The oceans are a major sink for CO2 produced by the burning of fossil fuels (Pauchauri et al., 2014) as well as for the heat produced by the greenhouse effect (Glecker et al., 2016). Oceans thus help to buffer multiple aspects of global climate change and their effects on marine and terrestrial ecosystems (Reid et al., 2009). Deep-sea ecological processes and characteristics, such as nutrient cycling, carbon sequestration, productivity, habitat provision, and trophic support, underlie the healthy functioning of ocean ecosystems and provide valuable ecosystem services to mankind (Thurber et al., 2014). For example, nutrients produced during the re-mineralization of organic matter at the deep seafloor are ultimately used by phytoplankton to produce organic matter that fuels secondary production. At the same time, organic-matter degradation and re-mineralisation contribute to carbon biogeochemical cycling in the ocean, and help to buffer the ocean against pH changes and the effects of ocean acidification (Berelson et al., 1997; Wenzhöfer et al., 2001; Cerrano et al., 2013). The health and sustainable functioning of the planet are therefore highly dependent on the deep sea (defined here as > 200 m), which accounts for more than 95% of the volume of the Earth’s oceans.
Atmospheric CO2 concentrations have risen from ~ 280 ppm during pre-industrial times to 407 ppm today as a result of the burning of fossil fuels, deforestation and the removal of other habitats that sequester carbon. Continued use of fossil fuels into the 21st century is predicted to lead to atmospheric CO2 levels > 900 ppm by 2100 (under Representative Concentration Pathway (RCP) 8.5; Meinshausen et al., 2011), though the precise level is highly dependent on the emission scenario (Pachauri et al., 2014). These rising atmospheric greenhouse gas concentrations have led to an increase in global average temperatures of ~ 0.2°C decade–1, much of which has been absorbed by the oceans, whilst the oceanic uptake of atmospheric CO2 has led to major changes in surface ocean pH (Levitus et al., 2000, 2005; Feely et al., 2008; Hoegh-Guldberg and Bruno, 2010; Mora et al., 2013; Roemmich et al., 2015).
The deep sea has experienced dramatic changes in physical and chemical variables in the geological past. For example, major expansion and strengthening of oxygen minimum zones (OMZs; O2 < 0.5 mL L–1) are known to have occurred during abrupt, decadal–centennial-scale warming events during the last deglaciation (Moffitt et al., 2015b; Praetorius et al., 2015). Similar fluctuations in OMZ intensity have occurred during the Dansgaard-Oeschger (D-O) events (millennial-scale abrupt climate oscillations) during the last glacial period (Cannariato and Kennett, 1999; Schmittner et al., 2007). Bottom-water temperature and current velocities have also fluctuated in relation to decadal–millennial scale climatic changes during the last de-glaciation and Holocene (Bianchi and McCave, 1999; Marchitto and deMenocal, 2003; Farmer et al., 2011; Cronin et al., 2012). Finally, deep-sea bottom temperatures have also exhibited systematic glacial-low, interglacial-high patterns during the Plio-Pleistocene (Dwyer et al., 1995; Martin et al., 2002; Sosdian and Rosenthal, 2009; Elderfield et al., 2010).
These paleo-environmental changes significantly modified deep-sea biodiversity and probably ecosystem functioning (Thomas and Gooday, 1996; Yasuhara and Cronin, 2008; Yasuhara et al., 2008, 2016). For example, in the bathyal Santa Barbara Basin (California margin), rapid, alternated shifts in the seafloor ecosystem have occurred in response to changes in OMZ intensity associated with D-O and de-glacial abrupt warming events (Cannariato et al., 1999; Moffitt et al., 2015a). North Atlantic fossil records during the last de-glaciation showed abrupt changes in deep-sea biodiversity associated with a rapidly changing climate (specifically deep-water circulation and temperature) over decadal to centennial time-scales (Yasuhara et al., 2008, 2014). Longer time-scale paleo-ecological studies have also shown systematic changes in faunal structure and biodiversity related to glacial-interglacial climate cycles over the last three million years, and more specifically to climate-driven changes in bottom temperature or particulate organic carbon (POC) flux, depending on the ocean and taxonomic group (Cronin et al., 1996; Cronin and Raymo, 1997; Yasuhara and Cronin, 2008; Yasuhara et al., 2009, 2012a; Yasuhara and Danovaro, 2016). Paleoecology (Yasuhara et al., 2015) has thus revealed the dynamic and sensitive nature of deep-sea ecosystem structure and biodiversity across a wide range of time scales in response to changing climatic conditions.
Many observational studies are showing that present-day climate change is already impacting deep-sea environments, as evidenced by increased deep-sea temperature (Purkey and Johnson, 2010), deoxygenation (Stramma et al., 2008, 2010, 2012; Keeling et al., 2010; Helm et al., 2011), lowered pH of intermediate deep-waters (Byrne et al., 2010), and altered POC flux to the seafloor (Ruhl and Smith, 2004; Smith et al., 2013). Despite emerging evidence that climate-driven changes in deep-sea environmental conditions may perturb the functioning of ocean-floor ecosystems (Danovaro et al., 2001; Smith et al., 2007, 2008; Dunlop et al., 2016; Yasuhara et al., 2016), our understanding of the extent to which projected physical and chemical changes will lead to deleterious ecological consequences is still very poor (Philippart et al., 2011). Given that deep-sea ecosystems are vitally important to the Earth system (Danovaro et al., 2014) and are at considerable risk from ongoing climate change (Mora et al., 2013; Jones et al., 2014; Levin and Le Bris, 2015), our goal in this paper is to understand and predict the nature and consequences of climate change at the deep seafloor until 2100. To this end, we describe the present status of four major environmental variables at the seabed that are likely to be altered by increased CO2 emissions to the atmosphere: temperature, oxygenation, pH and food supply (or POC flux). We then explore how these conditions may change by 2100, and assess how these changes are likely to modify benthic biodiversity and ecosystem functioning especially along continental margins, in polar regions, and at the abyssal seafloor. Finally, we briefly address how additional pressures (e.g., from bottom fishing, deep-sea mining) may further impact deep-seafloor ecosystems, and identify the possible societal implications of these combined changes.
To identify the present and future state of deep-sea ecosystems, we used a combination of expert opinion, current literature, and the output of the IPCC (Intergovernmental Panel on Climate Change) Fifth Assessment Report (AR5) models. To generate maps of present-day ocean conditions at the seafloor we: (1) derived the annual climatological means of bottom temperature and dissolved O2 based on the World Ocean Atlas 2013 (available from the NOAA National Oceanographic Data Center at http://www.nodc.noaa.gov/OC5/WOA09/pr_woa09.html); (2) calculated the current seafloor pH from total inorganic CO2, total alkalinity, temperature, salinity, and pressure using the program CO2SYS (data on seafloor carbon content and the program CO2SYS were available from the Global Ocean Data Analysis Project at http://cdiac.ornl.gov/oceans/; data on temperature, salinity and pressure were from the annual climatological mean from the World Ocean Atlas 2013); and (3) estimated POC flux to the global seafloor by first gathering data (from the portal www.science.oregonstate.edu/ocean.productivity) for global climatological monthly mean SeaWiFS (1998–2007) and MODIS (from 2008–2010) Level-3 chlorophyll-a concentration and Level-4 VGPM ocean primary productivity (Behrenfield and Falkowski, 1997). The export POC flux at the seafloor was then calculated using an equation from Lutz et al. (2007) based on the mean and seasonality (standard deviation/mean) of primary production, as well as the mean export depth below the euphotic zone over a 12-year period. The euphotic zone was calculated from the mean surface chlorophyll concentrations using the Case I model of Morel and Berthon (1989), while the export depth was calculated by subtracting the euphotic zone depth from the water depth. The ETOPO1 Global Relief Model from the NOAA National Geophysical Data Center (www.ngdc.noaa.gov/mgg/global/) was used for the global ocean depth.
Future ocean projections for the year 2100 were compiled from all available data generated by Earth Systems Models as part of the Coupled Model Inter-comparison Project Phase 5 (CMIP5) to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change (Taylor et al., 2012) as in Mora et al. (2013). A total of 31 Earth System Models from 18 centers in nine countries were collected and multi-model averages of temperature, pH, O2 and POC flux to the seafloor were reported as the inter-annual mean projections between 2090 and 2100 (Mora et al., 2013). Detailed descriptions of the accuracy and precision of multi-model estimates can be found in Mora et al. (2013).
Abyssal (water depth 3000–6000 m) and polar habitats can be characterized as cold, slightly alkaline, and well-oxygenated systems (Table 1; Figure 1). They are also quite dynamic environments with environmental conditions (e.g., temperature, POC flux) fluctuating over intra- and interannual timescales. At bathyal depths (200–3000 m) on continental margins, spatial gradients in temperature, pH, O2 and food supply can be much steeper (Table 1; Figure 1). One of the major differences between abyssal and bathyal regions is in terms of food supply to the seafloor, with abyssal regions being characterized by extreme food limitation over many thousands of kilometers (Table 1; Figure 1).
|Ocean basin||Depth zonea||Temperature (°C)||Dissolved oxygen (mL L–1)||pH||Seafloor POC flux (mg C m–2 d–1)|
|Atlantic||> 3 km||–0.07 to 4.39||4.6 to 6.92||7.98 to 8.11||0.69 to 10.41|
|0.2–3 km||–1.23 to 27.83||1.48 to 7.54||7.71 to 8.25||1.4 to 108.05|
|Pacific||> 3 km||0.7 to 3.84||2.01 to 4.9||7.72 to 8||0.56 to 10.68|
|0.2–3 km||–0.23 to 28.39||0.24 to 7.31||7.58 to 8.2||0.66 to 81.86|
|Indian||> 3 km||0.01 to 2.47||2.48 to 5.4||7.9 to 8.03||0.91 to 6.71|
|0.2–3 km||0.28 to 25.22||0.1 to 7.16||7.7 to 8.12||1.4 to 61.91|
|Southern||> 3 km||–0.62 to 1.62||4.15 to 6.07||7.94 to 8.03||0.16 to 2.67|
|0.2–3 km||–2.03 to 8.52||3.66 to 7.81||7.9 to 8.16||0.21 to 51.67|
|Arctic||> 3 km||–1.05 to –0.25||5.79 to 7.18||NAb||0.26 to 4.34|
|0.2–3 km||–1.77 to 7.87||2.67 to 8.13||NA||0.3 to 75.57|
Currently, temperatures at the abyssal seafloor at low to mid latitudes in the Atlantic, Indian and Pacific Oceans range between about 0.01 and 4°C (Table 1; Figure 1) (Yasuhara and Danovaro, 2016). Bathyal temperatures typically range from 2 to 8°C (Figure 1), though exceptionally warm seafloor temperatures can be found at upper bathyal depths, and in smaller enclosed seas such as the deep Red Sea (> 20°C; Roder et al., 2013), and bathyal and abyssal Mediterranean Sea (12–14°C). Sub-zero seafloor temperatures tend to occur only at high latitudes (Table 1; Figure 1).
