, ascribed at least partly to climate change (Baker-Austin et al., 2013; Baker-Austin et al., 2017). Extreme weather events such as flooding and tropical cyclones are also linked to increased incidences of Vibrio-related disease, suggested to be caused by the increased exposure of human populations to the pathogens during these extreme events (Baker-Austin et al., 2017). New evidence since AR5 therefore increases support for the linkages between warming, extreme weather events and increased risk of diseases caused by Vibrio bacteria (very high confidence).

Extrapolating from the observed relationship between environmental conditions and current Vibrio distributions, coastal areas that experience future warming, changes in precipitation and increases in nutrient inputs can be expected to see an increase in prevalence of Vibrio pathogens. These effects have been simulated in a global-scale model that relates occurrences of Vibrio with SST, pH, dissolved oxygen and chlorophyll a concentration under the SRES B1 scenario (Escobar et al., 2015). In the Baltic Sea, a nearly two-fold increase in the area suitable for Vibrio is projected between 2015 and 2050 for both RCP4.5 and RCP8.5 scenarios (relating to projected SST increase of 4°C‒5°C), resulting in an elevated risk of Vibrio infections (Semenza et al., 2017). Projected conditions of increased coastal flooding from storm surges and SLR (Section 5.2.2) will also increase exposure to waterborne disease (Ashbolt, 2019), such as Vibrio (medium confidence). However, uncertainty in the socioeconomic factors affecting the future vulnerabilities of human populations render quantitative projections of the magnitude of health impacts uncertain (Lloyd et al., 2016). 

Climate change–contaminant interactions can alter the bioaccumulation and amplify biomagnification of several contaminant classes (Boxall et al., 2009; Alava et al., 2018). This section assesses two types of contaminants that are of concern to environmental and human health as examples of other contaminants with similar properties (Alava et al., 2017). These two types of contaminants are the toxic and fat-soluble persistent organic pollutants (POPs), such as polychlorinated biphenyls (PCBs), as well as the neurotoxic and protein-binding organic form of mercury, methylmercury (MeHg) (Alava et al., 2017). POPs and MeHg are bioaccumulated by marine organisms and biomagnified in food webs, reaching exposure concentrations that become harmful and toxic to populations of apex predators such as marine mammals (Desforges et al., 2017; Desforges et al., 2018) (Figure 5.20). Human exposure to POPs and MeHg can lead to serious health effects (Ishikawa and Ikegaki, 1980; UNEP, 2013; Fort et al., 2015; Scheuhammer et al., 2015).

Inorganic forms of mercury are more soluble in low pH water, while higher temperature increases mercury uptake and the metabolic activity of bacteria, thereby increasing mercury methylation, uptake by organisms and bioaccumulation rates (Scheuhammer, 1991; Celo et al., 2006; López et al., 2010; Macdonald and Loseto, 2010; Riget et al., 2010; Corbitt et al., 2011; Krabbenhoft and Sunderland, 2013; Roberts et al., 2013; de Orte et al., 2014; McKinney et al., 2015), although there is limited evidence on the extent of exacerbation by ocean acidification expected in the 21st century. Increased melting of snow and ice from alpine ecosystems and mountains (Chapter 2) can also increase the release of POPs and MeHg from land-based sources into coastal ecosystems (Morrissey et al., 2005). Modelling projections for the Faroe Islands region suggest increased bioaccumulation of methyl mercury under climate change, with an average increases in MeHg concentrations in marine species of 1.6‒1.8% and 4.1‒4.7% under ocean warming scenarios of 0.8°C and 2.0°C, respectively, with an associated increase in potential human intake of mercury beyond levels recommended by the World Health Organization (Booth and Zeller, 2005). Foodweb modeling for the northeastern Pacific projects that concentrations of MeHg and PCBs in top predators could increase by 8% and 3%, respectively, by 2100 under RCP8.5 relative to current levels (Alava et al., 2018). Climate-related pollution risks are of particular concern in Arctic ecosystems and their associated indigenous communities because of the bioaccumulation of POPs and MeHg, causing long-term contamination of traditional seafoods (Marques et al., 2010; Tirado et al., 2010; Alava et al., 2017) of high dietary importance (Cisneros-Montemayor et al., 2016). 

Overall, climate change can increase the exposure and bioaccumulation of contaminants and thus the risk of impacts of POPs and MeHg on marine ecosystems and their dependent human communities as suggested by indirect evidence and model simulations (Marques et al., 2010; Tirado et al., 2010; Alava et al., 2017) (high agreement). However, there is limited evidence on observed increase in POPs and MeHg due to climate change. Apex predators and human communities that consume them, including Arctic communities and other coastal indigenous populations, are thus vulnerable to increase in exposure to these contaminants and the resulting health effects (medium evidence, medium agreement). 

The risk of microplastics has become a major concern for the ocean as they are highly persistent and have accumulated in many different marine environments, including the deep sea (Woodall et al., 2014; GESAMP, 2015; van Sebille et al., 2015; Waller et al., 2017; de Sá et al., 2018; Everaert et al., 2018; Botterell et al., 2019). There is limited evidence at present to assess their risk to marine ecosystems, wildlife and potentially humans through human consumption of seafood under climate change. 

Seafood provides protein, fatty acids, vitamins and other micronutrients essential for human health such as iodine and selenium (Golden et al., 2016). Over 4.5 billion people in the world obtain more than 15% of their protein intake from seafood, including algae and marine mammals as well as fish and shellfish (Béné et al., 2015; FAO, 2017)). Around 1.39 billion people obtain at least 20% of their supply of essential micronutrients from fish (Golden et al., 2016). SR15 concluded that global warming poses large risks to food security globally and regionally, especially in low-latitude areas, including fisheries (medium confidence) (Hoegh-Guldberg et al., 2018). This section builds on the assessment on observed and projected climate impacts on fish catches (Section 5.4.1.1) and further assess how such impacts interact with other climatic and non-climatic drivers in affecting food security through fisheries.

Many populations that are already facing challenges in food insecurity reside in low-latitude regions such as in the Pacific Islands and West Africa where maximum fisheries catch potential is projected to decrease under climate change security (Golden et al., 2016; Hilmi et al., 2017) (Section 5.4.1; Figure 5.21) and where land-based food production is also at risk (Blanchard et al., 2017) (medium confidence). Populations in these regions are also estimated to have the highest proportion of their micronutrient intake relative to the total animal sourced food (Golden et al., 2016) (ASF; Figure 5.21). This highlights their strong dependence on seafood as a source of nutrition that further elevates their vulnerability to food security from climate change impacts on seafood supply (high confidence). Modeling of seafood trade networks suggests that Central and West African nations are particularly vulnerable to shocks from decrease in seafood supply from international imports; thus their climate risks of seafood insecurity could be exacerbated by climate impacts on catches and seafood supply elsewhere (Gephart et al., 2016). In addition, experimental studies suggest that warming and ocean acidification reduce the nutritional quality of some seafood by reducing levels of protein, lipid and omega-3 fatty acids (Tate et al., 2017; Ab Lah et al., 2018; Lemasson et al., 2019). 

Non-climatic factors may exacerbate climate effects on seafood security. Over-exploitation of fish stocks reduces fish catches (Section 5.4.1.1) (Golden et al., 2016), whilst strong cultural dependence on seafood in many coastal communities may pose constraints in their adaptive capacity to changing fish availability (Marushka et al., 2019). The shift from traditional nutritious wild caught seafood-based diets of coastal indigenous communities, towards increased consumption of processed energy dense foods high in fat, refined sugar and sodium, due to social and economic changes (Kuhnlein and Receveur, 1996; Shannon, 2002; Charlton et al., 2016; Batal et al., 2017), has important consequences on diet quality and nutritional status (Thaman, 1982; Quinn et al., 2012; Luick et al., 2014). This has led to an increased prevalence of obesity, diabetes, and other diet-related chronic diseases (Gracey, 2007; Sheikh et al., 2011) as well as the related decrease in access to culturally or religiously significant food items. The risk of climate change on coastal communities through the ocean could therefore be increased by non-climatic factors such as economic development, trade, effectiveness of resource governance and cultural changes (high confidence). 

In summary, the food security of many coastal communities, particularly in low-latitude developing regions, is vulnerable to decreases in seafood supply (medium confidence) because of their strong dependence on seafood to meet their basic nutritional requirements (medium confidence), limited alternative sources of some of the essential nutrients obtained from seafood (medium confidence), and exposure to multiple hazards on their food security (high confidence). Although direct evidence from attribution analysis is not available, climate change may have already contributed to malnutrition by decreasing seafood supply in these vulnerable communities (low confidence) and reduce coastal Indigenous communities’ reliance on seafood-based diets (low confidence). Projected decreases in potential fish catches in tropical areas (high confidence) and a possible decrease in the nutritional content of seafood (low confidence) will further increase the risk of impacts on food security in low-latitude developing regions, with that risk being greater under high emission scenarios (medium confidence).

