Editing IPCC:AR6/SROCC/Chapter-5 (section)

=== 5.2.4 Impacts on Deep Seafloor Systems ===

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<span id="changes-on-the-deep-seafloor"></span>
==== 5.2.4.1 Changes on the Deep Seafloor ====

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The deep seafloor is assessed here as the vast area of the ocean bottom &gt;200 m deep, beyond most continental shelves (Levin and Sibuet, 2012; Boyd et al., 2019) (Figure 5.15). Below 200 m changes in light, food supply and the physical environment lead to altered benthic (seafloor) animal taxonomic composition, morphologies, lifestyles and body sizes collectively understood to represent the deep sea (Tyler, 2003).

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Most deep seafloor ecosystems globally are experiencing rising temperatures, declining oxygen levels, and elevated CO 2 , leading to lower pH and carbonate undersaturation (WGII AR5 30.5.7; Section 5.2.2.3). Small changes in exposure to these hazards by deep seafloor ecosystem have been confirmed by observation over the past 50 years. However, analysis using direct seafloor observations of these hazards over the past 15-29 years suggest that the environmental conditions are highly variable over time because of the strong and variable influences by ocean conditions from the sea surface (Frigstad et al., 2015; Thomsen et al., 2017). Such high environmental variability makes it difficult to attribute observed trends to anthropogenic drivers using existing datasets (Smith et al., 2013; Hartman et al., 2015; Soltwedel et al., 2016; Thomsen et al., 2017) ( ''high confidence'' ). Projections from global ESMs suggest large changes for temperature by 2100 and beyond under RCP8.5 (relative to present day variation) (Mora et al., 2013; Sweetman et al., 2017; FAO, 2019). The magnitude of the projected changes is lower under RCP2.6, and in some cases the direction of projected change to 2100 varies regionally under either scenario (FAO, 2019) ( ''high confidence'' ).

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<span id="open-ocean-seafloor---abyssal-plains-3000-6000-m"></span>
==== 5.2.4.2 Open Ocean Seafloor - Abyssal Plains (3000-6000 m) ====