Earth-system-model analyses suggest that some abyssal ecosystems are presently warming at rates of 0.01 to 0.1°C decade–1 (e.g., the Southern Ocean; Purkey and Johnson, 2010). Approximately 19% of the ocean’s heat uptake has gone into the deep ocean > 2000 m (Talley et al., 2015). Over the next 84 years, the highest temperature changes are likely to occur at the abyssal seafloor in the North Atlantic, Southern and Arctic Oceans (Mora et al., 2013) (Table 2; Figure 2). Bathyal depths are also likely to experience increasing temperatures (Soltwedel et al., 2005); Mora et al. (2013) modeled increases of 3.6, 4.4 and 3.7°C in the Pacific, Atlantic, and Arctic Oceans (e.g., Barents and Kara Sea), with lower temperature increases in the Indian and Southern Oceans (e.g., at sites of deep-water formation in the Weddell Sea) (Table 2; Figure 2) by 2100. These predictions of temperature change are in alignment with evidence that the deep Greenland Sea has warmed by 0.7°C since the 1950s (~ 0.01°C yr–1; Somavilla et al., 2013). Bathyal waters off Antarctica are also warming by 0.005–0.01°C yr–1 (Smith et al., 2007; Purkey and Johnson, 2010).
|Ocean basin||Depth zonea||Temperature (°C)||Dissolved oxygen (mL L–1)||pH||Seafloor POC flux (mg C m–2 d–1)|
|Atlantic||> 3 km||–0.37 to 0.98||–0.03 to 0||–0.13 to 0||–1.27 to –0.02|
|0.2–3 km||–0.32 to 4.41||–0.03 to 0.02||–0.37 to –0.01||–13.73 to 0.63|
|Pacific||> 3 km||0.02 to 0.47||–0.01 to 0||–0.06 to 0.01||–1.27 to 0.51|
|0.2–3 km||0.03 to 3.63||–0.05 to 0.01||–0.31 to 0||–5.19 to 2.37|
|Indian||> 3 km||–0.01 to 0.45||–0.01 to 0||–0.06 to 0||–0.64 to –0.04|
|0.2–3 km||–0.01 to 2.17||–0.01 to 0||–0.29 to 0||–3.7 to 0.4|
|Southern||> 3 km||0.09 to 0.74||–0.03 to 0||–0.09 to –0.01||–0.31 to 0.48|
|0.2–3 km||0.08 to 1.71||–0.04 to 0||–0.33 to –0.01||–1.95 to 4.78|
|Arctic||> 3 km||–0.13 to 0.76||–0.02 to 0||–0.14 to –0.01||–1.25 to 1.64|
|0.2–3 km||–0.18 to 3.71||–0.03 to 0.02||–0.37 to –0.01||–9.71 to 3.57|
Presently, much of the Atlantic Ocean is well oxygenated (Figure 1) relative to the North Indian and Pacific Oceans, where bottom water O2 concentrations are lower because of the biological removal of O2 as thermohaline circulation moves deep waters across ocean basins from the North and South Atlantic towards the North Pacific, in isolation from the surface ocean. Warming of the oceans will enhance thermal stratification and density gradients, which will reduce vertical mixing. Combined with a reduction in O2 solubility in warmer water, increased thermal stratification is predicted to create widespread ocean de-oxygenation (Keeling et al., 2010; Long et al., 2016), with the greatest effect in intermediate waters (100–1000 m; Stramma et al., 2012; Bopp et al., 2013). Already, distinct deep-water masses in the Southern Ocean (Helm et al., 2011), eastern North Atlantic (e.g., Sub-polar Mode Water, the Intermediate Water and the Mediterranean Outflow Water; Stendardo et al., 2015), and in the West Pacific (North Pacific Subtropical mode water; Takatani et al., 2012) display long-term decreases in O2 concentration. Bathyal seafloor habitats in the North Pacific, North Atlantic, Arctic and Southern Oceans could experience a reduction in bottom-water oxygenation by 0.03–0.05 mL L–1 by the year 2100 (Table 2; Figure 2), which represents a reduction in water column O2 levels by 0.7%–3.7% (Table 3; Figure 3). Significant OMZs are presently found along the continental margins of the East Pacific, Southeast Atlantic, West Pacific and North Indian Oceans (Levin, 2003; Helly and Levin, 2004). Ecosystems within and on the fringes of OMZs could be particularly affected by the O2 and warming changes predicted for bathyal environments (Table 3; Figures 2, 3; Keeling et al., 2010; Sperling et al., 2016). Though not resolved at the grid resolution shown in Figure 1, tropical and subtropical bathyal waters between 200 and 700 m, including those bathyal margins of all major eastern boundaries, have already lost considerable amounts of O2 over the last half-century, and many OMZs have expanded in volume (Stramma et al., 2008, 2010, 2012; Gilly et al., 2013; Deutsch et al., 2014).
|Ocean basin||Depth zonea||Dissolved oxygen||Seafloor POC flux|
|Atlantic||> 3 km||–0.44 to –0.02||–27.12 to –1.26|
|0.2–3 km||–0.68 to 2.05||–36.27 to 4.79|
|Pacific||> 3 km||–0.37 to 0.03||–31.8 to 9.8|
|0.2–3 km||–3.71 to 0.25||–50.4 to 6.54|
|Indian||> 3 km||–0.22 to –0.02||–39.65 to –1.27|
|0.2–3 km||–2.36 to 0.72||–54.61 to 3.87|
|Southern||> 3 km||–0.46 to –0.06||–18.81 to 21.2|
|0.2–3 km||–0.84 to –0.03||–14.63 to 53.16|
|Arctic||> 3 km||–0.31 to –0.03||–15.21 to 37.76|
|0.2–3 km||–1.08 to 0.24||–34.58 to 59.73|
In the abyssal realm, seafloor habitats under areas of deep-water formation (e.g., those in the North Atlantic and Southern Oceans) could experience a maximum decline in O2 concentration of 0.03 mL L–1 by 2100 (i.e., a 0.5% drop from current levels; Tables 2, 3; Figures 2, 3). Greenhouse warming may also exert an effect on abyssal O2 levels (as well as pH and temperature) by changing thermohaline circulation (Rahmstorf et al., 2015; Yamamoto et al., 2015). Reduced Atlantic Ocean overturning circulation will initially lead to lower O2 levels at the deep seafloor, and may alter the intensity of Pacific and Indian Ocean OMZs (Schmittner et al., 2007). However, over longer terms, deep-water oxygenation may also increase even if Atlantic meridional overturning circulation becomes weaker, as deep convection in the Weddell Sea and Antarctic Bottom Water becomes enhanced (Yamamoto et al., 2015).
The North Atlantic Ocean is currently the most alkaline in terms of seafloor pH, while the pH of the deep North Pacific Ocean is lower (Figure 1). This spatial gradient in pH reflects the age and isolation of the water masses, which accumulate CO2 released by biological respiration as they move through the ocean basins. There is also a contribution of excess atmospheric CO2 absorption introduced to deep-water masses from dense, cold CO2-rich surface waters at downwelling sites (e.g., North Atlantic), which then move through the oceans via meridional overturning circulation. Presently, studies in the Pacific Ocean have revealed that intermediate waters down to 500 m depth have experienced a decline in pH of 0.06 units between 1991 and 2006, with the greatest changes occurring around 25°N (Byrne et al., 2010). Model predictions from the North Atlantic have revealed that over 17% of the seafloor area below 500 m depth will experience pH reductions exceeding 0.2 units by 2100 because of subduction of high-CO2 waters by thermohaline circulation (Gehlen et al., 2014). These major pH reductions are projected to occur over important deep-sea features, such as seamounts and canyons (Gehlen et al., 2014). Bathyal seafloor habitats in other areas of the world’s oceans will also experience significant reductions in pH by the year 2100 (e.g., a decrease of 0.29 to 0.37 pH units) as a result of the entrainment of CO2-rich seawater to the seafloor at sites of bottom-water formation (Table 2; Figure 2). Since it is estimated to take 1000 years for water entrained at deep-water formation sites to circulate throughout the abyssal ocean basins, vast areas of the abyssal Indian and North Pacific Oceans will experience a lesser shift in pH by 2100 (maximum decrease in pH = 0.06; Table 2; Figure 2).
With the exception of certain continental margin habitats (e.g., OMZs, some canyons and seamounts) and chemosynthetic ecosystems, most of the deep sea, and particularly the abyss, is characterized by severe food limitation (POC flux of 1–2 g C m–2 yr–1; Watling et al., 2013) (Table 1; Figure 1). Currently, regions with the highest POC export lie at high latitudes, although transfer efficiency (the proportion of POC exported that arrives at the seafloor) is lowest here compared to lower latitudes. At lower latitudes, extensive mineralization takes place in the upper water column leading to less export from the euphotic zone, but transfer efficiency is higher, as most of the exported carbon tends to be refractory (Henson et al., 2012). Enhanced warming of the upper ocean is predicted to enhance stratification, reducing nutrient input to the upper euphotic zone and causing a shift in phytoplankton assemblages from large, fast-sinking diatoms (with low surface area:volume [SA:V] ratios) to slow-sinking picoplankton (with high SA:V ratios; Bopp et al., 2005). This shift is likely to reduce export flux to the seafloor, as well as transfer efficiency (Buesseler et al., 2007; Morán et al., 2010, 2015; Steinacher et al., 2010). Furthermore, freshening of Arctic regions by sea-ice meltwater and episodic input of large river runoff have been shown to reduce phytoplankton size and, by inference, export flux, a trend that has been projected to continue into the future (Li et al., 2009, 2013). The areas likely to be impacted by significant declines in POC flux owing to enhanced water column stratification lie in the North and South Pacific, North and South Atlantic, and North and South Indian Oceans (Tables 2, 3; Figures 2, 3). The abyssal and bathyal regions of the Indian Ocean are predicted to experience declines in POC flux by as much as 40% and 55%, respectively (Tables 2, 3; Figures 2, 3) by 2100. For abyssal seafloor habitats, these changes could be severe, as they are some of the most food-limited regions on the planet (Berelson et al., 1997; Danovaro et al., 2001; Smith et al., 1997, 2008). In contrast, the polar oceans are areas where POC flux is likely to rise. Vast abyssal and bathyal areas of the Arctic and Southern Oceans are predicted to experience a POC flux increase of upto 60 and 53%, respectively (Tables 2, 3; Figures 2, 3) as a result of longer ice-free periods (Comiso, 2010), though accounting for future shifts in phytoplankton size and their nutrient supply could modify this expectation (Tremblay et al., 2012; Arrigo, 2013). Localized increases in POC are also predicted for upwelling regions, such as coastal Chile and the west coast of the United States (Figures 2, 3) (Jones et al., 2014). These regions have already seen an increase in primary production and frequency of high POC flux to the deep seafloor because of increased wind stress and nutrient upwelling.
The continental margins of the world’s ocean are among the most heterogeneous and diverse of the ocean’s environments (Levin and Sibuet, 2012). Hydrographic, topographic and biotic influences create a multitude of seabed habitats and biomass/biodiversity hotspots, and can influence biodiversity–ecosystem functioning relationships (Zeppilli et al., 2016). These include broad expanses of organic-rich sediment, low O2 zones and OMZs, seamounts, banks, ridges, fjords, canyons, basins, coral and sponge reefs, organic falls, and areas of methane seepage. This heterogeneity supports the biodiversity responsible for a whole host of ecosystem functions and services (Levin and Dayton, 2009; Thurber et al., 2014). Because of the shallower depth of continental margin habitats and closer connections with land compared to abyssal habitats, continental margin ecosystems are likely to experience a greater degree of change in all environmental parameters compared to the abyssal seafloor (Tables 2, 3).