Climate change threatens key cultural dimensions of lives and livelihoods (Adger et al., 2012¹⁴⁵⁹), because people develop strong cultural ties and associate distinctive meanings with many natural places and biota in the form of traditions, customs and ways of life (Marshall et al., 2018¹⁴⁶⁰). These impacts have been felt both by indigenous and non-indigenous peoples. Recent estimates suggest that there are more than 1900 indigenous groups along the coastline with around 27 million people across 87 countries (Cisneros-Montemayor et al., 2016¹⁴⁶¹). AR5 concluded that climate change will affect the harvests of marine species with spiritual and aesthetic importance to indigenous cultures (Pörtner et al., 2014¹⁴⁶²). This section further assesses the effects of climate change on Indigenous knowledge and local knowledge and their transmission and the implication for well-being of people, complementing the assessment for Arctic indigenous people in Chapter 3. 

Indigenous knowledge is passed and appreciated over timeframes ranging from several generations to a few centuries (Cross Chapter Box 4 in Chapter 1). The adjustment of the transmission and the network of Indigenous knowledge on the ocean and coasts, and related perceptions and practice, implies a reworking of these knowledge systems where the individuals and the groups are actors in a narrative and historical construction (Roué, 2012¹⁴⁶³; Alderson-Day et al., 2015¹⁴⁶⁴). SLR is already transforming the seascape, such as the shape of shores in many low-lying islands in the Pacific, leading to modification or disappearance of geomorphological features that represent gods and mythological ancestors (Camus, 2017¹⁴⁶⁵; Kench et al., 2018¹⁴⁶⁶). These changing seascape also affects the mobility of people and residence patterns, and consequently, the structure and transmission of Indigenous knowledge (Camus, 2017¹⁴⁶⁷). The fear of SLR and climate change encourage security measures and the grouping of local people to the safest places, contributing to the erosion of indigenous culture and their knowledge about the ocean (Bambridge and Le Meur, 2018¹⁴⁶⁸), and impairment of opportunity for social elevation for some Pacific indigenous communities (Borthwick, 2016¹⁴⁶⁹). 

Climate change is also projected to shift the biogeography and potential catches of fishes and invertebrates (5.2.3.1, 5.3, 5.4.1.1) that form an integral part of the culture, economy and diet of many indigenous communities, such as those situated along the Pacific Coast of North America (Lynn et al., 2013¹⁴⁷⁰). Indigenous fishing communities that depend on traditional marine resources for food and economic security are particularly vulnerable to climate change through reduced capacity to conduct traditional harvests because of reduced access to, or availability of, resources (Larsen et al., 2014¹⁴⁷¹; Weatherdon et al., 2016¹⁴⁷²). Overall, the transmission of indigenous culture and knowledge is at risk because of SLR affecting sea- and land-scapes, the availability and access to culturally important marine species, and communities’ reliance on the ocean for their livelihood and their cultural beliefs (medium confidence). Strong attachment to traditional marine-based livelihoods has also been reported for non-indigenous communities in Canada (Davis, 2015), the USA (Paolisso et al., 2012¹⁴⁷⁴), Spain (Ruiz et al., 2012¹⁴⁷⁵) and Australia (Metcalf et al., 2015¹⁴⁷⁶). Reduction in populations of fish species that have supported livelihoods for generations, and deteriorations of iconic elements of seascapes are putting the well-being of these communities at risk (high confidence).

Other cultural values supported by the ocean are diverse. They include education, based on knowledge of marine environments. Such education can increase knowledge and awareness of climate change impacts and the efficacy of their mitigation (Meadows, 2011¹⁴⁷⁷); it can also influence the extent to which stewardship activities are adopted (von Heland et al., 2014; Wynveen and Sutton, 2015¹⁴⁷⁸; Bennett et al., 2018¹⁴⁷⁹), and can help develop new networks between coastal people and environmental managers for the purposes of planning and implementing new adaptation strategies (Wynveen and Sutton, 2015¹⁴⁸⁰). A critical element in reducing vulnerability to climate change is to educate people that they are an integral part of the Earth system and have a huge influence on the balance of the system. An important marine ecosystem service is to support such education (Malone, 2016¹⁴⁸¹). Thus, education can play a pivotal role in how climate change is perceived and experienced, and marine biodiversity and ecosystems play an important role in this. At the same time, climate change impacts on marine ecosystems (Sections 5.2.3, 5.2.4) can affect the role of the ocean in supporting such public education (medium evidence, high agreement, medium confidence).

The aesthetic appreciation of natural places is one of the fundamental ways in which people relate to their environment. AR5 noted that climate change may impact marine species with aesthetic importance that affect local and indigenous cultures, local economies and challenge cultural preservation (Pörtner et al., 2014¹⁴⁸²). Evidence since AR5 confirms that aesthetically appreciated aspects of marine ecosystems are important for supporting local and international economies (especially through tourism), human well-being, and stewardship. For example, Marshall et al. (2018) found that aesthetic values are a critically important cultural value for all cultural groups, and are important for maintaining sense of place, pride, identity and opportunities for inspiration, spirituality, recreation and well-being. However, climate change induced degradation and loss of biodiversity and habitats (Section 5.2.3, 5.2.4, 5.3) can also negatively impact the ecosystem features that are currently appreciated by human communities, such as coral reefs, mangroves, charismatic species (such as some marine mammals and seabirds) and geomorphological features (e.g., sandy beaches). There are also aesthetic and inspirational values of marine biodiversity and ecosystems that are important to the psychological and spiritual well-being of people, including film, literature, art and recreation (Pescaroli and Magni, 2015¹⁴⁸³). Other cultural dimensions that are becoming more widely acknowledged as potentially disturbed by climate change include the appreciation of scientific, artistic, spiritual, and health opportunities, as well as appreciation of biodiversity, lifestyle and aesthetics (Marshall et al., 2018¹⁴⁸⁴). Thus, climate change may also affect the way in which marine ecosystems support human well-being through cultural dimensions. However, the difficulties in evaluating the importance of aesthetic aspects of marine ecosystems, and in detecting and attributing of climate change impacts, result in such assessment having low confidence. 

Climate change affects human cultures and well-being differently. For example, Marshall et al. (2018)¹⁴⁸⁵ assessed the importance of identity, pride, place, aesthetics, biodiversity, lifestyle, scientific value and well-being within the Great Barrier Reef region by 8,300 people across multiple cultural groups. These groups included indigenous and non-indigenous local residents, Australians (non-local), international and domestic tourists, tourism operators, and commercial fishers. They found that all groups highly rated all (listed) cultural values, suggesting that these values are critically associated with iconic ecosystems. Climate change impacts upon the Great Barrier Reef, through increased temperatures, cyclones and SLR that cumulatively degrade the quality of the Reef, are therefore liable to result in cultural impacts for all groups. However, survey that assess the emotional responses to degradation of the Great Barrier Reef by similar stakeholder groups reported different levels of impacts among these groups (Marshall et al., 2019¹⁴⁸⁶). Therefore, many ocean and coastal dependent communities value marine ecosystems highly and climate impacts can affect their well-being, although the sensitivity to such impacts can vary among stakeholder groups (Marshall et al., 2019¹⁴⁸⁷) (low confidence).

Climate change may alter the environment too rapidly for cultural adaptation to keep pace. This is because the culture that forms around a natural environment can be so integral to people’s lives that disassociation from that environment can induce a sense of disorientation and disempowerment (Fisher and Brown, 2015¹⁴⁸⁸). The adaptive capacity of people to moderate or influence cultural impacts, and thereby reduce vulnerability to such impacts, is also culturally determined (Cinner et al., 2018¹⁴⁸⁹). For example, when a resource user such as a fisher, farmer, or forester is suddenly faced with the prospect that their resource-based occupation is no longer viable, they lose not only a means of earning an income but also an important part of their identity (Marshall et al., 2012¹⁴⁹⁰; Tidball, 2012¹⁴⁹¹). Loss of identity can, in turn, have severe economic, psychological and cultural impacts (Turner et al., 2008¹⁴⁹²). Climate change can quickly alter the quality of, or access to, a natural resource through degradation or coastal inundation, so that livelihoods and lifestyles are no longer able to be supported by that resource. When people are displaced from places that they value, there is strong evidence that their cultures are diminished, and in many cases endangered. There are no effective substitutions for, or adequate compensation for, lost sites of significance (Adger et al., 2012¹⁴⁹³). As sensitive marine ecosystems such as coral reefs and kelp forest are impacted by climate change at rapid rate (Section 5.3), these can lead to the loss of part of people’s cultural identity and values beyond the rate at which identify and values can be adjusted or substituted (medium confidence).