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Abyssal communities (3000–6000 m) cover over 50% of the ocean’s surface and are considered to be extremely food limited (Gage and Tyler, 1992; Smith et al., 2018 <sup>[[#fn:r684|684]]</sup> ). There is a strong positive relationship between surface primary production, export flux, and organic matter supply to the abyssal seafloor (Smith et al., 2008 <sup>[[#fn:r685|685]]</sup> ), with pulses of surface production reflected as carbon input on the deep seafloor in days to months (Thomsen et al., 2017 <sup>[[#fn:r686|686]]</sup> ). Both vertical and horizontal transport contribute organic matter to the sea floor (Frischknecht et al., 2018 <sup>[[#fn:r687|687]]</sup> ). Food supply to the seafloor regulates faunal biomass, explaining the strong positive relationships documented between surface production and seafloor faunal biomass in the Pacific Ocean (Smith et al., 2013 <sup>[[#fn:r688|688]]</sup> ), Gulf of Mexico (Wei et al., 2011 <sup>[[#fn:r689|689]]</sup> ) and north Atlantic Ocean (Hartman et al., 2015 <sup>[[#fn:r690|690]]</sup> ). Extended time series and broad spatial coverage reveal strong positive relationship between annual POC flux and abyssal sediment community oxygen consumption (Rowe et al., 2008 <sup>[[#fn:r691|691]]</sup> ; Smith et al., 2016a <sup>[[#fn:r692|692]]</sup> ). Observed reduction in in POC flux at the abyssal seafloor enhances the relative importance of the microbial loop and reduces the importance of benthic invertebrates in carbon transfer (Dunlop et al., 2016 <sup>[[#fn:r693|693]]</sup> ) (single study, ''limited evidence'' ). However, changes in the overlying mesopelagic and bathypelagic communities (see Section 5.2.3.2) will also affect food flux to the deep seafloor, as nekton and zooplankton transfer energy to depth through diel (daily day-night) vertical migrations, ontogenetic (life staged-based) migrations and falls of dead carcasses (Gage, 2003 <sup>[[#fn:r694|694]]</sup> ). Therefore, climate change impacts on organic carbon export from the epipelagic (Section 5.2.3.1) and deeper pelagic systems (Section 5.2.3.2) can affect the energy available to support the abyssal seafloor ecosystems ( ''medium confidence'' ). However, because observations on historical changes in POC flux in abyssal seafloor ecosystems are limited to a few locations, long-term records show high variability, and mechanistic understanding of factors affecting the biological carbon pump is incomplete, there is ''limited evidence'' that the abyssal seafloor ecosystem has already been affected by changes in POC flux as a result of climate change. The metabolic rate of deep seafloor ectotherms, and consequently their demand for food, increases with temperature. Thus, observed warming in deep sea ecosystems (Hoegh-Guldberg et al., 2014 <sup>[[#fn:r695|695]]</sup> ) (Section 5.2.2.2.1) is expected to increase the sensitivity of deep seafloor biota to decrease in food supplies associated with a change in POC flux ( ''high confidence'' ). However, there is ''limited evidence'' of observed changes in abyssal biota. Small deep sea biota demonstrate increased efficiency (effective use of food energy for growth and metabolism with minimal loss) at low food inputs (due to small size and dominance by prokaryotic taxa) (Gambi et al., 2017 <sup>[[#fn:r696|696]]</sup> ). Adaptation to low food availability in abyssal ecosystems may confer higher capacity to adjust to reduced food availability than for shallow biota ( ''limited evidence).'' Overall, the risk of impacts of climate change on abyssal ecosystems through reduction in food supplies from declining POC flux in the present day is low with ''low confidence'' .

The globally integrated export flux of carbon is projected to decrease in the open ocean in the 21st century under RCP2.6 (by 1.6–4.9%) and RCP8.5 (by 8.9–15.8%) relative to 2000 ( ''medium confidence'' ) (Section 5.2.2.6). This change in export flux of carbon is projected to yield declines in POC flux at the abyssal seafloor (representing food supply to benthos) of up to –27% in the Atlantic and up to –31 to –40% in the Pacific and Indian Oceans, with some increases in polar regions (Sweetman et al. 2017 <sup>[[#fn:r697|697]]</sup> ). In some models, additional dissolution of calcium carbonate due to ocean acidification further lowers POC flux, causing the projected export production declines to be up to 38% at the northeast Atlantic seafloor (Jones et al., 2014 <sup>[[#fn:r699|699]]</sup> ). Lower POC fluxes to the abyss reduce food supply and have been projected to cause a size-shift towards smaller organisms (Jones et al., 2014), resulting in rising respiration rates, lower biomass production efficiency, and lesser energy transfer to higher trophic levels (Brown et al., 2004 <sup>[[#fn:r700|700]]</sup> ) ( ''medium confidence'' ). Changes are projected to be largest for macrofauna and lesser and similar for megafauna and meiofauna (Jones et al., 2014) ( ''limited evidence'' , ''low confidence'' ). Projections using outputs from seven CMIP5 models suggest that 97.8 ± 0.6% (95% CI) of the abyssal seafloor area will experience a biomass decline by 2091–2100 relative to 2006–2015 under RCP8.5. The projected decreases in overall POC flux to the abyssal seafloor are projected to cause a 5.2–17.6% reduction in seafloor biomass in 2090–2100, relative to 2006–2015 under RCP8.5 (Jones et al., 2014 <sup>[[#fn:r701|701]]</sup> ). The projected impacts on abyssal seafloor biomass are significantly larger under RCP8.5 than RCP4.5 (Jones et al., 2014 <sup>[[#fn:r703|703]]</sup> ). However, existing estimates are based on total POC flux changes and do not account for changes in the type or quality of the sinking material, to which macrofaunal and meiofaunal invertebrates are highly sensitive (Smith et al., 2008 <sup>[[#fn:r704|704]]</sup> ; Smith et al., 2009 <sup>[[#fn:r705|705]]</sup> ; Tittensor et al., 2011 <sup>[[#fn:r706|706]]</sup> ). The projections also do not account for direct faunal responses to changes in temperature, oxygen or the carbonate system, all of which will influence benthic responses to changing food availability (AR5 Chapter 30.5.7), reducing to ''medium confidence'' the risk assessment that is based on these projections (Figure 5.16).