While few continental margin systems have been investigated directly in terms of the consequences of climate change, their strong gradients and regional variations have allowed significant understanding of the effects of temperature, pH, O2 and POC flux on deep-sea benthic ecosystems. Warming of surface waters along continental margins, and increased thermal stratification and reduced nutrient supply to the surface are likely to reduce both productivity and phytoplankton type and size (Smith et al., 2008; Morán et al., 2010, 2015), yielding reduced phytodetrital flux to the seabed (Bopp et al., 2005; Buesseler et al., 2008; Smith et al., 2008). Model outputs suggest that bathyal areas particularly prone to declining POC flux lie in the Norwegian and Caribbean Seas, NW and NE Atlantic, the eastern tropical Pacific, and bathyal Indian and Southern Oceans, which could experience as much as a 55% decline in POC flux by 2100 (Tables 2, 3; Figures 2, 3). Elevated seafloor temperatures at northerly latitudes (Figure 2) will lead to warming boundary currents and may trigger massive release of methane from gas hydrates buried on margins (Phrampus and Hornbach, 2012; Johnson et al., 2015) especially in the Arctic, with simultaneous effects on global climate, aerobic methane oxidation, water column de-oxygenation and ocean acidification (Biastoch et al., 2011; Boetius and Wenzhöfer, 2013). Along canyon-cut margins (e.g., the western Mediterranean), warming may additionally reduce density-driven cascading events, leading to decreased organic matter transport to the seafloor (Canals et al., 2006), though this very process is also likely to reduce physical disturbance at the seafloor. Greenhouse warming will also increase temperature differentials between land and oceans, and intensify wind-driven upwelling in eastern boundary currents, stimulating photosynthetic production at the surface (Bakun, 1990; Bakun et al., 2015; Wang et al., 2015). However, this new production will ultimately start to decay as it sinks and increase biogeochemical drawdown of O2. Upwelling will also bring low-O2, high-CO2 water onto the shelf and upper slope (Bakun, 1990; Feely et al., 2008; Bakun et al., 2010; Sydeman et al., 2014; Wang et al., 2015). Increased levels of precipitation on land will also alter terrestrial inputs, including sediments and organic debris, nutrients, and contaminants (Jaedicke et al., 2009; Caroletti and Barstad, 2010) that may smother seafloor sediments, and alter the trophic ecology of deep-sea habitats situated close to land (McLeod and Wing, 2007, McLeod et al., 2010).
Margin habitats are noted for dense, high biomass aggregations of structure-forming species, such as cold-water coral (CWC) reefs and coral ‘garden’ habitats (Coleman and Williams, 2002; Roberts et al., 2006; Levin and Dayton, 2009; Buhl-Mortensen et al., 2010; De Leo et al., 2010). CWC structures provide shelter from predation for a variety of fauna and, as such, can act as nursery grounds for commercially important species (Koslow et al., 2000, Baillon et al., 2012; Henry et al., 2013). The habitat complexity of these biogenic reefs also leads to high levels of biodiversity on the reefs (Henry et al., 2007). By altering internal currents, CWCs can also act as ecosystem engineers boosting organic matter deposition at the seafloor (van Oevelen et al., 2009; Soetaert et al., 2016). CWCs and other calcifying taxa (e.g., bivalves and echinoderms) may be susceptible to ocean acidification leading to brittle structures, enhanced susceptibility to predation and a loss of habitat as a result of lowered aragonite and calcite saturation states (Figures 2, 4C).
To date, studies of aragonitic, scleractinian CWC responses to ocean acidification have frequently examined short-term acclimation, with effects on coral bio-mineralization, growth, and skeletal strength only becoming evident in experiments run for periods of a year or more (e.g., Tittensor et al., 2010). Intriguingly, aragonitic CWC species are found close to and even below the aragonite saturation horizon (Roberts et al., 2009a; Findlay et al., 2014), raising the question of whether species adapted to lower saturation states may have inherent adaptations to future lower pH ocean conditions. However, with many of the known CWC reefs projected to be bathed in under-saturated water by the end of the century (Guinotte et al., 2006; Roberts et al., 2006) the accumulated biogenic reef structures will degrade over time, even if living corals persist (Hennige et al., 2015). This degradation could have implications for habitat provision with consequent effects on fish populations and fisheries production. Likely major impact zones include CWC reefs found in the northern Atlantic and Arctic Oceans, the Southern Ocean, and around New Zealand (Guinotte et al., 2006; Yesson et al., 2012) where deep-water pH could decrease by approximately 0.3–0.4 pH units by 2100 relative to current day values. Reduced food supply owing to lower POC fluxes could exacerbate these impacts because the metabolic cost of increased rates of calcification become greater as pH declines (Wood et al., 2008). Calcareous reef habitats in the northern Atlantic could therefore be especially hard-hit (Figure 2).
The expansion of low O2 zones will affect many aspects of deep-sea ecosystem structure and function (Gooday et al., 2010). Biodiversity declines as O2 levels decline, which can be manifested in multiple ways. Many species of octocorals (including gorgonians and pennatulaceans) provide habitat for a diverse array of associated invertebrates, but octocorals often decrease in abundance as O2 levels decline (e.g., Etnoyer and Morgan, 2005; Roberts et al., 2009b; Buhl-Mortensen et al., 2010). Reductions in octocoral abundance from de-oxygenation could significantly impact hard substratum availability, habitat complexity and benthic biodiversity (Roberts et al., 2009b; Hennige et al., 2015). Sediment-burrowing fauna will probably be increasingly excluded as water column O2 levels decline leading to a reduction in the mixed layer depth, and altered bioturbation rates and C-sequestration in sediments (Smith et al., 1997; Smith et al., 2000; Levin et al., 2009; Sperling et al., 2016; Figure 4B). This cascade of effects is highly likely at depths of 500–1000 m in parts of the Eastern Pacific, where OMZ expansion is projected to exceed thresholds for biodiversity (Sperling et al., 2016; Figures 2, 3).
Intolerant pelagic, demersal and benthic fish and invertebrate species that are mobile will experience habitat compression into shallower depths (McClatchie et al., 2010; Koslow et al., 2011; Sato et al., 2016), or adapt by migrating horizontally along continental slopes into higher-O2 environments. Hypoxia expansion over methane seeps may inhibit oxidizing symbionts that support dense mussel and tubeworm aggregations. These chemosynthetic aggregations typically enhance production and biodiversity on margins (Levin and Dayton, 2009; Cordes et al., 2010), and can provide critical nursery habitats (Treude et al., 2011). In contrast, hypoxia-tolerant taxa (e.g., squid and jellyfish) may expand their population sizes and distributions with consequences for food-web structure and pelagic-benthic coupling (Gilly et al., 2013; Lebrato et al., 2013; Sweetman et al., 2014a).
Single stressors like warming will limit tolerance windows for other stressors such as low O2 or low pH (Pörtner and Knust, 2007; Pörtner, 2012). Reductions in food supply and warming together with expansion of low O2 and pH zones, will increase the vulnerability of key habitats (e.g., CWC reefs) to anthropogenic disturbance (e.g., benthic trawling) and retard recovery of these fragile habitats from physical damage. Heavily-fished areas off the northern coast of Norway, which are also home to abundant CWC reefs (Fosså et al., 2002), could be especially sensitive, as they are predicted to experience an increase in temperature of 2–3°C and pH changes of –0.3 to –0.35 units, while also being subject to declining O2 (0.03 mL L–1: Figures 2, 3).
Benthic organisms inhabiting sediments along continental margins are responsible for most nitrogen cycling, while over 50% of carbon burial in the ocean occurs in continental margin sediments. In eastern current boundary systems (e.g., off the coast of Namibia), where O2 is already at sub-oxic levels, these regulating services (e.g., nitrogen removal, carbon sequestration) are currently, and will continue to be highly sensitive to small changes in oxygenation (Deutsch et al., 2011). For example, expansion of low O2 waters could easily shift carbon-processing pathways by favoring chemosynthesis and by increasing the relative importance of bacteria, protozoa (e.g., foraminifera) and metazoan meiofauna in biogeochemical cycling relative to larger taxa (Levin et al., 2003; Diaz and Rosenberg, 2008; Woulds et al., 2009; Sweetman et al., 2016), which would impact energy flow to upper trophic levels (Figure 4B).
Major changes in the upper ocean resulting from global warming are likely to include increased sea-surface temperatures and thermal stratification, and reduced nutrient upwelling over vast areas of the open ocean (Bopp et al., 2001; Gregg et al., 2003; Fischlin et al., 2007; Smith et al., 2008). Ocean acidification is also predicted to reduce microbial production of nitrate from ammonium (Beman et al., 2011), which could have major consequences for oceanic primary production because a significant fraction of the nitrate used by phytoplankton is generated by nitrification at the ocean surface (Yool et al., 2007). Major consequences of such changes over regional scales will probably include (1) reductions in primary production combined with (2) shifts from diatom-dominated (low SA:V ratio) phytoplankton assemblages with high POC-export efficiencies to picoplankton communities (high SA:V ratio) characterized by low export efficiencies (Smith et al., 2008; Morán et al., 2010; Morán et al., 2015). In addition, reductions in calcification from lowered pH in surface waters could reduce phytoplankton sinking rates through loss of ballast (Hofmann and Schellnhuber, 2009), though this effect will depend on the ratio of the fraction of ballasted vs. un-ballasted fractions of the sinking POC. Our model outputs suggest that seafloor POC flux will decline in most oceanic areas with exceptions off Peru, the northern coast of Chile, and the Southern and Arctic Oceans (Tables 2, 3; Figures 2, 3). The continued reduction in the extent of sea ice in the Arctic is expected to lead to increased photosynthetic primary production and POC flux there (Jones et al., 2014), which could benefit fauna whose energetic demands increase as a result of ocean acidification (e.g., calcifying taxa). Reductions in seafloor POC flux will be most drastic, on a percentage basis, in the oceanic gyres and equatorial upwelling zones, with the northern and southern Pacific Ocean and southern Indian Ocean gyres experiencing as much as a 32–40% decline in POC flux (Tables 2, 3; Figures 2, 3). Recent studies have suggested that the NE Atlantic Ocean could also undergo similar reductions in POC flux (Jones et al., 2014). Because the quantity and quality of POC flux is an important ecological forcing factor in the abyss, abyssal ecosystems will be highly sensitive to such changes (Smith et al. 2008, 2009; Tittensor et al., 2011). For example, 3-fold reductions in POC flux (e.g., from 1.5 to 0.5 g C m–2 yr–1), which might occur in the equatorial Pacific (Laws et al., 2000), are predicted to halve benthic microbial and nematode biomass (Figure 4D). These POC flux changes could also lead to a 5-fold decline in macrofaunal biomass, and cause dramatic reductions in the sediment mixed-layer depth, benthic respiration, and bioturbation intensity (Smith et al., 2008; Jones et al., 2014; Figure 4D). Such a decrease in POC flux would also mean a decline in the diversity of nematodes and macrofauna, which are thought to be key functional components of abyssal seafloor ecosystems. This decline in diversity is particularly likely as these groups of organisms tend to rely heavily on detrital matter sinking to the seafloor for their energy requirements (Danovaro et al., 2008; Smith et al., 2008; Jones et al., 2014).
Holothurians are often the prominent abyssal benthic megafauna (Lauerman and Kaufmann, 1998) and play an important role in organic carbon processing and bioturbation. Holothurian population dynamics have been linked to POC flux (Smith et al., 2008) and are considered indicators of climate change impacts on abyssal ecosystems (Glover et al., 2010). Holothurian community dynamics have been examined in detail at long-term time-series sites in relation to shifts in surface phytodetritus input linked to the North Atlantic Oscillation (Porcupine Abyssal Plain, NE Atlantic; Billett et al., 2001) and to the Northern Oscillation Index and Bakun Upwelling Index (Station M, NE Pacific; Ruhl et al., 2014). At both sites, variation in these climate indices were correlated with increased pulses of POC, which resulted in significant increases in abundance of holothurian species, particularly those species able to rapidly use phytodetrital material and successfully reproduce and recruit (Billett et al., 2001; Smith et al., 2006, 2009; Huffard et al., 2016). These studies confirm the predictions of basic macro-ecology and spatial-gradient studies that climate change fluctuations can cause temporal changes in food inputs leading to changes in overall macro- and megafaunal biomass and community structure in terms of size distributions and dominance (Ruhl et al., 2008, 2014). Thus, it is a reasonable expectation that macro- and megafaunal communities will shift in relation to future climatically linked changes in POC flux to the abyssal seafloor (Figure 4D). Episodic pulses provide food supply to sustain benthic communities over periods of deficit (Smith et al., 2013, 2014); the predicted reduction in POC input over large abyssal areas will likely increase these deficits with a significant impact on faunal communities and their role in ecosystem functioning (Figure 4D).