Redistribution of marine species in response to direct and indirect effects of climate change may also disrupt existing marine resource sharing and associated governance (Miller and Russ, 2014¹⁴⁹⁴; Pinsky et al., 2018¹⁴⁹⁵). These effects have contributed to disputes in international fisheries management for North Atlantic mackerel (Spijkers and Boonstra, 2017¹⁴⁹⁶) and Pacific salmon (Miller and Russ, 2014¹⁴⁹⁷). These disagreements have stressed diplomatic relations in some cases (Pinsky et al., 2018¹⁴⁹⁸). Decreases and fluctuations in fish stock abundance and fish catches have also contributed to past disputes (Belhabib et al., 2016¹⁴⁹⁹; Pomeroy et al., 2016¹⁵⁰⁰; Blasiak et al., 2017¹⁵⁰¹). Under climate change, shifts in abundance and distribution of fish stocks are projected to intensify in the 21st century (Sections 5.2.3, 5.3, 5.4.1.1). Stocks may locally increase and decrease elsewhere. New or increased fishing opportunities may be created when exploited fish stocks shift their distribution into a country’s waters where their abundance was previously too low to support viable fisheries (Pinsky et al., 2018¹⁵⁰²). The number of new transboundary stocks occurring in exclusive economic zones worldwide was projected to be around 46 and 60 under RCP2.6 and RCP8.5, respectively, by 2060 relative to 1950‒2014 (Pinsky et al., 2018¹⁵⁰³). However, such alteration of the sharing of resources between countries would challenge existing international fisheries governance regimes and, without sufficient adaptation responses, increase the potential for disputes in resource allocation and management (Belhabib et al., 2018¹⁵⁰⁴; Pinsky et al., 2018¹⁵⁰⁵). Overall, projected climate change impacts on fisheries in the 21st century increase the risk of potential conflicts among fishery area users and authorities or between two different communities within the same country (Ndhlovu et al., 2017¹⁵⁰⁶; Shaffril et al., 2017¹⁵⁰⁷; Spijkers and Boonstra, 2017¹⁵⁰⁸) (medium confidence), exacerbated through competing resource exploitation from international actors and mal-adapted policies (low confidence). Such risks can be reduced by appropriate fisheries governance responses that are discussed in Sections 5.5.2 and 5.5.3.

Global gross revenues from marine fisheries were around 150 billion in 2010 USD (Swartz et al., 2013; Tai et al., 2017). Capture fisheries provide full-time and part-time jobs for an estimated 260 ± 6 million people in the 2000s period, of whom 22 ± 0.45 million are small‐scale fishers (Teh and Sumaila, 2013¹⁵¹¹). Small-scale fisheries are important for the livelihood and viability of coastal communities worldwide (Chuenpagdee, 2011¹⁵¹²). AR5 concluded with low confidence that climate change will lead to a global decrease in revenue with regional differences that are driven by spatial variations of climate impacts on and the flexibility and capacities of food production systems (Pörtner et al., 2014¹⁵¹³). AR5 also highlighted the high vulnerability of mollusc aquaculture to ocean acidification. For example, the oyster industry in the Pacific has lost nearly 110 million USD in annual revenue due to ocean acidification (Ekstrom et al., 2015¹⁵¹⁴). This section examines the rapidly growing literature assessing the risks of climate change on fisheries and aquaculture sectors, and the potential interaction between climatic and non-climatic drivers on the economics of fisheries. However, new evidence on observed economic impacts of climate change on fisheries since AR5 is limited. 

Since AR5, projections on climate change impacts on the economics of marine fisheries have incorporated a broader range of social-economic considerations. Driven by shifts in species distributions and maximum catch potential of fish stocks (Section 5.4.1), if the ex-vessel price of catches remains the same, marine fisheries maximum revenue potential are projected to be negatively impacted in 89% of the world’s fishing countries under the RCP8.5 scenario by the 2050s relative to the current status, with projected global decreases of 10.4 ± 4.2% and 7.1 ± 3.5% under RCP8.5 and RCP2.6, respectively, by 2050 relative to 2000 (Lam et al., 2016). While the projected changes in revenues are sensitive to price scenarios (Lam et al., 2016¹⁵¹⁵), future maximum revenue potential is reduced under high emission scenarios (Sumaila et al., 2019¹⁵¹⁷). For example, when the elasticity of seafood price in relation to their supply was modelled explicitly, fisheries maximum revenue potential under a 1.5°C atmospheric warming scenario was projected to be higher than for 3.5°C warming by 7.4% (13.1 billion USD) ± 2.3%, across projections from three CMIP5 models (Sumaila et al., 2019¹⁵¹⁸). Accounting for the subsequent impacts on the dependent communities and relative to the 1.5°C warming scenario, that study also projected a decrease in seafood workers’ incomes of 7.8% (3.7 billion USD) ± 2.3% and an increase in households’ seafood expenditure by the global population of 3.2% (6.3 billion USD) ± 3.9% annually under a 3.5°C warming scenario (Sumaila et al., 2019¹⁵¹⁹). 

Fisheries management strategies and fishing effort affect the realised catch and economic benefits of fishing (Barange, 2019¹⁵²⁰). Modelling analysis of fish stocks with available data worldwide showed that for RCP6.0, adaptation of fisheries by accommodating shifts in species distribution and abundance, as well as rebuilding existing overexploited or depleted fish stocks, is projected to lead to substantially higher global profits (154%), harvest (34%), and biomass (60%) in the future, relative to a no adaptation scenario. However, the total profit, harvest and biomass are negatively affected even with the full adaptation scenario under RCP8.5 (Gaines et al., 2018¹⁵²¹). Overall, climate change impacts on the abundance, distribution and potential catches of fish stocks (see Section 5.3.1) are expected to reduce the maximum potential revenues of global fisheries (high agreement, medium evidence, medium confidence). These impacts on fisheries will increase the risk of impacts on the income and livelihoods of people working in these economic sectors by 2050 under high greenhouse gas emission scenarios relative to low emission scenario (high confidence). Rebuilding overexploited or depleted fisheries can help improve economic efficiency and reduce climate risk, provided that emissions are greatly reduced (medium confidence).

The economic implications of climate change on fisheries vary between regions and countries because of the differences in exposure to revenue changes and the sensitivity and adaptive capacity of the fishing communities to these changes (Hilmi et al., 2015¹⁵²²). Regions where the maximum potential revenue is projected to decrease coincide with areas where indicators such as human development index suggest high economic vulnerability to climate change (Barbier, 2015¹⁵²³; Lam et al., 2016¹⁵²⁴). Many coastal communities in these regions rely heavily on fish and fisheries as a major source of animal proteins, nutritional needs, income and job opportunities (FAO, 2019). Negative impacts on the catch and total fisheries revenues for these countries are expected to have greater implications for jobs, economies, food and nutritional security than the impacts on regions with high Human Development Index (Allison et al., 2009¹⁵²⁵; Srinivasan et al., 2010¹⁵²⁶; Golden et al., 2016¹⁵²⁷; Blasiak et al., 2017¹⁵²⁸). Climate change impacts to coral reefs and other fish habitats, as well as to targeted fish and invertebrate species themselves are expected to reduce harvests from small-scale, coastal fisheries by up to 20% by 2050, and by up to 50% by 2100, under RCP8.5 (Bell et al., 2018a¹⁵²⁹). Therefore, climate risk to communities that are strongly dependent on fisheries associated with ecosystems that are particularly sensitive to climate change such as coral reefs will have be particularly high (Cinner et al., 2016¹⁵³⁰) (high confidence). 

Climate change may also worsen non-climate related socioeconomic shocks and stresses, and hence is an obstacle to economic developments (Hallegatte et al., 2015¹⁵³¹). Climate risk on the economics of fishing is projected to be higher for tropical developing countries where existing adaptive capacity to the risk is lower, thereby challenging their sustainable economic development (high confidence). However, observed impacts are not yet well documented (Lacoue-Labarthe et al., 2016¹⁵³²) , and there are many uncertainties relating to how climate change would affect the dynamics of fishing costs, with consequent adjustment of fishing effort that might intensify or lessen the overcapacity issue. Studies have attempted to project how fishers may respond to changes in fish distribution and abundance by incorporating different management systems (Haynie and Pfeiffer, 2012¹⁵³³; Galbraith et al., 2017¹⁵³⁴). However, the impacts of climate change on management effectiveness and trade practices is still inadequately understood (Galbraith et al., 2017¹⁵³⁵). 