Regionally, while reductions in POC flux are projected at low and mid latitudes in the Pacific, Indian and Atlantic Oceans, increases are projected at high latitudes associated in part with reduction in sea ice cover (Yool et al., 2013 <sup>[[#fn:r707|707]]</sup> ; Rogers, 2015 <sup>[[#fn:r708|708]]</sup> ; Sweetman et al., 2017 <sup>[[#fn:r709|709]]</sup> ; Yool et al. 2017 <sup>[[#fn:r710|710]]</sup> ; FAO 2019 <sup>[[#fn:r711|711]]</sup> ) (see Chapter 3) ( ''medium confidence'' ). Notably, Arctic and Southern Ocean POC fluxes at the abyssal seafloor are projected to increase by up to 38% and 21%, respectively by 2100 under RCP8.5 (Sweetman et al., 2017 <sup>[[#fn:r712|712]]</sup> ). While an increase in food supply may yield higher benthic biomass at high latitudes, warmer temperatures and reduced pH projected for the polar regions (Chapter 3) would elevate faunal metabolic demands, likely diminishing the benefit of elevated food supply to an unknown extent (Sweetman et al., 2017 <sup>[[#fn:r713|713]]</sup> ). Overall, given the limited food availability for fauna in the abyssal plains and the projected warming (Section 5.2.2.2.2) that increases the demand for food to support the elevated metabolic rates, the projected decrease in influx of organic matter and seafloor biomass will result in high risks of impacts to abyssal ecosystems by the end of the 21st century under RCP8.5 ( ''medium confidence'' ) (Figure 5.16). The risk of impacts is projected to be substantially lower under RCP4.5 or RCP2.6 ( ''high confidence'' ). The impacts on abyssal seafloor ecosystems affect functions that are important to support ecosystem services (see Section 5.4.1). For example, smaller-sized organisms exhibit reduced bioturbation intensity and depth of mixing causing reduced carbon sequestration (Smith et al., 2008 <sup>[[#fn:r714|714]]</sup> ) (Figure 5.15).

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'''Figure 5.15'''

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'''Figure 5.15 | A conceptual diagram illustrating how climate drivers are projected to modify deep sea ecosystems as discussed in Section 5.2.4.'''
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[[File:31366bf5606fb7d2d2844e775476da50 IPCC-SROCC-CH_5_15-3000x2465.jpg]]

Figure 5.15 | A conceptual diagram illustrating how climate drivers are projected to modify deep sea ecosystems as discussed in Section 5.2.4.

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==== 5.2.4.3 Bathyal Ecosystems (200–3000 m) ====

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Bathyal ecosystems consist of numerous geomorphic features with steep topography (Figure 5.15). These include continental slopes covering 5.2% of the seafloor, over 9400 steep-sided canyons, and &gt;9000 conical seamounts (submarine volcanos which are mainly inactive), as well as guyots and ridges which together cover ~6% of the seafloor (Harris et al., 2014) ''.'' Seamounts and canyons support high animal densities and biomass including cold water coral, sponge and bryozoan reefs, exhibit high secondary production supported by locally enhanced primary production and intensified water flow, function as diversity hotspots and serve as stepping stones for larval dispersal (Rowden et al., 2010). Canyons transport particulate organic matter, migrating plankton and coarse material from the shelf, and are sites where intensified mixing and advection of water masses occurs (De Leo et al., 2010; Levin and Sibuet, 2012; Fernandez-Arcaya et al., 2017).  Slopes, canyons and seamounts exhibit strong vertical temperature, oxygen and pH gradients generating sharp ecological zonation (Levin and Sibuet, 2012), thus changes in exposures are expected to alter the distributions of their communities (Figure 5.15, 5.16 '') (medium confidence).  ''