POC flux to the seafloor, and its degree of seasonality, also strongly influence benthic foraminiferal abundance, diversity and assemblage composition (Altenbach et al., 1999; Sun et al., 2006; Corliss et al., 2009; Gooday et al., 2012). There are well-established decreases in the relative abundance of calcareous taxa (as well as calcite saturation) compared to agglutinated taxa, and in labile organic matter flux, with increasing water depth. Thus, one consequence of reduced POC flux may be enhanced abundances of agglutinated relative to calcareous foraminifera (Gooday, 2003; Cornelius and Gooday, 2004). Shoaling of the carbonate compensation depth (CCD) over time as a result of increased CO2 levels would also have important consequences for the diversity of abyssal benthic foraminifera, especially in regions where the depth of the seafloor lies close to the present depth of the CCD (e.g., the eastern Clarion Clipperton Zone, Pacific Ocean). This is an area where foraminifera are overwhelmingly dominated by agglutinated taxa already (Saidova, 1965; Nozawa et al., 2006). A shift from calcareous to agglutinated foraminifera would likely impact deep-sea function (e.g., deep-sea carbon cycling) in addition to altering the biogeographical distribution of fauna (Gooday et al., 2012). Agglutinated foraminifera, particularly forms such as komokiaceans, which are a dominant faunal component in the abyssal deep sea (Tendal and Hessler, 1977), are believed to have a lower metabolic rate and to be less active in carbon processing than calcareous foraminifera (Gooday et al., 2008).
All of these changes predicted for the abyssal zone are likely to fundamentally alter the structure of abyssal ecosystems, as well as the functions that they provide. Changes to microbial and faunal biomass, as well as shifts in biodiversity resulting from changes in POC flux (Figure 4D), and the complex interactions among benthic organisms, have the potential to feed back over long timescales to a range of intertwined functions, such as carbon cycling, which is highly dependent on benthic biomass and diversity (Thurber et al., 2014).
The Arctic Ocean and Antarctic coastal seas have particular characteristics in common, notably seasonality in solar radiation, sea-ice cover and temperature, that can modulate surface primary production. Despite these commonalities, the oceanographic and physiographic settings and the geological histories of the Arctic and Antarctic regions are very different. There are also major differences in their faunal characteristics (Dayton, 1990; Clarke and Johnston, 2003), as well as contrasting macro-ecological patterns (Brandt et al., 2009; Yasuhara et al., 2012b). For example, the Arctic seafloor has many more surface-burrowing species, such as echiurans, polychaetes, echinoderms and crustaceans, relative to the Antarctic (Dayton, 1990), but deep-sea diversity is generally much lower in the Arctic (Culver and Buzas, 2000). This difference is thought to reflect more glacial disturbance and insufficient time for recolonization in the Arctic, as well as fluctuations in surface productivity and reduced circulation during glacial episodes (Culver and Buzas, 2000).
The Arctic and many areas of the Antarctic (e.g., western Antarctic Peninsula or WAP) are predicted to undergo more surface-water warming than other parts of the Earth over the next century, which will affect surface production, sea-ice cover, and hence food availability and quality for deep-sea benthic organisms (see Ingels et al., 2012, for Antarctic coastal to deep-sea ecosystems). Parts of the Antarctic Peninsula, including the WAP, are already experiencing the greatest increase in mean annual atmospheric temperature on Earth (Chapman and Walsh, 2007; Clarke et al., 2007; Solomon et al., 2007; Smale and Barnes, 2008), and temperatures at the seafloor in the Southern Ocean are predicted to rise by as much as 0.7°C at abyssal depths and 1.7°C at bathyal depths by 2100 (Table 2). Field and modeling studies have also revealed rapid atmospheric and surface-water warming in the Arctic Ocean during recent decades (Overland et al., 2004; Spielhagen et al., 2011). Bathyal Arctic waters are following this rapid warming trend (Soltwedel et al., 2005), and temperatures at both bathyal and abyssal depths could increase by as much as 0.1–3.7°C relative to present-day temperatures by 2100 (Table 2; Figure 2). A recent study has shown that deep-sea benthic Archaea can be more sensitive to temperature shifts than their bacterial counterparts; changes in deep-water temperature may thus alter the relative importance of Archaea in benthic ecosystem processes at polar latitudes (Danovaro et al., 2016). Warming at polar latitudes will also open up new habitat for invasive species (Figure 4A). For example, in the Arctic, the commercially important snow crab has extended its range to the north (Alvsvåg et al., 2009; Bluhm et al., 2009), and warming appears to have led to the appearance of highly predacious, temperature-sensitive king crabs (Lithodidae) in waters of the WAP (Smith et al., 2011). These invasive top predators can threaten the diversity of epi- and infaunal communities, as well as physically disturb large areas of soft sediment (Smith et al., 2007). On the other hand, warming-induced extensions of the ranges of temperate–subpolar benthic species into polar oceans (Bluhm et al., 2011) may increase benthic diversity, both in the short-term (Węsławski et al., 2011) and the long-term (Rasmussen et al., 2003; Yasuhara and Danovaro, 2016), although these invasive species are likely to displace less competitive, endemic species unable to cope with increased temperatures (Figure 4A).
Although open ocean polar regions are likely to experience increased primary productivity and POC flux, the production and transport of organic matter to the seafloor will probably initially decline in deep-sea habitats located close to land. As polar and sub-polar regions become warmer, glacial meltwater and erosion of melting tundra (Węsławski et al., 2011) will enhance water column turbidity in coastal zones (Grange and Smith, 2013; Sahade et al., 2015), reducing water column light levels and phytoplankton production, which could affect near-shore deep-sea systems. The increased sedimentation in deep coastal areas, particularly fjords, may also smother or clog the breathing and feeding apparatus of sessile suspension-feeding fauna. Ophiuroids, capitellid polychaetes and other opportunists may be favored by increased sediment inputs. In deep Arctic fjords, high sediment fluxes already create large areas of burial disturbance, which can negatively impact trophic complexity, diversity and productivity of benthic assemblages while also inducing O2 stress (Syvitski, 1989; Włodarska-Kowalczuk et al., 2005; Renaud et al., 2007). In time, continued warming will reduce sediment fluxes into many high-latitude fjords as a result of glacial retreat onto land, potentially increasing benthic productivity and biodiversity (Syvitski et al., 1989).
Changing ice regimes will impact glacial and ice-sheet calving, with ramifications for physical disturbance in the deep sea. Large icebergs can scour the sediment down to 400 m on the Antarctic shelf. This disturbance leads to scale-dependent recolonization of scoured areas and an increased input of dropstones (Smale and Barnes, 2008). These processes will enhance seafloor heterogeneity and create hard substrates for sessile megafauna (Schulz et al., 2010; Meyer et al. 2015, 2016). Dropstones also create diverse microhabitats for meiofauna, allowing for greater trophic and functional diversity around stones (Hasemann et al., 2013; Gooday et al., 2015). In the longer term, iceberg scouring and dropstone deposition will tend to elevate diversity on regional scales through (re)colonization processes, although the immediate effect of scouring will be local elimination of many species (Gutt et al., 1996; Gutt and Piepenburg, 2003). In addition, recent evidence suggests that iceberg production followed by melting might significantly elevate local nutrient levels, driving greater primary production and POC flux to the seafloor in deeper waters (Smith et al., 2007), though this will probably decline after initial increases under continued warming. Thus, the immediate direct impacts on seafloor communities will be relatively short-lived, but the wider effects may be longer lasting. Finally, melting of icebergs and glacial ice may lead to freshening of surface waters leading to enhanced stratification of the upper water column and the release of essential nutrients and trace metals such as iron (Wang et al., 2014). Together with decreased sea-ice cover these factors may act to increase primary production and POC flux. However, increased respiration (owing to increased temperatures) might result in local hypoxia, especially in isolated intra-shelf basins and fjords such as those found along the WAP.
Changes in the quantity and the quality of POC flux to the seafloor will have impacts on ecosystem structure and function (Figure 4D). Present-day reductions in sea-ice and ice-shelf cover (Comiso, 2010) are leading to changes in upper-ocean pelagic dynamics (e.g., increasing surface primary production, and generating shifts from krill to salps; Loeb et al., 1997; Arrigo et al., 2008, 2013; Arrigo and van Dijken, 2011). Under high O2 conditions at shallow depths, metazoans tend to outcompete bacteria in terms of organic matter processing when carbon input to the seafloor increases (van Nugteren et al., 2009; Sweetman et al., 2014b). If the same holds true for the deep seafloor, then elevated POC fluxes caused by reduced sea-ice cover could trigger a switch from dominance of benthic organic matter processing by bacteria to dominance by metazoans with consequences for energy flow to upper-trophic levels (Figure 4D). Increases in the abundance, biomass and diversity of benthic communities, the depth of bioturbation, the prevalence of large, habitat-forming taxa (sponges, benthic cnidarians), and the extension of species ranges downslope into deeper water (De Rijk et al., 2000) are other likely consequences of enhanced POC flux (Figure 4D). However, minor increases in temperature will also increase overall metabolic rates of benthic and pelagic communities. In Antarctic regions, one of the key processes that is thought to govern the deep-water soft sediment communities is captured in the FOODBANCS hypothesis (Smith and DeMaster, 2008), according to which concentrated summer food pulses and slow microbial enzymatic activity caused by the cold temperatures provide a long-lasting food bank that supports benthic metazoan communities (Mincks et al., 2005). With increasing temperatures, bacterial recycling will be enhanced, potentially leading to a non-linear increase in the overall metabolism of the benthic community and increased food limitation in deeper seas (Figure 4A). This effect could explain why some groups like meiofauna may prefer to feed on bacterial food sources in polar areas (Ingels et al., 2010).
Another important stressor, ocean acidification (OA), will be enhanced at high latitudes because of the higher capacity of seawater to absorb CO2 at low temperatures (Orr et al., 2005; Table 2; Figure 2). While some areas will experience increased POC flux resulting from a loss of sea-ice cover that could help alleviate some OA effects (Wood et al., 2008), the additional carbon deposited will increase respiratory CO2 production with impacts on carbonate dissolution. Carbonate dissolution is an issue of particular concern around the Antarctic continent, where it exerts a strong influence on the distributions of benthic foraminifera and other calcareous fauna. Here, the CCD is currently ‘multi-bathic’, ranging from a few hundred meters in intra-shelf basins to several thousand meters over abyssal plains. The depression of the CCD will increase the energy required for organism calcification (Wood et al., 2008), impacting many of the archetypal Antarctic fauna (e.g., molluscs, echinoderms, foraminifera and ostracods). The exact nature of these effects, however, needs detailed examination because, as with all of these factors, there will be synergies that could either exacerbate or ameliorate the stress induced by another perturbation.