Tourism is one of the largest sectors in the global economy. Between 1995‒1998 and 2011‒2014, the average total contribution of tourism to global GDP increased from 69 billion USD (6.8%) to 166 billion USD (8.5%) respectively, and generated more than 21 million jobs between 2011‒2014 (UNCTAD, 2018¹⁵³⁶). Coastal tourism and other marine-related recreational activities contributes substantially to the tourism sector (Cisneros-Montemayor et al., 2013¹⁵³⁷; O’Malley et al., 2013¹⁵³⁸; Spalding et al., 2017¹⁵³⁹; Giorgio et al., 2018¹⁵⁴⁰; UNWTO, 2018¹⁵⁴¹). For example, it is estimated that around 121 million people a year participated in marine-based recreational activities, generating  47 billion in 2003 USD in expenditures and supporting one million jobs (Cisneros-Montemayor and Sumaila, 2010¹⁵⁴²). Tourism is one of the main industries that provides opportunities for social and economic development (Jiang and DeLacy, 2014¹⁵⁴³), and marine tourism is particularly important for many coastal developing countries and Small Island Developing States (SIDS). AR5 identified the tourism sector in the Caribbean region as particularly vulnerable to climate change effects, due to hurricanes, whilst SR15 concluded that warming will directly affect climate-dependent tourism markets on a worldwide basis (medium confidence) (Hoegh-Guldberg et al., 2018¹⁵⁴⁴). This assessment provides updates since AR5 and SR15. 

Empirical modelling of future risks to tourism is based on projected climate impacts (Section 5.3) for relevant coastal ecosystems, including degradation or loss of beach and coral reef assets (Weatherdon et al., 2016¹⁵⁴⁵) (Section 4.3.3.6.2). These projections are developed from the relationship between the economic benefits generated from coral reef related tourism with observed characteristics of coral reefs, the characteristics of tourism activities. Based on scenarios of projected future warming and decreases in coral reef coverage, a global loss of tourism and recreation value in the near-future (2031‒2050) of 2.57–2.95 billion yr-1 in 2000 USD is projected under RCP2.6, and of 3.88‒5.80 billion yr-1 in 2000 USD under RCP8.5 (Chen et al., 2015¹⁵⁴⁶). Opinion surveys in four countries suggest that if severe coral bleaching persists in the Great Barrier Reef, tourism in adjacent areas could greatly decline, from 2.8 million to around 1.7 million visitors per year, equivalent to more than 1 billion AUS (~0.69 billion USD using exchange rate in 2019), that is, in tourism expenditure and with potential loss of around 10,000 jobs (Swann and Campbell, 2016¹⁵⁴⁷). 

Many coastal tourism destinations are exposed to risks of flooding, SLR and coastal squeeze on coastal ecosystems (Lithgow et al., 2019¹⁵⁴⁸) (Section 5.3); there are also other climate related-risks. Droughts, which are projected to be more frequent, will also impact the tourism industry (and local food security) through water and food shortages (Pearce et al., 2018¹⁵⁴⁹). If climate change and ocean acidification reduce the seafood supply, the attractiveness of coastal regions for tourists will also decrease (Wabnitz et al., 2017¹⁵⁵⁰). North Atlantic hurricanes and tropical storms have increased in intensity over the last 30 years, with climate projections indicating an increasing trend in hurricane intensity (Chapter 6). Three major Caribbean storms, Harvey, Irma and Maria, occurred in 2017, with loss and damage to the tourism industries of Dominica, the British Virgin Islands, and Antigua and Barbuda estimated at 2.2 billion USD, and environmental recovery costs estimated at 6.8 million USD (UNDP, 2017¹⁵⁵¹). Pacific tourist destinations, which tend to focus on nature-based and marine activities, are also at high risk of extreme events and other climate change impacts (Klint et al., 2015¹⁵⁵²). However, global tourism has a high carbon footprint (flights, cruises, etc.) (Lenzen et al., 2018¹⁵⁵³), so any reduction in the intensity of this sector would help mitigate climate change. 

Evidence from recent studies on projected climate risks on recreational fishing is equivocal, with the direction of impacts depending on the location, species targeted and societal context. For example: poleward range shifts of marine fish (Section 5.2.3) could yield new opportunities for recreational fishing in mid- to high-latitude regions (DiSegni and Shechter, 2013¹⁵⁵⁴); projected increases in air temperature may enable longer fishing days in some area (Dundas and von Haefen, 2015¹⁵⁵⁵); and extreme events may alter the composition of recreational fishing catches (Santos et al., 2016¹⁵⁵⁶). Since climate risks to recreational fishing vary largely depending on the responses of the targeted species to climate-related pressures, there is low confidence in the overall risk to the activity. 

Overall, evidence since AR5 and SR15 confirms that climate impacts to coastal ecosystems would increase risks to coastal tourism, particularly under high emission scenarios (medium confidence). Economic impacts will be greatest for those developing countries where tourism is the main source of foreign revenue (medium to high evidence).

The integrity of ecosystems and their services can affect the value of human assets, particularly coastal properties and infrastructure (Hoegh-Guldberg et al., 2018¹⁵⁵⁷). Climate change is expected to have negative impacts on coastal properties and their value through the loss and damage caused by SLR, increased storm intensity (hurricanes and cyclones), heat waves, floods, droughts and other extreme events, particularly in tropical SIDS (Chapter 4). Natural disasters already cost Pacific Island Countries and Territories between 0.5‒6.6% of GDP yr-1 (World Bank, 2017¹⁵⁵⁸), with localised damages and losses from individual storms far exceeding these estimates (e.g., 64% of Vanuatu’s GDP for Cyclone Pam in 2015). The impacts of natural disasters on Jamaica’s coastal transport infrastructure are currently estimated to be a significant proportion of their GDP, and such costs are projected to increase substantially in the next few decades under climate change (UNCTAD, 2017¹⁵⁵⁹; Monioudi et al., 2018¹⁵⁶⁰). In 2015, tropical storm Erika devastated Dominica causing 483 million USD in damages and losses (mostly related to transport, housing and agriculture), equivalent to 90% of Dominica’s GDP (World Bank, 2017¹⁵⁶¹). For the USA, Ackerman and Stanton (2007) forecast that annual real estate losses due to climate change could increase from 0.17% of GDP in 2025 to 0.36% in 2100, with Atlantic and Gulf Coast states being the most vulnerable. Other North American studies have shown that informed coastal property owners are willing to initially invest in infrastructure to counter climate change impacts (McNamara and Keeler, 2013¹⁵⁶²); however, they would avoid further investment if adaptation costs increase substantially and there are greater risks of long-term impacts (Putra et al., 2015¹⁵⁶³).

The impacts of changing marine ecosystems and ecosystem services on the value of human assets need to consider the risk perception, future development and adaptation responses of human communities (Section 5.5.2, Chapter 4) (Bunten and Kahn, 2014¹⁵⁶⁴). For example, the potential for climate impacts on the value of coastal real estate will depend on the changing insurance market or the cost of adaptation measures, which in turn depend on the willingness to pay by asset holders and wider society, including local and national governments. Further research is needed to discount valuations for potential losses that may occur in the future but with uncertain occurrence, and to improve real estate loss estimates over local to regional scales. 

Marine ecosystem services contribute to climate moderation and coastal defenses (Section 5.4.1.2). However, while the above studies in this section acknowledge the contribution of many climate impacts on real estate and infrastructure through ecosystem losses and degradation, often they are not accounted for in quantitative economic impact assessments. Overall, there is high confidence that SLR, increases in storm intensity and other extreme events will impact the values of coastal real estates and infrastructure, particularly in tropical SIDS, through the risk and impacts of direct physical damages. However, there is low confidence that impacts due to underlying loss and damage of ecosystems and their services are being similarly accounted for.

Estuaries, shelf seas and a wide range of other intertidal and shallow-water habitats (Section 5.3) play an important role in the global carbon cycle through their primary production by rooted plants, seaweeds (macroalgae) and phytoplankton, and also by processing riverine organic carbon. However, the natural carbon dynamics of these systems have been greatly changed by human activities (Regnier et al., 2013¹⁶⁴⁸; Cloern et al., 2016¹⁶49; Day and Rybczyk, 2019¹⁶⁵⁰) (high confidence). Direct anthropogenic impacts include coastal land-use change (Ramesh et al., 2015¹⁶⁵¹; Li et al., 2018a¹⁶⁵²); indirect effects include increased nutrient delivery and other changes in river catchments (Jiao et al., 2011¹⁶⁵³; Regnier et al., 2013¹⁶⁵⁴), and marine resource exploitation in shelf seas (Bauer et al., 2013¹⁶⁵⁵). There is high confidence that these human-driven changes will continue, reflecting coastal settlement trends and global population growth (Barragán and de Andrés, 2015¹⁶⁵⁶). 