In some regions, observational records document changing conditions in bathyal ecosystems (Levin, 2018; Section 5.2.2.4). In the Northeast Pacific continental slopes associated with the California Current ecosystem, observations over the past 25 years show high variability but an overall trend of decreasing ocean oxygen and pH levels with oxygen declines of up to 40% and pH declines of 0.08 units in California . (Goericke et al., 2015) (high agreement, robust evidence, high confidence). Large oxygen declines are linked to past warming events on continental margins, over multiple time scales from 1–100 ky (Dickson et al., 2012; Moffitt et al., 2015). Studies across modern oxygen gradients on slopes reveal that suboxic (5–10 µMol kg -1 O 2 ) values lead to loss of biodiversity of fish (Gallo and Levin, 2016), invertebrates (Levin, 2003; Gallo and Levin, 2016; Sperling et al., 2016), and protozoans (Bernhard and Reimers, 1991; Gooday et al., 2000; Moffitt et al., 2014) (high confidence). Shoaling oxyclines on continental slopes have altered depth distributions of multiple co-occurring echinoid species over the past 25 years (Sato et al., 2017) and can reduce the growth rate, and change the skeletal structure and biochemical composition of a common sea urchin (Sato et al., 2018). In central Pacific oceanic canyons, fish abundance and diversity are reduced at 4 to 5 times higher oxygen concentrations than on continental slopes (&lt;31 µMol kg -1 O 2 ) (De Leo et al., 2012). Low oxygen on continental slopes causes reductions in faunal body size and bioturbation (Diaz and Rosenberg, 1995; Levin, 2003; Middelburg and Levin, 2009; Sturdivant et al., 2012), simplification of trophic structure reducing energy flow to upper trophic levels (Sperling et al., 2013), shifts in carbon processing pathways from metazoans to protozoans (Woulds et al., 2009), and reduced colonisation potential (Levin et al., 2013). These changes are expected to lead to altered ecosystem structure and function, with lower carbon burial (Smith et al., 2000; Levin and Dayton, 2009) (medium confidence). Both carbon sequestration and nitrogen recycling are highly sensitive to small changes in oxygenation within the suboxic zone (Deutsch et al., 2011).

Bathyal species adapted to OMZs where CO 2 levels are characteristically high, appear less vulnerable to the negative impacts of ocean acidification (Taylor et al., 2014). Benthic foraminifera, which are often the numerically dominant deep sea taxon, show no significant effect of short-term exposure to ocean acidification on survival of multiple species (Dissard et al., 2010; Haynert et al., 2011; Keul et al., 2013; McIntyre-Wressnig et al., 2014; Wit et al., 2016) and in fact hypoxia in combination with elevated pCO 2 favors survival of some foraminifera (Wit et al., 2016). However, lower pH exacerbates shallow foraminiferal sensitivity to warming (Webster et al., 2016). Limited evidence suggests that combined declines in pH and oxygen may lead to increase in some agglutinating taxa and a decrease in carbonate-producing foraminifera, including those using carbonate cement (van Dijk et al., 2017). Exposure to acidification (0.4 unit pH decrease) reduces fecundity and embryo development rate in a bathyal polychaete. Where both oxygen and CO 2 stress occur together on bathyal slopes, oxygen can be the primary driver of change (Taylor et al. 2014; Sato et al. 2018). Nematodes are sensitive to changes in temperature (Danovaro et al., 2001; Danovaro et al., 2004; Yodnarasri et al., 2008) and elevated CO 2 (Barry et al., 2004; Fleeger et al., 2006; Fleeger et al., 2010). There is low agreement about the direction of meiofaunal responses among studies, reflecting opposing responses in different regions. However, there is high agreement that meiofauna are sensitive to change in environment and food supply ( ''medium confidence'' ). Additional research is needed across all taxa on how hypoxia and pH interact (Gobler and Baumann, 2016).