Many of the changes outlined above could affect both Arctic and Antarctic seafloor ecosystems, although their intensity will probably vary between the poles and among areas within polar regions that are warming at different rates (e.g., WAP versus the eastern Ross Sea). Moreover, different organism groups (e.g., bacteria versus metazoans), and different life stages of the same species (e.g., larvae versus adults), may respond differently and display different degrees of sensitivity to environmental changes (Ingels et al., 2012).
Climate mitigation by the deep ocean may ultimately compromise many of the ecosystem services we value. The large fraction of the planet covered in deep waters guarantees that most carbon sequestration and significant nitrogen cycling in the ocean occurs here. Presently the ocean absorbs approximately 25% of industrial area CO2 emissions, and 93% of the heat; much of this absorption occurs in deep waters below 200 m (Levin and Le Bris, 2015). Non-market supporting services are provided by deep-sea ecosystems in the form of habitat provision, nursery grounds, trophic support, refugia, and biodiversity functions provided by assemblages on seamounts, coral and sponge reefs, banks, canyons, slopes, fjords and other settings (Armstrong et al., 2012; Mengerink et al., 2014; Thurber et al., 2014; Levin and Le Bris, 2015). The extensive species, genetic, enzymatic, and biogeochemical diversity hosted by the deep ocean also holds the potential for new pharmaceutical and industrial applications, as well as keys to adaptation to environmental change. Because huge expanses of the deep ocean will be exposed to changing environmental conditions as a result of climate change (Mora et al., 2013; this study), the societal impacts of climate change in the deep sea will undoubtedly be widespread, complex and dynamic. Some effects will be direct; for example, we expect alterations in the distributions and health of open-ocean and deep-sea fish populations and commercially exploited stocks. This impact will result from warming-induced changes in metabolism (Deutsch et al., 2015) and body size (Cheung et al., 2013) linked to latitudinal or depth shifts in species distributions, in addition to vertical habitat compression from OMZ expansions (Prince and Goodyear, 2006; Stramma et al., 2010, 2012; Yasuhara and Danovaro, 2016). Less clear are the impacts of acidification stress and multiple stressors on deep-sea fish populations and fisheries production (Rosa and Seibel, 2008). These effects can be indirect, via shifts in food availability or species interactions (e.g., predation and competition). Direct effects from depletion of O2 levels and rising water temperatures over the next century may also impact embryonic survival rates of vulnerable deep-sea oviparous (egg-laying) elasmobranchs that currently deposit their capsules at the seafloor in very narrow oceanographic niches with distinct O2, salinity and temperature conditions (Henry et al., 2016).
At the same time, the deep sea is fast becoming a target area for increased exploitation of key resources and the dumping of pollutants (Mengerink et al., 2014). For example, oil and gas are currently being extracted to depths > 3000 m on continental margins, with increasing risk for leakage and spills (Cordes et al., 2016). Pressure from fishing has also increased since the mid-20th century on seamounts and along continental margins, and there is now strong evidence that many deepwater fish species (e.g., rockfish, Greenland halibut, lings and tusks, orange roughy, sablefish and blue grenadier) have been severely exploited through trawling and longlining, with some species having been fished to commercial extinction (e.g., pelagic armourhead; Koslow et al. 2000). There is also extensive interest in mineral mining at hydrothermal vent systems along mid ocean ridges and back arc basins, bathyal seamounts and polymetallic nodule areas at abyssal depths, as well as for phosphorites on margins (Mengerink et al., 2014; Thurber et al., 2014; Wedding et al., 2015; Amon et al., 2016; Levin et al., 2016; Vanreusel et al., 2016). Deep-sea sediments and organisms are also sinks for a variety of chemical pollutants, including persistent organochlorine pollutants (e.g., PCBs), which may have toxic effects for a variety of fauna at high concentrations (Froescheis et al., 2000; Looser et al., 2000; Ramirez Llodra et al., 2011).
Many of the areas that will be targeted for resource extraction lie in areas that will be most heavily impacted by climate change (e.g., the Arctic for oil, gas and fisheries, the equatorial open ocean for nodule mining, upwelling margins for phosphorites, and the North Atlantic for fishing). Although some animals (e.g., fishes) may be able to adapt to a limited degree to these anthropogenic stressors, for example, by accelerating their growth and reaching maturity at earlier ages (Koslow et al., 2000), the slow population growth rates and long generation times that characterize the evolution of many deep-sea organisms will ultimately limit their adaptation to stress. Additional stresses imposed by low O2, low pH, elevated temperature and reduced food supply are likely to further reduce the resilience of individual species and ecosystems to anthropogenic stressors, and slow rates of recovery. For example, slowed growth of carbonate skeletons in the face of acidification will reduce recovery of biogenic habitats from bottom fishing disturbance, and delayed development under hypoxic conditions and nutritional stress with declining food availability could impact communities recovering from oil pollution and mining impacts, which will further compromise ecosystem structure and function in the deep sea. Thus, a key challenge going forward is to understand the synergies between different stressors associated with climate change and those linked to direct human impacts, such as physical disturbance from bottom trawling or mining plumes, overfishing, oil spills and more.
There is a growing need for multi-sector ocean governance in the deep sea. Efforts to reform the single-sector approach to marine spatial management should incorporate expected climate-induced changes in temperature, oxygenation, pH, and POC flux to the seabed in the development of regional and international management scenarios. By the end of the 21st century, managers will also need to account for background changes in ecosystem functions and services that may not relate to anthropogenic climate change. Naturally occurring interannual and multidecadal shifts in regional ocean regimes such as the Pacific El Niño-Southern Oscillation, the North Atlantic Oscillation, and the Atlantic Multidecadal Oscillation, for example, are bimodal oscillations that cycle between phases of warmer and cooler sea surface temperatures. These oscillations can produce effects that mimic CO2-induced changes (e.g., altered upwelling regimes of deep nutrient-rich waters with effects on POC flux; Merino and Monreal-Gomez, 2009). Observations of climate variables underpin much of the knowledge and modeling described here. However, these observations are decidedly sparse in the deep sea and need to be enhanced, through sustained deep-ocean observing (e.g., via Deep Argo). Meeting this need will require enhanced international co-operation across disciplines as enshrined by the Global Ocean Observing System, and the associated, nascent Deep Ocean Observing Strategy. An integrated approach would allow policy-makers to react to changing distributions and resilience of marine resources, and would underpin adaptive multi-sectoral marine spatial plans to help safeguard deep-sea ecosystems.
The deep sea provides our global society a diversity of ecological and ecosystem services, which are likely to expand in the coming decades. At the same time, a number of co-occurring stressors are likely to impact the ecology of deep-sea communities and the ways in which these communities influence the long-term carbon cycle. As society makes critical decisions about the use and conservation of deep-ocean ecosystems, it is important that we recognize the vulnerability of life on the ocean floor to climate-related stressors, and the direct influence that the surface climate can exert on the world’s largest biome.
This is a contribution from the Deep-Ocean Stewardship Initiative (DOSI), and International Network for scientific investigation of DEEP-sea ecosystems (INDEEP) project that was awarded by the Total Foundation.
We thank the Norwegian Research Council for awarding funding to A.K. Sweetman, L.A. Levin, A.R. Thurber and C.R. Smith to run the workshop “CLIDEEP – Workshop to explore the impacts of climate change on deep-sea pelagic and benthic ecosystems” (NFR grant No. 216598) at Friday Harbor Laboratories, University of Washington, where the foundations for this paper were laid. A.K. Sweetman D.O.B. Jones and R. Danovaro acknowledge funding from the European Union Seventh Framework Programme (FP7/2007–2013) under grant agreement 603418 (MIDAS), and the European Union Horizon 2020 research and innovation programme under grant agreement 689518 (MERCES). L.-A. Henry and J.M. Roberts acknowledge funding from the European Union’s Horizon 2020 research and innovation programme under grant agreement No 678760 (ATLAS): this output reflects only the authors’ views and the European Union cannot be held responsible for any use that may be made of the information contained therein.
The authors have no competing interests to declare.
Altenbach AV Pflaumann U Schiebel R Thies A Timm S et al. 1999. Scaling percentages and distributional patterns of benthic foraminifera with flux rates of organic carbon. J Foraminiferal Res 29: 173–185.
Alvsvåg J, Agnalt AL and Jørstad KE 2009. Evidence for a permanent establishment of the snow crab (Chionoecetes opilio) in the Barents Sea. Biol Invasions 11: 587–595, DOI: http://dx.doi.org/10.1007/s10530-008-9273-7
Amon DJ Ziegler AF Dahlgren TG Glover AG Goineau A et al. 2016. First insights into the abundance and diversity of abyssal megafauna in a polymetallic-nodule region in the eastern Clarion-Clipperton Zone. Sci Rep 6: 30492.DOI: http://dx.doi.org/10.1038/srep30492
Armstrong CW, Foley NS, Tinch R and van den Hove S 2012. Services from the deep: Steps towards valuation of deep sea goods and services. Ecosys Serv 2: 2–13, DOI: http://dx.doi.org/10.1016/j.ecoser.2012.07.001
Arrigo KR 2013. The changing Arctic Ocean. Elem Sci Anth 1(1): 000010.DOI: http://dx.doi.org/10.12952/journal.elementa.000010
Arrigo KR and van Dijken GL 2011. Secular trends in Arctic Ocean net primary production. J Geophys Res 116: C09011.DOI: http://dx.doi.org/10.1029/2011JC007151
Arrigo KR, van Dijken G and Pabi S 2008. Impact of a shrinking Arctic ice cover on marine primary production. Geophys Res Lett 35: L19603.DOI: http://dx.doi.org/10.1029/2008GL035028
Baillon S, Hamel JF, Wareham VE and Mercier A 2012. Deep cold-water corals as nurseries for fish larvae. Front Ecol Environ 10: 351–356, DOI: http://dx.doi.org/10.1890/12002
Bakun A 1990. Global climate change and intensification of coastal ocean upwelling. Science 247: 198–201, DOI: http://dx.doi.org/10.1126/science.247.4939.198
Bakun A, Field D, Redondo-Rodriguez A and Weeks S 2010. Greenhouse gas, upwelling-favorable winds, and the future of coastal ocean upwelling ecosystems. Global Change Biol 16: 1213–28, DOI: http://dx.doi.org/10.1111/j.1365-2486.2009.02094.x