Policy recognition of the mitigation benefits of coastal ecosystems requires quantitative information on their actual and potential carbon uptake and storage at the local and national scale, within an international framework for carbon accounting (Crooks et al., 2011¹⁶⁵⁷; Hejnowicz et al., 2015¹⁶⁵⁸). Such methods are being developed for coastal habitats structured by rooted plants (Needelman et al., 2018¹⁶⁵⁹; Troxler et al., 2018¹⁶⁶⁰; Needelman et al., 2019¹⁶⁶¹), considered here as ‘coastal vegetation’, linked to protocols for verification of longterm carbon removal and financial incentives (Crooks et al., 2011¹⁶⁶²; Hejnowicz et al., 2015¹⁶⁶³) and building on techniques used for managing terrestrial carbon sinks (Ahmed and Glaser, 2016b¹⁶⁶⁴; Aziz et al., 2016¹⁶⁶⁵). Proposals to apply carbon accounting to seaweeds, the water column and shelf sea sediments (Krause-Jensen and Duarte, 2016¹⁶⁶⁶; Zhang et al., 2017¹⁶⁶⁷) are less well-developed. 

Mangrove, salt marsh and seagrass habitats are widely recognised as blue carbon ecosystems with mitigation potential (Chmura et al., 2003¹⁶⁶⁸; Duarte et al., 2005¹⁶⁶⁹; Kennedy et al., 2010¹⁶⁷⁰; McLeod et al., 2011¹⁶⁷¹). Although covering only ~0.1% of the Earth’s surface, these three ecosystems together have been estimated to support 1–10% of global marine primary production (Duarte et al., 2017¹⁶⁷²). More than 150 countries contain at least one of these ecosystems; 71 countries contain all three (Herr and Landis, 2016¹⁶⁷³), and 74 countries mention such coastal wetlands (five specifically as blue carbon) in their Nationally Determined Contributions (NDCs) to the Paris Agreement (Martin et al., 2016a¹⁶⁷⁴; Gallo et al., 2017¹⁶⁷⁵).

These three vegetated coastal habitats are characterised by high, yet variable, organic carbon storage in their soils and sediments on a per unit area basis (high confidence). In the humid tropics, mangrove below-ground organic carbon is typically 500–1000 tC ha–1 (Donato et al., 2011¹⁶⁷⁶; Alongi and Mukhopadhyay, 2015¹⁶⁷⁷; Howard et al., 2017¹⁶⁷⁸)), although only ~50 tC ha–1 in arid regions (Almahasheer et al., 2017¹⁶⁷⁹). Australian salt marshes show particularly wide variation in organic carbon storage, ranging from 15–1000 tC ha–1 (top 1 m) with mean of 165 tC ha–1 (Kelleway et al., 2016¹⁶⁸⁰; Macreadie et al., 2017b¹⁶⁸¹). For seagrass meadows, storage values are typically 400–1600 tC ha–1 but can exceed 2000 tC ha–1 (Serrano et al., 2014¹⁶⁸²). These accumulations have occurred over decadal to millennial time scales (McKee et al., 2007¹⁶⁸³; Lo Iacono et al., 2008¹⁶⁸⁴). Such blue carbon stock values are similar to freshwater wetlands and peat, but higher than for most forest soils (Laffoley and Grimsditch, 2009¹⁶⁸⁵; Pan et al., 2011¹⁶⁸⁶) (high confidence). 

When vegetated coastal ecosystems are disturbed, a proportion of their stored carbon is released back to the atmosphere, along with other greenhouse gases (Marba and Duarte, 2009¹⁶⁸⁶; Duarte et al., 2010¹⁶⁸⁶; Pendleton et al., 2012¹⁶⁸⁸; Lovelock et al., 2017¹⁶⁸⁹). Globally, around 25–50% of vegetated coastal habitats have already been lost or degraded due to coastal agricultural developments, urbanisation and other human disturbance during the past 100 years (McLeod et al., 2011¹⁶⁹⁰). The highest historical losses (60–90%) have occurred in Europe and China (Jickells et al., 2015¹⁶⁹¹; Gu et al., 2018¹⁶⁹²; Li et al., 2018a¹⁶⁹³). Current losses are estimated at 0.2–3.0% yr-1, depending on vegetation type and location (FAO et al., 2014; Alongi and Mukhopadhyay, 2015¹⁶⁹⁴; Atwood et al., 2017¹⁶⁹⁵) (medium confidence). Associated global carbon emissions are estimated at 0.04–0.28 GtC yr–1  (Pendleton et al., 2012¹⁶⁹⁶); 0.06–0.61 GtC yr–1 (Howard et al., 2017¹⁶⁹⁷); 0.10–1.46 GtC yr–1 (Lovelock et al., 2017¹⁶⁹⁸); and 0.007 GtC yr–1 (mangroves only) (Taillardat et al., 2018¹⁶⁹⁹). This range of values reflects uncertainties regarding the global rate of habitat loss, and the proportion of carbon remineralised to CO2.

Mitigation through emission reduction can therefore be achieved by habitat protection, to greatly reduce or end the human-driven loss of mangrove, salt marsh and seagrass ecosystems. Such action could potentially produce nationally-significant mitigation (>1% of fossil fuel emissions) for several countries (Taillardat et al., 2018¹⁷⁰⁰). However, there are still many uncertainties in quantifying carbon release due to habitat degradation and loss (Lovelock et al., 2017¹⁷⁰¹), and hence in determining emission reductions. Furthermore, this mitigation option is not available to those countries where habitat loss is not currently occurring, for example, in Bangladesh (Taillardat et al., 2018¹⁷⁰²). Since legal structures already exist in many countries to protect coastal wetlands, the main policy need may be the enforcement of national regulation and site-specific MPAs (Miteva et al., 2015¹⁷⁰³; Herr et al., 2017¹⁷⁰⁴; Howard et al., 2017¹⁷⁰⁵).

The alternative mitigation approach using coastal blue carbon ecosystems is to enhance the natural carbon uptake of such habitats, not only by increasing their spatial coverage through habitat restoration and new habitat creation, but also by taking management measures to maximise the carbon uptake and storage for existing coastal ecosystems. Such measures include reducing anthropogenic nutrient inputs and other pollutants; restoring hydrology, by removing barriers to tidal flow and sediment delivery; and reinstating predators (to reduce carbon loss caused by some bioturbators) (Macreadie et al., 2017a¹⁷⁰⁶). Per unit area of habitat created, restored or rehabilitated, such actions may offer high rates of carbon removal: widely-quoted values are 226 ± 39 gC m-1 yr-1 for mangroves, 218 ± 24 gC m-1 yr-1 for salt marsh and 138 ± 38 gC m-1 yr-1 for seagrass ecosystems (McLeod et al., 2011¹⁷⁰⁷; Isensee et al., 2019¹⁷⁰⁸). 

Around 90 restoration and rehabilitation projects for mangroves have been documented (López-Portillo et al., 2017¹⁷⁰⁹), with associated development of a range of restoration evaluation methods (Zhao et al., 2016a¹⁷¹⁰). Salt marsh restoration is reviewed by Adam (2019)¹⁷¹¹ and seagrass restoration by van Katwijk et al. (2016). Consistent conclusions, supported by other studies (Bayraktarov et al., 2016¹⁷¹²; Wylie et al., 2016¹⁷¹³) are that: natural regeneration increases the likelihood of longterm survival; higher success rates are achieved with strong stakeholder engagement; and it is critical that the (human) factors causing original loss and degradation have been properly addressed (high confidence).

Quantification of the climatic benefits of such actions is, however, not straightforward. Measurements of carbon burial rates show high site-specific variability, being strongly affected by a wide range of environmental factors for mangroves (Adame et al., 2017¹⁷¹⁴; Schile et al., 2017¹⁷¹⁵), seagrasses (Lavery et al., 2013¹⁷¹⁶) and salt marshes (Kelleway et al., 2017b¹⁷¹⁷). The reliable determination of sediment accumulation rates is a key consideration, with associated uncertainties not fully reflected in the McLeod et al. (2011)¹⁷¹⁸ estimates given above. In particular, geochemical-based studies have indicated that seagrass carbon burial may have been greatly overestimated (Johannessen and Macdonald, 2016¹⁷¹⁹). These issues are contentious (Johannessen and Macdonald, 2018a¹⁷²⁰; Johannessen and Macdonald, 2018b¹⁷²¹; Macreadie et al., 2018¹⁷²²; Oreska et al., 2018¹⁷²³); their scientific resolution is highly desirable. Additional complexities relating to the mitigation role of coastal blue carbon ecosystems include the following:

• Emissions of other greenhouse gases also need to be taken into account (Keller, 2019b¹⁷²⁴). Methane release from mangrove habitats can reduce the scale of their climatic benefits by 18–22% (Adams et al., 2012¹⁷²⁵; Chen and Ganapin, 2016¹⁷²⁶; Chmura et al., 2016¹⁷²⁷; Rosentreter et al., 2018¹⁷²⁸; Cameron et al., 2019¹⁷²⁹) and nitrous oxide and methane together may offset salt marsh CO2 uptake by 24–31% (Adams et al., 2012¹⁷³⁰). Nitrous oxide emissions are strongly affected by nutrient loading (Chmura et al., 2016¹⁷³¹); under pristine conditions, mangroves can provide a sink rather than a source (Maher et al., 2016¹⁷³²). Note that values of the ‘offset’ depend on the metrics used for determining CO2 equivalents.
• Carbonate formation, releasing CO2, may also reduce the benefits of carbon storage by similar proportions (Howard et al., 2017¹⁷³³; Macreadie et al., 2017a¹⁷³⁴; Kennedy et al., 2018¹⁷³⁵; Saderne et al., 2019¹⁷³⁶).
• Lateral transfers are not well-quantified. Whilst some of the carbon stored in coastal marine sediments may be recalcitrant carbon from terrestrial or atmospheric sources (and should therefore be excluded) (Chew and Gallagher, 2018¹⁷³⁷), export of dissolved organic carbon, inorganic carbon and alkalinity may be considered as additional sequestration (Maher et al., 2018¹⁷³⁸; Santos et al., 2019¹⁷³⁹). 
• The permanence of vegetated coastal systems, even if well-protected, cannot be assumed under future temperature regimes (Ward et al., 2016¹⁷⁴⁰; Duke et al., 2017¹⁷⁴¹; Jennerjahn et al., 2017¹⁷⁴²; Nowicki et al., 2017¹⁷⁴³)
• Responses to future SLR are also uncertain and complex (Kirwan and Megonigal, 2013¹⁷⁴⁴; Spencer et al., 2016¹⁷⁴⁵) . However, impacts are not necessarily negative: carbon sequestration capacity may increase where totally new habitats are created (Barnes, 2017¹⁷⁴⁶), or if mangroves replace salt marshes (Kelleway et al., 2016¹⁷⁴⁷). 

In summary, a combination of both conservation and restoration of mangrove, salt marsh and seagrass habitats can contribute to national mitigation effort for those countries with relatively large coastlines where such ecosystems naturally occur (Murdiyarso et al., 2015¹⁷⁴⁸; Atwood et al., 2017¹⁷⁴⁹). However, the associated current uncertainties in quantifying relevant carbon storage and flows are expected to be problematic for reliable measurement, reporting and verification (high confidence).

At the global scale, synthesis studies have estimated the potential additional sequestration achieved by cost effective coastal blue carbon restoration as ~0.05 GtC yr–1 (Griscom et al., 2017¹⁷⁵⁰) and 0.04 GtC yr–1 (National Academies of Sciences, Engineering, and Medicine, 2019), assuming that a relatively high proportion of vegetated ecosystems can be re-instated to their 1980–1990 extents. These values compare to current net anthropogenic emissions from all sources of 10.0 GtC yr–1 (Le Quéré et al., 2018), and are consistent with the ‘very low’ scores by (Gattuso et al., 2018) for the climate mitigation benefits of conserving and restoring coastal vegetation (Figure 5.23). Coastal ecosystem restoration could theoretically achieve higher sequestration, around ~0.2 GtC yr-1 (Griscom et al., 2017¹⁷⁵²), but would be challenging, because of the semi-permanent and on-going nature of most coastal land-use change, such as human settlement, conversion to agriculture and aquaculture, shoreline hardening and port development (Gittman et al., 2015¹⁷⁵³; Li et al., 2018a¹⁷⁵⁴). 

Restoration costs could also be an important constraint for large-scale application. Based on published data from 246 observations, Bayraktarov et al. (2016)¹⁷⁵⁵ estimated median total costs for restoration of one hectare of mangrove, salt marsh and seagrass habitat to be ~2,508, 151,129 and 383,672 respectively, in 2010 USD. For each ecosystem, there was high variability in costs according to the economy of the country where the restoration projects were carried out, and the restoration technique applied. Assessment of coastal conservation and restoration costs is also given in Section 4.4.2.3, in Box 5.5 (in the context of coral reef restoration costs) and Section 5.5.2.5. 

Measures to protect and restore coastal blue carbon habitats provide many other societal benefits in addition to climate regulation (Section 5.4.1). In particular, there is high confidence that coastal wetlands benefit local fisheries, enhance biodiversity, give storm protection, reduce coastal erosion, improve water quality and support local livelihoods (Costanza et al., 2008¹⁷⁵⁶; Spalding et al., 2014¹⁷⁵⁷). Coastal ecosystems may keep pace with sufficiently gradual SLR, and may be more cost-effective in flood protection than hard infrastructure like seawalls (Temmerman et al., 2013¹⁷⁵⁸; Möller, 2019¹⁷⁵⁹). Coastal blue carbon can therefore be considered as a ‘no regrets’ mitigation option at the national level in many countries, in addition to (not a replacement for) more effective mitigation measures. Additional research is needed over the full range of environmental conditions to improve knowledge and understanding of the complex carbon dynamics of coastal vegetation and associated systems, to enable well-quantified and cost-effective carbon sequestration enhancement (Vázquez-González et al., 2017¹⁷⁶⁰; Windham-Myers et al., 2019¹⁷⁶¹).

Seaweeds do not directly transfer carbon to marine sediments, unlike the rooted coastal vegetation considered above (Howard et al., 2017¹⁷⁶²). Nevertheless, seaweed detritus can deliver carbon to sedimentary sites (Hill et al., 2015¹⁷⁶³) and may provide a source of refractory dissolved organic (Krause-Jensen and Duarte, 2016¹⁷⁶⁴). Recent studies indicate that globally important amounts of carbon may be involved in these processes (Krause-Jensen and Duarte, 2016¹⁷⁶⁵; Krause-Jensen et al., 2018¹⁷⁶⁶; Smale et al., 2018¹⁷⁶⁷). There is, however, currently low confidence that enhancement of natural seaweed production can provide a significant mitigation response, due to large uncertainties relating to sequestration duration and effectiveness. Such considerations relate to transport pathways, the fate of material transported to deeper water, and the timescales of its subsequent return to the atmosphere over decadal to century timescales. 

Seaweed aquaculture is inherently more manageable as a mitigation response (N‘Yeurt et al., 2012¹⁷⁶⁸; Chung et al., 2013¹⁷⁶⁹; Chung et al., 2017¹⁷⁷⁰; Duarte et al., 2017¹⁷⁷¹). If linked to biofuel or biogas production (N‘Yeurt and Iese, 2014¹⁷⁷²; Moreira and Pires, 2016¹⁷⁷³; Sondak et al., 2017¹⁷⁷⁴), there would be potential to reduce emissions (as an alternative to fossil fuels); if also linked to carbon capture and storage (Hughes et al., 2012¹⁷⁷⁵), it may be possible to achieve negative emissions (net CO2 removal from the atmosphere). Full life cycle analyses are needed to assess the energy efficiency of such approaches, and the viability of scaling them up to climatically-important levels, taking account of associated environmental and socioeconomic implications.

A different mitigation option using seaweeds relates to their use as a dietary supplement for ruminants to suppress methane production. In vitro studies have given promising results (Dubois et al., 2013¹⁷⁷⁶; Machado et al., 2016¹⁷⁷⁷; Machado et al., 2018¹⁷⁷⁸). However, because the potential scale of real-world benefits have yet to be quantified, there is low confidence in this approach as a mitigation option. 

Land-based nutrient management could, in theory, be used to enhance carbon storage in coastal seas and deeper waters, by increasing the amount of refractory dissolved organic carbon (Jiao et al., 2011¹⁷⁷⁹; Jiao et al., 2014b¹⁷⁸⁰; Jiao et al., 2018b¹⁷⁸¹). This idea is supported by a statistical analysis of the relationship between organic carbon and nitrate in various natural environments (Taylor and Townsend, 2010¹⁷⁸²) as well as by experimental results in estuarine and offshore waters (Yuan et al., 2010¹⁷⁸³; Jiao et al., 2011¹⁷⁸⁴; Jiao et al., 2014b¹⁷⁸⁵). Delivery of nutrients from agricultural fertilisers and sewage discharge to coastal waters may currently promote the microbial breakdown of river-derived terrestrial dissolved organic carbon, reducing carbon storage (Liu et al., 2014¹⁷⁸⁶). Thus reducing nutrient inputs in the future may expand carbon storage by favouring the microbial carbon pump, in addition to the multiple co-benefits of reduced nutrient loads related to HABs, oxygenation and ocean acidification (Miranda et al., 2013¹⁷⁸⁷; Jiao et al., 2018a¹⁷⁸⁸; Zhang et al., 2018¹⁷⁸⁹). Although there is some evidence for the impact of dissolved organic carbon variations on global scale climate (Rothman et al., 2003¹⁷⁹⁰) the benefits of this approach have yet to be determined quantitatively and uncertainties remain regarding the longevity of removal and associated carbon accounting (measurement, reporting and verification). Until such issues are better resolved, there is low confidence that stimulation of refractory dissolved organic carbon production could provide an operational long-term mitigation measure.