Continental slopes, seamounts and canyons (200–2500 m) are projected to experience significant warming, pH decline, oxygen loss and decline in POC flux by 2081–2100 (compared to 1951–2000) under RCP8.5 (Table 5.5). In contrast, the average changes are projected to be 30–50% less under RCP2.6 (Table 5.5) by 2081-2100. Most ocean regions at bathyal depths (200–2500 m) except the Southern and Arctic Oceans are predicted to experience on average declining export POC flux under RCP8.5 by 2081–2100 (Yool et al., 2017; FAO, 2019) with the largest declines of 0.7–8.1 mg C m -2 d -1 in the Northeast Atlantic (FAO, 2019). There is a strong macroecological relationship between depth, export POC flux, biomass and zonation of macrobenthos on continental slopes (Wei et al., 2011), such that lower POC fluxes will alter seafloor community biomass and structure ( ''medium confidence'' ) (See also Section 5.2.4.1). This is modified on the local scale by near-bottom currents, which alter sediment grain size, food availability, and larval dispersal (Wei et al., 2011).

Declines in faunal biomass (6.1 ± 1.6% 95% C.I) are predicted for 96.6 ± 1.2% of seamounts under RCP8.5 by 2091–2100 relative to 2006–2015, driven by a projected 13.8 ± 3.3% drop in POC flux (Jones et al., 2014). The majority (85%) of mapped canyons are projected to experience comparable benthic biomass declines (Jones et al., 2014). By 2100 under RCP8.5, pH reductions exceeding -0.2 pH units are projected in ~ 23% of north Atlantic deep sea canyons and 8% of seamounts (Gehlen et al., 2014), with potential negative consequences for their cold water coral habitats (See Box 5.2).

Mean temperature (warming) signals are projected to emerge from background variability before 2040 in canyons of the Antarctic, northwest Atlantic, and South Pacific (FAO, 2019). Enhanced stratification and change in the intensity and frequency of downwelling processes under atmospheric forcing (including storms and density-driven cascading events would alter organic matter transported through canyons (Allen and Durrieu de Madron, 2009) ( ''low confidence'' ). Changes in the quantity and quality of transferred particulate organic matter, as well as physical disturbance during extreme events cause a complex combination of positive and negative impacts at different depths along the canyon floor (Canals et al., 2006; Pusceddu et al., 2010). Canyons and slopes are recognised as hosting many methane seeps and other chemosynthetic habitats (e.g., whale and wood falls) supported by massive transport of terrestrial organic matter (Pruski et al., 2017); their climate vulnerabilities are discussed below.

Seamounts have been proposed to serve as refugia for cold water corals facing shoaling aragonite saturation horizons (Tittensor et al., 2011), but could become too warm for deep-water corals in some regions (e.g., projections off Australia) (Thresher et al., 2015) ( ''one study, low confidence'' ). Seamounts are major spawning grounds for fishes; reproduction on seamounts may be disrupted by warming (Henry et al., 2016) ( ''one study, low confidence'' ). In the north Atlantic, models suggest seamounts are an important source of cold water coral larvae that maintain resilience under shifting NAO conditions (Fox et al., 2016), thus loss of suitable seamount habitat may have far-reaching consequences (Gehlen et al., 2014) ( ''low confidence'' ) (also see Box 5.2).