Behrenfeld MJ and Falkowski PG 1997. Photosynthetic rates derived from satellite-based chlorophyll concentration. Limnol Oceanogr 42: 1–20, DOI: http://dx.doi.org/10.4319/lo.1997.42.1.0001
Beman JM, Chow CE, King AL, Fuhrman JA, Andersson A, Bates NR, Bopp BN and Hutchins DA 2011. Global declines in oceanic nitrification rates as a consequence of ocean acidification. Proc Natl Acad Sci USA 108(1): 208–213, DOI: http://dx.doi.org/10.1073/pnas.1011053108
Berelson WM Anderson RF Dymond J Demaster D Hammond DE et al. 1997. Biogenic budgets of particle rain, benthic remineralization and sediment accumulation in the equatorial Pacific. Deep-Sea Res 44: 2251–2281, DOI: http://dx.doi.org/10.1016/S0967-0645(97)00030-1
Bianchi GG and McCave IN 1999. Holocene periodicity in North Atlantic climate and deep-ocean flow south of Iceland. Nature 397: 515–517, DOI: http://dx.doi.org/10.1038/17362
Biastoch A Treude T Rüpke LH Riebesell U Roth C et al. 2011. Rising Arctic Ocean temperatures cause gas hydrate destabilization and ocean acidification. Geophys Res Lett 38: L08602.DOI: http://dx.doi.org/10.1029/2011GL047222
Billett DSM, Bett BJ, Reid WDK, Boorman B and Priede IG 2001. Long-term change in the abyssal NE Atlantic: The ‘Amperima Event’ revisited. Deep-Sea Res II 57: 1406–1417, DOI: http://dx.doi.org/10.1016/j.dsr2.2009.02.001
Bluhm BA, Gradinger R and Hopcroft RR 2011. Arctic Ocean diversity: Synthesis. Mar Biodiv 41: 1–4, DOI: http://dx.doi.org/10.1007/s12526-010-0080-x
Bluhm BA, Iken K, Hardy SM, Sirenko BI and Holladay BA 2009. Community structure of epibenthic megafauna in the Chukchi Sea. Aqu Biol 7: 269–293, DOI: http://dx.doi.org/10.3354/ab00198
Boetius A and Wenzhöfer F 2013. Seafloor oxygen consumption fueled by methane from cold seeps. Nature Geosci 6: 725–734, DOI: http://dx.doi.org/10.1038/ngeo1926
Bopp L, Aumont O, Cadule P, Alvain S and Gehlen M 2005. Response of diatoms distribution to global warming and potential implications: A global model study. Geophys Res Lett 32: L19606.DOI: http://dx.doi.org/10.1029/2005GL023653
Bopp L Monfray P Aumont O Dufresne J-L Le Treut H et al. 2001. Potential impact of climate change on marine export production. Global Biogeochem Cycles 15: 81–99, DOI: http://dx.doi.org/10.1029/1999GB001256
Bopp L Resplandy L Orr JC Doney SC Dunne JP et al. 2013. Multiple stressors of ocean ecosystems in the 21st century: Projections with CMIP5 models. Biogeosci 10: 6225–6245, DOI: http://dx.doi.org/10.5194/bg-10-6225-2013
Brandt A, Linse K and Schüller M 2009. Bathymetric distribution patterns of Southern Ocean macrofaunal taxa: Bivalvia, Gastropoda, Isopoda and Polychaeta. Deep-Sea Res I 56: 2013–2025, DOI: http://dx.doi.org/10.1016/j.dsr.2009.06.007
Buesseler KO Doney SC Karl DM Boyd PW Caldeira K et al. 2008. Ocean iron fertilization: Moving forward in a sea of uncertainty. Science 319: 162.DOI: http://dx.doi.org/10.1126/science.1154305
Buesseler KO Lamborg CH Boyd PW Lam PJ Trull TW et al. 2007. Revisiting carbon flux through the ocean’s twilight zone. Science 316: 567–570, DOI: http://dx.doi.org/10.1126/science.1137959
Buhl-Mortensen L Vanreusel A Gooday AJ Levin LA Priede IG et al. 2010. Biological structures as a source of habitat heterogeneity and biodiversity on the deep ocean margins. Mar Ecol 31: 21–50, DOI: http://dx.doi.org/10.1111/j.1439-0485.2010.00359.x
Byrne RH, Mecking S, Feely RA and Liu X 2010. Direct observations of basin-wide acidification of the North Pacific Ocean. Geophys Res Lett 37: L02601.DOI: http://dx.doi.org/10.1029/2009GL040999
Canals M Puig P Durrieu de Madron X Heussner S Palanques A et al. 2006. Flushing submarine canyons. Nature 444: 354–357, DOI: http://dx.doi.org/10.1038/nature05271
Cannariato KG and Kennett JP 1999. Climatically related millennial-scale fluctuations in strength of California margin oxygen-minimum zone during the past 60 ky. Geology 27: 975–978, DOI: http://dx.doi.org/10.1130/0091-7613(1999)027<0975:CRMSFI>2.3.CO;2
Cannariato KG, Kennett JP and Behl RJ 1999. Biotic response to late Quaternary rapid climate switches in Santa Barbara Basin: Ecological and evolutionary implications. Geology 27: 63–66, DOI: http://dx.doi.org/10.1130/0091-7613(1999)027<0063:BRTLQR>2.3.CO;2
Caroletti GN and Barstad I 2010. An assessment of future extreme precipitation in western Norway using a linear model. Hydrol Earth Sys Sci 14: 2329–2341, DOI: http://dx.doi.org/10.5194/hess-14-2329-2010
Cerrano C Cardini U Bianchelli S Corinaldesi C Pusceddu A et al. 2013. Red coral extinction risk enhanced by ocean acidification. Sci Rep 3: 1457.DOI: http://dx.doi.org/10.1038/srep01457
Chapman WL and Walsh JE 2007. A synthesis of Antarctic temperatures. J Climate 20: 4096–4117, DOI: http://dx.doi.org/10.1175/JCLI4236.1
Cheung WWL Sarmiento JL Dunne J Frölicher TL Lam VWY et al. 2013. Shrinking of fishes exacerbates impacts of global ocean changes on marine ecosystems. Nature Clim Change 3: 254–258, DOI: http://dx.doi.org/10.1038/nclimate1691
Clarke A Murphy EJ Meredith MP King JC Peck LS et al. 2007. Climate change and the marine ecosystem of the western Antarctic Peninsula. Phil Trans Roy Soc B 362: 149–166, DOI: http://dx.doi.org/10.1098/rstb.2006.1958
Coleman FC and Williams SL 2002. Overexploiting marine ecosystem engineers: Potential consequences for biodiversity. Trends Ecol Evol 17: 40–44, DOI: http://dx.doi.org/10.1016/S0169-5347(01)02330-8
Cordes EE Cunha MR Galéron J Mora C Olu-Le Roy K et al. 2010. The influence of geological, eochemical, and biogenic habitat heterogeneity on seep diversity. Mar Ecol 31: 51–65, DOI: http://dx.doi.org/10.1111/j.1439-0485.2009.00334.x
Cordes EE Jones DOB Schlacher TA Amon DJ Bernadino AF et al. 2016. Environmental impacts of the deep-water oil and gas industry: A review to guide management strategies. Front Environ S 4: 58.DOI: http://dx.doi.org/10.3389/fenvs.2016.00058
Corliss BH, Brown CW, Sun X and Showers WJ 2009. Deep-sea benthic diversity linked to seasonality of pelagic productivity. Deep-Sea Res I 56: 835–841, DOI: http://dx.doi.org/10.1016/j.dsr.2008.12.009
Cornelius N and Gooday AJ 2004. ‘Live’ (stained) deep-sea benthic foraminiferans in the western Weddell Sea: Trends in abundance, diversity, and taxonomic composition along a depth transect. Deep-Sea Res II 51: 1571–1602, DOI: http://dx.doi.org/10.1016/j.dsr2.2004.06.024
Cronin TM Dwyer GS Farmer J Bauch HA Spielhagen RF et al. 2012. Deep Arctic Ocean warming during the last glacial cycle. Nature Geosci 5: 631–634, DOI: http://dx.doi.org/10.1038/ngeo1557
Cronin TM and Raymo ME 1997. Orbital forcing of deep-sea benthic species diversity. Nature 385: 624–627, DOI: http://dx.doi.org/10.1038/385624a0
Cronin TM, Raymo ME and Kyle KP 1996. Pliocene (3.2–2.4 Ma) ostracode faunal cycles and deep ocean circulation, North Atlantic Ocean. Geology 24: 695–698, DOI: http://dx.doi.org/10.1130/00917613(1996)024<0695:PMOFCA>2.3.CO;2
Culver SJ and Buzas MA 2000. Global latitudinal species diversity gradient in deep-sea benthic foraminifera. Deep-Sea Res I 47: 259–275, DOI: http://dx.doi.org/10.1016/S0967-0637(99)00055-2
Danovaro R, Dell’Anno A, Fabiano M, Pusceddu A and Tselepides A 2001. Deep-sea ecosystem response to climate changes: the eastern Mediterranean case study. Trends Ecol Evol 16: 505–510, DOI: http://dx.doi.org/10.1016/S0169-5347(01)02215-7
Danovaro R Gambi C Dell’Anno A Corinaldesi C Fraschetti S et al. 2008. Exponential decline of deep-sea ecosystem functioning linked to benthic biodiversity loss. Curr Biol 18: 1–8, DOI: http://dx.doi.org/10.1016/j.cub.2007.11.056
Danovaro R, Molari M, Corinaldesi C and Dell’Anno A 2016. Macroecological drivers of archaea and bacteria in benthic deep-sea ecosystems. Sci Adv 2: e1500961.DOI: http://dx.doi.org/10.1126/sciadv.1500961
Danovaro R, Snelgrove PVR and Tyler P 2014. Challenging the paradigms of deep-sea ecology. Trends Ecol Evol 29: 465–475, DOI: http://dx.doi.org/10.1016/j.tree.2014.06.002
Dayton PK 1990. Polar benthos, in, Polar oceanography In: London: Academic, pp. 631–685, DOI: http://dx.doi.org/10.1016/B978-0-08-092595-0.50010-6
De Leo FC, Smith CR, Rowden AA, Bowden DA and Clark MA 2010. Submarine canyons: Hotspots of benthic biomass and productivity in the deep sea. Proc Roy Soc B 277: 2783–2792, DOI: http://dx.doi.org/10.1098/rspb.2010.0462
De Rijk S, Jorissen FJ, Rohling EJ and Troelstra SR 2000. Organic flux on bathymetric zonation of Mediterranean benthic Foraminifera. Mar Micropaleo 40: 151–166, DOI: http://dx.doi.org/10.1016/S03778398(00)00037-2
Deutsch C Berelson W Thunell R Weber T Tems C et al. 2014. Centennial changes in North Pacific anoxia linked to tropical trade winds. Science 345: 665–668, DOI: http://dx.doi.org/10.1126/science.1252332
Deutsch C, Brix H, Ito T, Frenzel H and Thompson L 2011. Climate-forced variability of ocean hypoxia. Science 333: 336–339, DOI: http://dx.doi.org/10.1126/science.1202422
Deutsch C, Ferrel A, Seibel B, Pörtner H-O and Huey RB 2015. Climate change tightens a metabolic constraint on marine habitats. Science 348: 1132–1135, DOI: http://dx.doi.org/10.1126/science.aaa1605
Diaz RJ and Rosenberg R 2008. Spreading dead zones and consequences for marine ecosystems. Science 321: 926–929, DOI: http://dx.doi.org/10.1126/science.1156401
Dunlop KM van Oevelen D Ruhl HA Huffard CL Kuhnz LA et al. 2016. Carbon cycling in the deep eastern North Pacific benthic food web: Investigating the effect of organic carbon input. Limnol Oceanogr, DOI: http://dx.doi.org/10.1002/lno.10345
Dwyer GS Cronin TM Baker PA Raymo ME Buzas JS et al. 1995. North Atlantic deepwater temperature change during late Pliocene and late Quaternary climatic cycles. Science 270: 1347–1351, DOI: http://dx.doi.org/10.1126/science.270.5240.1347
Elderfield H Greaves M Baker S Hall IR Tripati A et al. 2010. A record of bottom water temperature and seawater δ18O for the Southern Ocean over the past 440 kyr based on Mg/Ca of benthic foraminiferal Uvigerina spp. Quat Sci Rev 29: 160–169, DOI: http://dx.doi.org/10.1016/j.quascirev.2009.07.013
Etnoyer P and Morgan LE 2005. Habitat-forming deep-sea corals in the Northeast Pacific Ocean In: Freiwald, A and Roberts, JM eds. Cold-water corals and ecosystems. Berlin: Springer-Verlag, pp. 331–343, DOI: http://dx.doi.org/10.1007/3-540-27673-4_16
Farmer JR Cronin TM de Vernal A Dwyer GS Keigwin LD et al. 2011. Western Arctic Ocean temperature variability during the last 8000 years. Geophys Res Lett 38: L24602.DOI: http://dx.doi.org/10.1029/2011GL049714
Feely RA, Sabine CL, Hernandez-Ayon JM, Ianson D and Hales B 2008. Evidence for upwelling of corrosive “acidified” water onto the continental shelf. Science 320: 1490–1492, DOI: http://dx.doi.org/10.1126/science.1155676
Findlay HS Hennige SJ Wicks LC Moreno NJ Malcolm E et al. 2014. Fine-scale nutrient and carbonate system dynamics around cold-water coral reefs in the northeast Atlantic. Sci Rep 4: 3671.DOI: http://dx.doi.org/10.1038/srep03671
Fischlin A Midgley GF Price J Leemans R Gopal B et al. 2007. Ecosystems, their properties, goods, and services In: Parry, ML, Canziani, OF, Palutikof, JP, van der Linden, PJ and Hanson, CE eds. Climate change 2007: impacts, adaptation and vulnerability. Contribution of Working group II to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge: Cambridge University Press, pp. 211–272.