Anthropogenic sediment disturbance, through fishing, dredging and the installation of offshore structures, affects the security of carbon storage in shelf sea sediments (Hale et al., 2017¹⁷⁹¹). Management of such activities might therefore increase carbon retention, over relatively large areas of shelf seas (Avelar et al., 2017¹⁷⁹²; Luisetti et al., 2019¹⁷⁹³). However, there is a lack of data and understanding of the complex processes that affect carbon storage in the potentially mobile fraction of marine sediments (van de Velde et al., 2018¹⁷⁹⁴); exceptions are provided by Hu et al. (2016)¹⁷⁹⁵ and Diesing et al. (2017)¹⁷⁹⁶. Due to these uncertainties, there is currently low confidence that control of sediment disturbance can be used for climate mitigation.

There is theoretically greater potential for carbon removal by ‘enhanced weathering’ using mineral additions to coastal waters (and the open ocean) (Rau, 2011; Renforth and Henderson, 2017¹⁷⁹⁷). These approaches are based on increasing the naturally-occurring uptake of CO2 by carbonates (e.g., calcite and dolomite) or silicate minerals (such as olivine). Such rock-weathering currently sequesters ~0.25 GtC yr–1, on land and at sea (Taylor et al., 2015) and provides the longterm control of atmospheric CO2 concentrations. It could be enhanced by adding ground minerals to beaches (Montserrat et al., 2017¹⁸⁰¹) or the sea surface. Other geochemical approaches for adding alkalinity that are less directly based on natural processes (Rau et al., 2012¹⁸⁰²; GESAMP, 2019¹⁸⁰³) are not considered here.

Enhanced weathering methods might be used to reduce local impacts, for example, for coral reefs (Albright et al., 2016b¹⁷⁹⁸; Feng et al., 2016¹⁷⁹⁹), as well as contributing to wider mitigation of climate change. However, their climatic benefits would be difficult to quantify, with other constraints on their development and deployment relating to the governance, cost and uncertain environmental impacts of large-scale application (Gattuso et al., 2018¹⁸⁰⁰). The combination of these factors results in low confidence that enhanced weathering can provide a viable and acceptable climate mitigation approach.

There are many studies on biological climate change adaptation responses (Crozier and Hutchings, 2014¹⁸⁴⁴; Miller et al., 2017¹⁸⁴⁵; Diamond, 2018¹⁸⁴⁶). Sections 5.2.3 and 5.3.3 discuss three main types of biological adaptation, broadly defined: evolutionary (genetic) adaptation through natural selection; phenotypic plasticity (acclimatisation), within an organism’s lifetime; and individual or population mobility towards more favourable conditions. There are, however, expected to be limits to such natural adaptation, and large variations between species and populations (Gienapp and Merilä, 2018¹⁸⁴⁷). 

An accurate understanding of climate change impacts upon species, their sensitivity and adaptive capacity and consequent ecological effects (considering both indirect as well as direct impacts) is used to estimate extinction risk, so that an appropriate management response can be developed (Butt et al., 2016¹⁸⁴⁸). EbA takes these complex interactions into account (Hobday et al., 2015¹⁸⁴⁹), including the disruptive impacts of alien invasive species (Ondiviela et al., 2014¹⁸⁵⁰; Wigand et al., 2017¹⁸⁵¹) . Effective adaptation action, therefore, contains a broader consideration than historical conservation practices (medium evidence, high agreement), including the development of international collaborations and databases to improve ocean-scale understanding of climate change impacts (Okey et al., 2014¹⁸⁵²; Young et al., 2015¹⁸⁵³). A key knowledge gap relates to the critical thresholds for irreversible change for species (Powell et al., 2017¹⁸⁵⁴).

Coral reefs are currently threatened by the continuous global degradation of warm water coral reef ecosystems and the failure of traditional conservation actions to revive most of the degrading reefs (Rinkevich, 2008¹⁸⁵⁵; Miller and Russ, 2014¹⁸⁵⁶). Interventions to rehabilitate degraded coral reef ecosystems can be categorised as preventive (‘passive’ restoration) or adaptive (‘active’ restoration) (Miller and Russ, 2014¹⁸⁵⁷; Linden and Rinkevich, 2017¹⁸⁵⁸) (see Box 5.5). 

Inspired by silviculture (forestation) approaches to terrestrial ecosystem restoration, studies (Rinkevich, 1995¹⁸⁵⁹; Rinkevich, 2005¹⁸⁶⁰; Rinkevich, 2006¹⁸⁶¹; Rinkevich, 2008¹⁸⁶²; Bongiorni et al., 2011¹⁸⁶³) have proposed a two step restoration strategy for warm water coral reefs termed gardening of denuded coral reefs. In the first step, a large pool of coral colonies (derived from coral nubbins and fragments, and from sexually derived spat) are farmed in underwater nurseries, preferably on mid-water floating devices installed in sheltered zones, in which coral material can be cultured for up to several years. In the second step, nursery-grown coral colonies, together with recruited associated biota, are transplanted to degraded reef sites (Shafir and Rinkevich, 2008¹⁸⁶⁴; Mbije et al., 2010¹⁸⁶⁵; Shaish et al., 2010b¹⁸⁶⁶; Shaish et al., 2010a¹⁸⁶⁷; Bongiorni et al., 2011¹⁸⁶⁸; Horoszowski-Fridman et al., 2011¹⁸⁶⁹; Linden and Rinkevich, 2011¹⁸⁷⁰; Mbije et al., 2013¹⁸⁷¹; Cruz et al., 2014¹⁸⁷²; Chavanich et al., 2015¹⁸⁷³; Horoszowski-Fridman et al., 2015¹⁸⁷⁴; Lirman and Schopmeyer, 2016¹⁸⁷⁵; Montoya Maya et al., 2016¹⁸⁷⁶; Ng et al., 2016; Lohr and Patterson, 2017; Rachmilovitz and Rinkevich, 2017¹⁸⁷⁸). Active restoration of coral reefs, while still in its infancy and facing a variety of challenges (Rinkevich, 2015b¹⁸⁷⁹; Hein et al., 2017¹⁸⁸⁰), has been suggested to potentially improve the ecological status of degraded coral reefs and the socioeconomic benefits that the reefs provide (Rinkevich, 2014¹⁸⁸¹; Rinkevich, 2015b¹⁸⁸²; Linden and Rinkevich, 2017¹⁸⁸³). 

Ecological engineering approaches may promote coral reef adaptation (Rinkevich, 2014¹⁸⁸⁴; Forsman et al., 2015¹⁸⁸⁵; Coelho et al., 2017¹⁸⁸⁶; Horoszowski-Fridman and Rinkevich, 2017¹⁸⁸⁷; Linden and Rinkevich, 2017¹⁸⁸⁸; Rachmilovitz and Rinkevich, 2017¹⁸⁸⁹). They also include: augmenting functional diversity, including that of the microbiome (Casey et al., 2015¹⁸⁹⁰; Horoszowski-Fridman and Rinkevich, 2017¹⁸⁹¹; Shaver and Silliman, 2017¹⁸⁹²); transplantating whole habitats (Shaish et al., 2010b¹⁸⁹³; Gómez et al., 2014¹⁸⁹⁴); and enhancing genetic diversity (Iwao et al., 2014¹⁸⁹⁵; Drury et al., 2016¹⁸⁹⁶; Horoszowski-Fridman and Rinkevich, 2017¹⁸⁹⁷). Active restoration can contribute to reef rehabilitation in all major reef regions (Rinkevich, 2014¹⁸⁹⁸; Rinkevich, 2015b¹⁸⁹⁹). However, there is limited evidence on how resistant these manipulated corals are to global change drivers (Shaish et al., 2010b¹⁹⁰⁰; Shaish et al., 2010a¹⁹⁰¹) or how the nursery time affects biological traits like reproduction in coral transplants (Horoszowski-Fridman et al., 2011¹⁹⁰²). Coral epigenetics may also be used as an adaptive management tool for reef rehabilitation (low confidence), as suggested by studies on coral adaptation (Brown et al., 2002¹⁹⁰³; Horoszowski-Fridman et al., 2011¹⁹⁰⁴; Palumbi et al., 2014¹⁹⁰⁵; Putnam and Gates, 2015¹⁹⁰⁶; Putnam et al., 2016¹⁹⁰⁷). 

Research on active coral reef restoration (Box 5.5) suggests the potential to help rehabilitate degraded coral reefs, provided that the underlying drivers of the impacts are mitigated (high confidence). Ongoing and new research in active coral reef restoration may further improve active reef restoration outcomes (Box 5.5) (low confidence). However, these coral reef restoration options may be ineffectual if global warming exceeds 1.5oC relative to pre-industrial levels (Hoegh-Guldberg et al., 2018¹⁹⁰⁸; IPCC, 2018¹⁹⁰⁹).  