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'''Table 5.5'''

Projected climate changes from the present to 2081–2100 given as mean (min, max) at the deep seafloor for continental slopes, canyons, seamounts and cold water corals mapped from 200–2500 m under RCP8.5 and RCP2.6 Projections are based on three 3D, fully coupled earth system models (ESMs) (as part of CMIP5): the Geophysical Fluid Dynamics Laboratory’s ESM 2G (GFDL-ESM-2G); the Institut Pierre Simon Laplace’s CM6-MR (IPSL-CM5A-MR); and (iii) the Max Planck Institute’s ESM-MR (MPI-ESM-MR). Export flux at 100 m was converted to export POC flux at the seafloor (epc) using the Martin curve following the equation: epc = epc100 (depth/ export depth)-0.858. Projections were made onto the (i) slope from a global ocean basin mask from World Ocean Atlas 2013 V2 (NOAA, 2013), (ii) global distribution of submarine canyons with canyon heads shallower than 1500 m (Harris and Whiteway, 2011); (iii) global distribution of seamounts with summits between 200–2500 m (Kim et al. 2011); and (iv) global occurrence of cold water corals between 200–2500 m (Freiwald et al. 2017).

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{| class="wikitable"
|-
| ''' '''
| '''  Temperature '''
'''        (''' '''o''' '''C)'''

| '''      pH '''
| '''DO '''
'''(''' ''µ'' '''Mol kg''' '''-1''' ''')'''

| '''POC flux'''
'''(mgC m''' '''-2''' '''d''' '''-1''' ''')'''

|-
| ''' '''
| '''RCP2.6'''
| ''' RCP2.6'''
|-
| Continental slopes
| +0.30 (–0.44, + 2.30)
| –0.06 (–0.19, –0.02)
| –3.1 (–49.3, +61.7)
| –0.39 (–16.0, +3.9)
|-
| Canyons
| +0.31 (–0.27, +1.76)
| -0.05 (-0.13, +0.01)
| –3.5 (–44.7, +29.3)
| –0.33 (–10.53, +3.53)
|-
| Seamounts
| +0.13 (+0.01, +0.67)
| -0.02 (-0.11, +0.005)
| –3.46 (–18.9, +4.1)
| –0.15 (–2.20, +1.33)
|-
| Cold water corals
| +4.3 (–0.29, +1.85)
| -0.07 (-0.13, 0.0)
| –3.5 (–25.6, +24.7)
| –0.7 (–10.5, +3.4)
|-
|
|-
|
| '''RCP8.5'''
|-
| Continental slopes
| +0.75 (–8.4, +4.4)
| –0.14 (–0.44, –0.02)
| –10.2 (–67.8, +53.8)
|  –0.66 (–33.33, +10.3)
|-
| Canyons
| +0.19 (–0.03, +1.14)
| -0.11 (-0.35, +0.02)
| –0.8 (–28.8, +10.1)
| –0.80 (–28.76, +10.07)
|-
| Seamounts
| +0.66 (–0.75, +3.19)
| -0.03 (-0.19, +0.001)
| –0.50 (–7.2, +3.0)
| –0.50 (–7.18, +2.98)
|-
| Cold water corals
| +0.96 (–0.42, +3.84)
| -0.15 (-0.39, +0.001)
| –10.6 (–59.2, +11.1)
| –1.69 (–20.1, +4.6)
|}
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==== 5.2.4.4 Chemosynthetic Ecosystems ====