Fosså JH, Buhl-Mortensen P and Furevik DM 2002. The deep-water coral Lophelia pertusa in Norwegian waters: distribution and fishery impacts. Hydrobiologia 471: 1–12, DOI: http://dx.doi.org/10.1023/A:1016504430684
Froescheis O, Looser R, Cailliet GM, Jarman WM and Ballschmiter K 2000. The deep-sea as a final global sink for semivolatile persistent organic pollutants? Part I: PCBs in surface and deep-sea dwelling fish of the North and South Atlantic and the Monterey Bay Canyon (California). Chemosphere 40: 651–660, DOI: http://dx.doi.org/10.1016/S0045-6535(99)00461-0
Gehlen M Séférian R Jones DOB Roy T Roth R et al. 2014. Projected pH reductions by 2100 might put deep North Atlantic biodiversity at risk. Biogesci 11: 6955–6967, DOI: http://dx.doi.org/10.5194/bg-11-6955-2014
Gilly WF, Beman JM, Litvin SY and Robison BH 2013. Oceanographic and biological effects of shoaling of the oxygen minimum zone. Ann Rev Mar Sci 5: 393–420, DOI: http://dx.doi.org/10.1146/annurev-marine-120710-100849
Glecker PJ, Durack PJ, Stouffer RJ, Johnson GC and Forest CE 2016. Industrial-era global ocean heat uptake doubles in recent decades. Nature Clim Change 6: 394–398, DOI: http://dx.doi.org/10.1038/nclimate2915
Glover AG Gooday AJ Bailey DM Billett DSM Chevaldonne P et al. 2010. Temporal change in deep-sea benthic ecosystems: A review of the evidence from recent time-series studies. Adv Mar Biol 58: 1–95, DOI: http://dx.doi.org/10.1016/B978-0-12-381015-1.00001-0
Gooday AJ 2003. Benthic foraminifera (Protista) as tools in deep-water palaeoceanography: Environmental influences on faunal characteristics. Adv Mar Biol 46: 1–90, DOI: http://dx.doi.org/10.1016/S0065-2881(03)46002-1
Gooday AJ Bett BJ Escobar E Ingole B Levin LA et al. 2010. Habitat heterogeneity and its influence on benthic biodiversity in oxygen minimum zones. Mar Ecol 31: 125–147, DOI: http://dx.doi.org/10.1111/j.1439-0485.2009.00348.x
Gooday AJ, Bett BJ, Jones DOB and Kitazato H 2012. The influence of productivity on abyssal foraminiferal biodiversity. Mar Biodiv 42: 415–431, DOI: http://dx.doi.org/10.1007/s12526-012-0121-8
Gooday AJ, Goineau A and Voltski I 2015. Abyssal foraminifera attached to polymetallic nodules from the eastern Clarion Clipperton Fracture Zone: a preliminary description and comparison with North Atlantic dropstone assemblages. Mar Biodiv 45: 391–412, DOI: http://dx.doi.org/10.1007/s12526-014-0301-9
Gooday AJ, Nomaki H and Kitazato H 2008. Modern deep-sea benthic foraminifera: A brief review of their biodiversity and trophic diversity In: Austin, WEN and James, RH eds. Biogeochemical controls on palaeoceanographic environmental proxies. London: Geological Society, 303pp. 97–119, DOI: http://dx.doi.org/10.1144/SP303.8
Grange LJ and Smith CR 2013. Megafaunal communities in rapidly warming fjords along the West Antarctic Peninsula: Hotspots of abundance and beta diversity. PLoS ONE 8: e77917.DOI: http://dx.doi.org/10.1371/journal.pone.0077917
Gregg WW, Conkright ME, Ginoux P, O’Reilly JE and Casey NW 2003. Ocean primary production and climate: Global decadal changes. Geophys Res Lett 30: L016889.DOI: http://dx.doi.org/10.1029/2003GL016889
Guinotte JM Orr J Cairns S Freiwald A Morgan L et al. 2006. Will human-induced changes in seawater chemistry alter the distribution of deep-sea scleractinian corals?. Front Ecol Environ 4: 141–146, DOI: http://dx.doi.org/10.1890/1540-9295(2006)004[0141:WHCISC]2.0.CO;2
Gutt J and Piepenburg D 2003. Scale-depended impact on diversity of Antarctic benthos caused by grounding of icebergs. Mar Ecol Prog Ser 253: 77–83, DOI: http://dx.doi.org/10.3354/meps253077
Gutt J, Starmans A and Dieckman G 1996. Impact of iceberg scouring on polar benthic habitats. Mar Ecol Prog Ser 137: 311–316, DOI: http://dx.doi.org/10.3354/meps137311
Hasemann C Bergmann M Kanzog C Lochthofen N Sauter E et al. 2013. Effects of dropstone-induced habitat heterogeneity on Arctic deep-sea benthos with special reference to nematode communities. Mar Biol Res 9: 229–245, DOI: http://dx.doi.org/10.1080/17451000.2012.739694
Helly J and Levin LA 2004. Global distribution of naturally occurring marine hypoxia on continental margins. Deep-Sea Res 51: 1159–1168, DOI: http://dx.doi.org/10.1016/j.dsr.2004.03.009
Helm KP, Bindoff NL and Church JA 2011. Observed decreases in oxygen content of the global ocean. Geophys Res Lett 38: L23602.DOI: http://dx.doi.org/10.1029/2011GL049513
Hennige SJ Wicks LC Kamenos NA Perna G Findlay HS et al. 2015. Hidden impacts of ocean acidification to live and dead coral framework. Proc Roy Soc B 282: 20150990.DOI: http://dx.doi.org/10.1098/rspb.2015.0990
Henry L-A Moreno NJ Hennige SJ Wicks LC Vad J et al. 2013. Cold-water coral reef habitats benefit recreationally valuable sharks. Biol Conserv 161: 67–70, DOI: http://dx.doi.org/10.1016/j.biocon.2013.03.002
Henry L-A and Roberts JM 2007. Biodiversity and ecological composition of macrobenthos on coldwater coral mounds and adjacent off-mound habitat in the bathyal Porcupine Seabight, NE Atlantic. Deep-Sea Res I 54: 654–672, DOI: http://dx.doi.org/10.1016/j.dsr.2007.01.005
Henry L-A, Stehmann M, De Clippele L, Golding N and Roberts JM 2016. Seamount egg-laying grounds of the deepwater skate Bathyraja richardsoni (Garrick 1961). J Fish Biol 89: 1473–1481, DOI: http://dx.doi.org/10.1111/jfb.13041
Henson SA, Sanders R and Madsen E 2012. Global patterns in efficiency of particulate organic carbon export and transfer to the deep ocean. Global Biogeochem Cycles 26: GB1028.DOI: http://dx.doi.org/10.1029/2011GB004099
Hoegh-Guldberg O and Bruno JF 2010. The impact of climate change on the world’s marine ecosystems. Science 328: 1523–1528, DOI: http://dx.doi.org/10.1126/science.1189930
Hofmann M and Schellnhuber HJ 2009. Oceanic acidification affects marine carbon pump and triggers extended marine oxygen holes. Proc Nat Acad Sci USA 106: 3017–3022, DOI: http://dx.doi.org/10.1073/pnas.0813384106
Huffard CL, Kuhnz LA, Lemon L, Sherman AD and Smith KL 2016. Demographic indicators of change in a deposit-feeding abyssal holothurian community (station M, 4000 m). Deep-Sea Res 109: 27–39, DOI: http://dx.doi.org/10.1016/j.dsr.2016.01.002
Ingels J Van den Driessche P De Mesel I Vanhove S Moens T et al. 2010. Preferred use of bacteria over phytoplankton by deep-sea nematodes in polar regions. Mar Ecol Prog Ser 406: 121–133, DOI: http://dx.doi.org/10.3354/meps08535
Ingels J Vanreusel A Brandt A Catarino AI David B et al. 2012. Possible effects of global environmental changes on Antarctic benthos: a synthesis across five major taxa. Ecol Evol 2: 453–485, DOI: http://dx.doi.org/10.1002/ece3.96
Jaedicke C, Lied K and Kronholm K 2009. Integrated database for rapid mass movements in Norway. Nat Hazards Earth Sys Sci 9: 469–479, DOI: http://dx.doi.org/10.5194/nhess-9-469-2009
Johnson HP, Miller UK, Salmi MS and Solomon EA 2015. Analysis of bubble plume distributions to evaluate methane hydrate decomposition on the continental slope. Geochem Geophys Geosys 16: 3825–3839, DOI: http://dx.doi.org/10.1002/2015GC005955
Jones DOB Yool A Wei C-L Henson SA Ruhl HA et al. 2014. Global reductions in seafloor biomass in response to climate change. Global Change Biol 20: 1861–1872, DOI: http://dx.doi.org/10.1111/gcb.12480
Keeling RF, Körtzinger A and Gruber N 2010. Ocean deoxygenation in a warming world. Annu Rev Mar Sci 2: 199–229, DOI: http://dx.doi.org/10.1146/annurev.marine.010908.163855
Koslow JW Boehlert GW Gordon JDM Haedrich RL Lorance P et al. 2000. Continental slope and deep-sea fisheries: Implications for a fragile ecosystem. ICES J Mar S 57: 548–557, DOI: http://dx.doi.org/10.1006/jmsc.2000.0722
Koslow JA, Goericke R, Lara-Lopez A and Watson W 2011. Impact of declining intermediate-water oxygen on deepwater fishes in the California Current. Mar Ecol Prog Ser 436: 207–218, DOI: http://dx.doi.org/10.3354/meps09270
Lauerman LM and Kaufmann RS 1998. Deep-sea epibenthic echinoderms and a temporally varying food supply: Results from a one year time series in the NE Pacific. Deep-Sea Res II 45: 817–842, DOI: http://dx.doi.org/10.1016/S0967-0645(98)00004-6
Laws EA, Falkowski PG, Smith WO, Ducklow H and McCarthy JJ 2000. Temperature effects on export production in the open ocean. Global Biogeochem Cycles 14: 1231–1246, DOI: http://dx.doi.org/10.1029/1999GB001229
Lebrato M Molinero J-C Cartes JE Lloris D Mélin F et al. 2013. Sinking jelly-carbon unveils potential environmental variability along a continental margin. PLoS ONE 8: e82070.DOI: http://dx.doi.org/10.1371/journal.pone.0082070
Levin LA and Dayton PK 2009. Ecological theory and continental margins: where shallow meets deep. Trends Ecol Evol 24: 606–617, DOI: http://dx.doi.org/10.1016/j.tree.2009.04.012
Levin LA and Le Bris N 2015. The deep ocean under climate change. Science 350: 766–768, DOI: http://dx.doi.org/10.1126/science.aad0126
Levin LA, Mengerink K, Gjerde KM, Rowden AA, Van Dover CL, Clark MR, Ramirez-Llodra E, Currie B, Smith CR, Sato KN, Gallo N, Sweetman AK, Lily H, Armstrong CW and Brider J 2016. Defining “Serious Harm” to the marine environment in the context of Deep-Seabed Mining. Mar Policy 74: 245–259, DOI: http://dx.doi.org/10.1016/j.marpol.2016.09.032