Mangroves provide significant ecosystem services, including localised coastal protection from extreme storm events (Section 5.4.1), supporting services through increased sedimentation rates (Hayden and Granek, 2015¹⁹⁶⁰) and provisioning services for local communities, for example, habitats for nurseries to support fisheries. Mangroves provide limited carbon mitigation, in terms of global emissions reduction, and substantial job creation (Table 5.7) co-benefits (for example through Reducing Emissions from Deforestation and Forest Degradation programmes) when managed properly (Section 5.4.1, 5.5.1.1), and there is evidence of their value in supporting aquaculture and fishery initiatives (Huxham et al., 2015¹⁹⁶¹; Ahmed and Glaser, 2016a¹⁹⁶²). 

Mangrove EbA responses most commonly reported included ecosystem restoration (Sierra-Correa and Cantera Kintz, 2015¹⁹⁶³; Romañach et al., 2018¹⁹⁶⁴) and management such as mangroves re-planting through community participation programmes (Nanlohy et al., 2015¹⁹⁶⁵; Nguyen et al., 2017¹⁹⁶⁶; Triyanti et al., 2017¹⁹⁶⁷). Mangrove EbA has been reported to provide multiple co-benefits in terms of improvement in support for coastal physical processes, including: shoreline stabilisation (Hayden and Granek, 2015¹⁹⁶⁸; Nanlohy et al., 2015¹⁹⁶⁹); ecological functioning (Sierra-Correa and Cantera Kintz, 2015¹⁹⁷⁰; Miller et al., 2017¹⁹⁷¹) with improved ecosystem services (Alongi, 2015¹⁹⁷²; Nanlohy et al., 2015¹⁹⁷³; Palacios and Cantera, 2017¹⁹⁷⁴); carbon mitigation (5.5.1.1); supporting livelihoods (Nanlohy et al., 2015¹⁹⁷⁵; Nguyen et al., 2017¹⁹⁷⁶); and reductions in coastal infrastructure damage and community vulnerability to climate change impacts. Managed retreat to counter coastal squeeze (Section 5.3) through improved governance, creation of finance and land use planning can allow mangroves to move up the shoreline contour or down the latitudinal gradient (Sierra-Correa and Cantera Kintz, 2015¹⁹⁷⁷; Ward et al., 2016¹⁹⁷⁸; Romañach et al., 2018¹⁹⁷⁹). Therefore, mangrove EbA responses can strengthen coastal ecosystem services through shoreline stabilisation and provide multiple co-benefits for coastal communities, like job creation and improved access to ecosystem services (high confidence).

There are, however, examples where community mangrove restoration projects have resulted in maladaptive outcomes, in which the resulting ecosystem degradation could not provide the ecosystem services required (Nguyen et al., 2017¹⁹⁸⁰; Romañach et al., 2018¹⁹⁸¹). Such maladaptation can be a result of poor governance processes or a lack of community compliance with restoration plans. These examples emphasise the value of designing effective governance to implement adaptation responses with broad community participation to improve the climate risk reduction outcomes and co-benefits (Sierra-Correa and Cantera Kintz, 2015¹⁹⁸²; Nguyen et al., 2017¹⁹⁸³) (medium evidence, high agreement).

Mangrove and other coastal ecosystems restoration and management can be applied through reducing non-climatic hazards (Gilman et al., 2008¹⁹⁸⁴; Ataur Rahman and Rahman, 2015¹⁹⁸⁵; Sierra-Correa and Cantera Kintz, 2015¹⁹⁸⁶; Ahmed and Glaser, 2016a¹⁹⁸⁷; Nguyen et al., 2017¹⁹⁸⁸; Romañach et al., 2018¹⁹⁸⁹). Coastal and catchment development, including wetland transformation and degradation (Miloshis and Fairfield, 2015¹⁹⁹⁰; Schaeffer-Novelli et al., 2016¹⁹⁹¹; Watson et al., 2017a¹⁹⁹²; Schuerch et al., 2018¹⁹⁹³), the disruption of physical processes impacting sedimentation rates (Watson et al., 2017a¹⁹⁹⁴) and coastal squeeze compound coastal climate change impacts like erosion, flooding and saltwater intrusion (Ondiviela et al., 2014¹⁹⁹⁵; Miloshis and Fairfield, 2015¹⁹⁹⁶; Schaeffer-Novelli et al., 2016¹⁹⁹⁷; Wigand et al., 2017¹⁹⁹⁸) (Section 5.3). This reduces the ability of these ecosystems to provide protection from wave and storm impacts, whilst positive feedbacks may occur that cause a net release of carbon into the atmosphere, for example, in salt marshes (Wong et al., 2014a¹⁹⁹⁹) (Section 5.4.1). In some cases, effective interventions requires management at a broad spatial scale that includes a variety of ecosystems, for example, including ecosystems like mussel beds on the seaward side of seagrass beds to reduce wave energy and erosion (Ondiviela et al., 2014²⁰⁰⁰). Where sediment accretion matches the SLR rate, wetlands and salt marshes provide effective coastal protection and other important ecosystem services (high confidence).

Coastal dune systems are widely transformed globally. Human disturbance and the limited stabilising ability of dune vegetation are key causes of degradation (Onaka et al., 2015²⁰⁰¹; Ranasinghe, 2016²⁰⁰²; MacDonald et al., 2017²⁰⁰³; Pranzini, 2017²⁰⁰⁴; Salgado and Martinez, 2017²⁰⁰⁵; Vikolainen et al., 2017²⁰⁰⁶; Gracia et al., 2018²⁰⁰⁷), while restoration efforts can be supported by both hard (Sutton-Grier et al., 2015²⁰⁰⁸; Pranzini, 2017²⁰⁰⁹) and soft (Sutton-Grier et al., 2015²⁰¹⁰; Vikolainen et al., 2017²⁰¹¹) engineering responses. Reduced coastal erosion (Sánchez-Arcilla et al., 2016²⁰¹²; Goreau and Prong, 2017²⁰¹³; Vikolainen et al., 2017²⁰¹⁴; Carro, 2018²⁰¹⁵; Gracia et al., 2018²⁰¹⁶) and flood risk (Onaka et al., 2015²⁰¹⁷; MacDonald et al., 2017²⁰¹⁸; Nehren et al., 2017²⁰¹⁹) through maintaining dunes as natural buffers against wave energy (Nehren et al., 2017²⁰²⁰) can increase resilience to climate change impacts (Sutton-Grier et al., 2015²⁰²¹; Magnan and Duvat, 2018²⁰²²). Engineered responses and sand replenishment are considered complementary approaches (Onaka et al., 2015²⁰²³; Martínez et al., 2017²⁰²⁴). Section 4.4.4.1 provides an overview of sediment-based adaptation response measures, including cost estimates for beach nourishment and dune maintenance, a discussion of co-benefits and drawbacks of combining hard and soft infrastructure measures, and challenges with sourcing sediment for beach replenishment. In some cases dune restoration and sand replenishment projects have not been successful, due to fire damage (Shumack and Hesse, 2017²⁰²⁵) or the rapid loss of sand within replenishment schemes due to coastal processes and stakeholder rejection of adaptation activities (Pranzini, 2017²⁰²⁶). Coastal dune restoration and beach replenishment are effective responses against coastal erosion and flooding, where sufficient materials and space to implement are available (medium confidence).

There is a growing body of literature regarding the effectiveness and economics of EbA. In addition to building resilience to climate change, EbA is expected to bring a wide range of co-benefits that include increasing ecological complexity, with multiple ecosystem services, and other economic co-benefits (Perkins et al., 2015²⁰²⁷; Perry, 2015²⁰²⁸; Moosavi, 2017²⁰²⁹; Scarano, 2017²⁰³⁰). The cost-effectiveness of EbA approaches varies between marine ecosystem types; for example, coral reefs (Perkins et al., 2015²⁰³¹; Beetham et al., 2017²⁰³²; Elliff and Silva, 2017²⁰³³; Beck et al., 2018²⁰³⁴; Comte and Pendleton, 2018²⁰³⁵) and salt-marshes (Ondiviela et al., 2014²⁰³⁶; Miloshis and Fairfield, 2015²⁰³⁷; Schaeffer-Novelli et al., 2016²⁰³⁸; Wigand et al., 2017²⁰³⁹) performed best at reducing wave heights, whilst salt marshes and mangroves were two to five times cheaper than submerged breakwaters for wave heights of less than half a meter. Although low regrets, win-win approaches like EbA are supported in the literature (Watkiss et al., 2014²⁰⁴⁰; Barange et al. 2018²⁰⁴¹), syntheses of experience from context-specific practical implementation of EbA and assessment of their cost-effectiveness are limited (Narayan et al., 2016). Therefore, EbA can be a cost-effective approach for securing climate change-related ecosystem services with multiple co-benefits (medium evidence, high agreement).