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Despite having nutrition derived largely from chemosynthetic sources fueled by fluids from the earth’s interior, hydrothermal vent and methane seep ecosystems are linked to surface ocean environments and water-column processes in many ways that can expose them to aspects of climate change ( ''medium confidence'' ). The reliance of vent and seep mussels on surface-derived photosynthetic production to supplement chemosynthetic food sources (Riou et al., 2010 <sup>[[#fn:r778|778]]</sup> ; Riekenberg et al., 2016 <sup>[[#fn:r779|779]]</sup> ; Demopoulos et al., 2019 <sup>[[#fn:r780|780]]</sup> ), and in some cases as a cue for synchronised gametogenesis (sperm and egg production) (Dixon et al., 2006 <sup>[[#fn:r781|781]]</sup> ; Tyler et al., 2007 <sup>[[#fn:r782|782]]</sup> ) can make them vulnerable to changing amounts or timing of POC flux to the deep seabed in most areas except high latitudes, or to changes in timing of surface production (see Section 5.2.2.5) ( ''limited evidence'' ) Most of the large, habitat-forming (foundation) species at vents and seeps such as mussels, tubeworms, and clams require oxygen to serve as electron acceptor for aerobic hydrogen-, sulfide- and methane oxidation (Dubilier et al., 2008 <sup>[[#fn:r783|783]]</sup> ) and appear unable to grow under dysoxic conditions (&lt;5–10 µmol kg –1 O 2 ) (Sweetman et al., 2017 <sup>[[#fn:r784|784]]</sup> ) ( ''medium confidence'' ). The distributions of these taxa at seeps could be constrained by climate-driven expansion of midwater oxygen minima (Stramma et al., 2008 <sup>[[#fn:r785|785]]</sup> ; Schmidtko et al., 2017 <sup>[[#fn:r786|786]]</sup> ), which is occurring at water depths where seep ecosystems typically occur on continental margins (200–1000 m). Rising bottom temperatures or shifting of warm currents on continental margins could increase dissociation of buried gas hydrates on margins (Phrampus and Hornbach, 2012 <sup>[[#fn:r787|787]]</sup> ) ( ''low confidence'' ) potentially intensifying anaerobic methane oxidation (which produces hydrogen sulfide) (Boetius and Wenzhoefer, 2013 <sup>[[#fn:r788|788]]</sup> ) and expanding cover of methane seep communities ( ''limited'' evidence). Larvae of vent species such as bathymodiolin mussels, alvinocarid shrimp, and some limpets that develop in or near surface waters (Herring and Dixon, 1998 <sup>[[#fn:r789|789]]</sup> ; Arellano et al., 2014 <sup>[[#fn:r790|790]]</sup> ), are likely to be exposed to warming waters, decreasing pH and carbonate saturation states, and in some places, reduced phytoplankton availability (Section 5.2.2), causing reduced calcification and growth rates (as in shallow water mussel larvae, Frieder et al. (2014)) ( ''limited evidence, low confidence'' ). Larvae originating at vents or seeps beneath upwelling regions may also be impaired by effects of hypoxia associated with expanding OMZ (Stramma et al., 2008 <sup>[[#fn:r791|791]]</sup> ) during migration to the surface ( ''limited evidence)'' . Warming and its effects on climate cycles have the potential to alter patterns of larval transport and population connectivity through changes in circulation (Fox et al., 2016 <sup>[[#fn:r792|792]]</sup> ) or surface generated mesoscale eddies (Adams et al., 2011 <sup>[[#fn:r793|793]]</sup> ) ( ''limited evidence'' ; ''low confidence'' ). Climate-induced changes in the distribution and cover of vent and seep foundation species may involve alteration of attachment substrate, food and refuge for the many habitat-endemic species that rely on them (Cordes et al., 2010 <sup>[[#fn:r794|794]]</sup> ) and for the surrounding deep sea ecosystems which interact through transport of nutrients and microbes, movement of vagrant predators and scavengers, and plankton interactions (Levin et al., 2016 <sup>[[#fn:r795|795]]</sup> ) ( ''limited evidence'' ; ''low confidence'' ). There is, however, insufficient analysis of faunal symbiont and nutritional requirements, life histories, larval transport and cross-system interaction to quantify the extent of the consequences described above under future climate conditions.

<div id="section-5-2-4-4chemosynthetic-ecosystems-block-2" class="box"></div>

<span id="box-5.2-cold-water-corals-and-sponges"></span>