Levin LA, Rathburn AE, Gutiérrez D, Muñoz P and Shankle A 2003. Bioturbation by symbiont bearing annelids in near-anoxic sediments: implications for biofacies model and paleo oxygen assessments. Palaeogeogr Palaeoclimatol Palaeoecol 199: 129–140, DOI: http://dx.doi.org/10.1016/S0031-0182(03)00500-5
Levin LA and Sibuet M 2012. Understanding continental margin biodiversity: a new imperative. Ann Rev Mar Sci 4: 79–112, DOI: http://dx.doi.org/10.1146/annurev-marine-120709-142714
Levin LA, Whitcraft CR, Mendoza GF, Gonzalez JP and Cowie G 2009. Oxygen and organic matter thresholds for benthic faunal activity on the Pakistan margin oxygen minimum zone (700–1100 m). Deep-Sea Res II 56: 449–471, DOI: http://dx.doi.org/10.1016/j.dsr2.2008.05.032
Levitus S, Antonov J and Boyer T 2005. Warming of the world ocean, 1955–2003. Geophys Res Lett 32: L02604.DOI: http://dx.doi.org/10.1029/2004GL021592
Levitus S, Antonov JI, Boyer TP and Stephens C 2000. Warming of the world ocean. Science 287: 2225–2229, DOI: http://dx.doi.org/10.1126/science.287.5461.2225
Li WKW, Carmack EC, McLaughlin FA, Nelson RJ and Williams WJ 2013. Space-for-time substitution in predicting the state of picoplankton and nanoplankton in a changing Arctic Ocean. J Geophys Res 118: 5750–5759, DOI: http://dx.doi.org/10.1002/jgrc.20417
Li WKW, McLaughlin FA, Lovejoy C and Carmack EC 2009. Smallest algae thrive as the Arctic Ocean freshens. Science 326: 539–539, DOI: http://dx.doi.org/10.1126/science.1179798
Loeb V Siegel V Holm-Hansen O Hewitt R Fraser W et al. 1997. Effects of sea-ice extent and krill or salp dominance on the Antarctic food web. Nature 387: 897–900, DOI: http://dx.doi.org/10.1038/43174
Long MC, Deutsch C and Ito T 2016. Finding forced trends in oceanic oxygen. Global Biogeochem Cycles 30: 381–397, DOI: http://dx.doi.org/10.1002/2015GB005310
Looser R, Froescheis O, Cailliet GM, Jarman WM and Ballschmiter K 2000. The deep-sea as a final organic sink of semivolatile persistent organic pollutants? Part II: Organochlorine pesticides in surface and deep-sea dwelling fish of the North and South Atlantic and the Monterey Bay Canyon (California). Chemosphere 40: 661–670, DOI: http://dx.doi.org/10.1016/S0045-6535(99)00462-2
Lutz MJ, Caldeira K, Dunbar RB and Behrenfeld MJ 2007. Seasonal rhythms of net primary production and particulate organic carbon flux to depth describe the efficiency of biological pump in the global ocean. J Geophys Res 112: C10011.DOI: http://dx.doi.org/10.1029/2006JC003706
Marchitto TM and deMenocal PB 2003. Late Holocene variability of upper North Atlantic Deep Water temperature and salinity. Geochem Geophys Geosys 4: 12.DOI: http://dx.doi.org/10.1029/2003GC000598
Martin PA Lea DW Rosenthal Y Shackleton NJ Sarnthein M et al. 2002. Quaternary deep sea temperature histories derived from benthic foraminiferal Mg/Ca. Earth Planet Sci Lett 198: 193–209, DOI: http://dx.doi.org/10.1016/S0012-821X(02)00472-7
McClatchie S, Goericke R, Cosgrove R and Vetter R 2010. Oxygen in the Southern California Bight: multidecadal trends and implications for demersal fisheries. Geophys Res Lett 37: L044497.DOI: http://dx.doi.org/10.1029/2010GL044497
McLeod RJ and Wing SR 2007. Hagfish in the New Zealand fjords are supported by chemoautotrophy of forest carbon. Ecology 88: 809–816, DOI: http://dx.doi.org/10.1890/06-1342
McLeod R, Wing S and Skilton J 2010. High incidence of invertebrate–chemoautotrophic symbioses in benthic communities of New Zealand fjords. Limnol Oceanogr 55: 2097–2106, DOI: http://dx.doi.org/10.4319/lo.2010.55.5.2097
Meinshausen M Smith SJ Calvin K Daniel JS Kainuma MLT et al. 2011. The RCP greenhouse gas concentrations and their extensions from 1765 to 2300. Clim Change 109: 213–241, DOI: http://dx.doi.org/10.1007/s10584-011-0156-z
Mengerink KJ Van Dover CL Ardron J Baker M Escobar-Briones E et al. 2014. A call for deep-ocean stewardship. Science 344: 696–698, DOI: http://dx.doi.org/10.1126/science.1251458
Merino M and Monreal-Gómez MA 2009. Ocean currents and their impact on marine life In: Marine Ecology. Oxford: Encyclopedia of Life Support Systems (EOLSS) Developed under the Auspices of the UNESCO, Eolss Publishers, pp. 47–52.
Meyer KS, Sweetman AK, Young CM and Renaud PE 2015. Environmental factors structuring Arctic mega-benthos – a case study from a shelf and two fjords. Front Mar Sci 2: 22.DOI: http://dx.doi.org/10.3389/fmars.2015.00022
Meyer KS, Young CM, Sweetman AK, Taylor J, Bergmann M and Soltwedel T 2016. Rocky islands in a sea of mud: Biotic and abiotic factors structuring deep-sea dropstone communities. Mar Ecol Prog Ser 556: 45–57, DOI: http://dx.doi.org/10.3354/meps11822
Mincks SL, Smith CR and DeMaster DJ 2005. Persistence of labile organic matter and microbial biomass in Antarctic shelf sediments: Evidence of a sediment “food bank.”. Mar Ecol Prog Ser 300: 3–19, DOI: http://dx.doi.org/10.3354/meps300003
Moffitt SE, Hill TM, Roopnarine PD and Kennett JP 2015a. Response of seafloor ecosystems to abrupt global climate change. Proc Nat Acad Sci USA 112: 4684–4689, DOI: http://dx.doi.org/10.1073/pnas.1417130112
Moffitt SE Moffitt RA Sauthoff W Davis CV Hewett K et al. 2015b. Paleoceanographic insights on recent Oxygen Minimum Zone expansion: lessons for modern oceanography. PLoS ONE 10: e0115246.DOI: http://dx.doi.org/10.1371/journal.pone.0115246
Mora C Wei C-L Rollo A Amaro T Baco AR et al. 2013. Biotic and human vulnerability to projected changes in ocean biogeochemistry over the 21st century. PLoS Biology 11: e1001682.DOI: http://dx.doi.org/10.1371/journal.pbio.1001682
Morán XAG Alonso-Sáez L Nogueira E Ducklow HW González N et al. 2015. More, smaller bacteria in response to ocean’s warming?. Proc Roy Soc B 282: 20150371.DOI: http://dx.doi.org/10.1098/rspb.2015.0371
Morán XAG, Lopez-Urrutia A, Calvo-Diaz A and Li WKW 2010. Increasing importance of small phytoplankton in a warmer ocean. Global Change Biol 16: 1137–1144, DOI: http://dx.doi.org/10.1111/j.1365-2486.2009.01960.x
Morel A and Berthon JF 1989. Surface pigments, algal biomass profiles, and potential production of the euphotic layer: Relationships reinvestigated in view of remote-sensing applications. Limnol Oceanogr 34: 1545–1562, DOI: http://dx.doi.org/10.4319/lo.1918.104.22.1685
Nozawa F, Kitazato H, Tsuchiya M and Gooday AJ 2006. ‘Live’ benthic foraminifera at an abyssal site in the equatorial Pacific nodule province: abundance, diversity and taxonomic composition. Deep-Sea Res I 51: 1406–1422, DOI: http://dx.doi.org/10.1016/j.dsr.2006.06.001
Orr JC Fabry VJ Aumont O Bopp L Doney SC et al. 2005. Anthropogenic ocean acidification over the twenty-first century and its impact on calcifying organisms. Nature 437: 681–686, DOI: http://dx.doi.org/10.1038/nature04095
Overland JE, Spillane MC, Percival DB, Wang M and Mofjeld HO 2004. Seasonal and regional variation of pan-arctic surface air temperature over the instrumental record. J Climate 17: 3263–3282, DOI: http://dx.doi.org/10.1175/1520-0442(2004)017<3263:SARVOP>2.0.CO;2
Pachauri RK Allen MR Barros VR Broome J Cramer W et al. 2014. Climate Change 2014: Synthesis Report. Contribution of Working Groups I, II and III to the Fifth Assessment Report of the Intergovernmental Panel on Climate Change In: Pachauri, R and Meyer, L eds. Geneva, Switzerland: IPCC, p. 151.978-92-9169-143-2
Philippart CJM Anadón R Danovaro R Dippner JW Drinkwater KF et al. 2011. Impacts of climate change on European marine ecosystems: observations, expectations and indicators. J Exp Mar Biol Ecol 400: 52–69, DOI: http://dx.doi.org/10.1016/j.jembe.2011.02.023
Phrampus BJ and Hornbach MJ 2012. Recent changes to the Gulf Stream causing widespread gas hydrate destabilization. Nature 490: 527–530, DOI: http://dx.doi.org/10.1038/nature11528
Praetorius SK Mix AC Walczak MH Wolhowe MD Addison JA et al. 2015. North Pacific deglacial hypoxic events linked to abrupt ocean warming. Nature 527: 362–366, DOI: http://dx.doi.org/10.1038/nature15753
Prince ED and Goodyear CP 2006. Hypoxia-based habitat compression of tropical pelagic fishes. Fish Oceanogr 15: 451–464, DOI: http://dx.doi.org/10.1111/j.1365-2419.2005.00393.x
Pörtner HO 2012. Integrating climate-related stressor effects on marine organisms: Unifying principles linking molecule to ecosystem-level changes. Mar Ecol Prog Ser 470: 273–290, DOI: http://dx.doi.org/10.3354/meps10123
Pörtner HO and Knust R 2007. Climate change affects marine fishes through the oxygen limitation of thermal tolerance. Science 315: 95–97, DOI: http://dx.doi.org/10.1126/science.1135471
Purkey SG and Johnson GC 2010. Warming of global abyssal and deep Southern Ocean waters between the 1990s and 2000s: Contributions to global heat and sea level rise budgets. J Clim 23: 6336–6351, DOI: http://dx.doi.org/10.1175/2010JCLI3682.1
Rahmstorf S Box JE Feulner G Mann ME Robinson A et al. 2015. Exceptional twentieth-century slowdown in Atlantic Ocean overturning circulation. Nature Clim Change 5: 475–480, DOI: http://dx.doi.org/10.1038/nclimate2554
Ramirez LE Tyler PA Baker MC Bergstad OA Clark MR et al. 2011. Man and the last great wilderness: Human impact on the deep sea. PLoS ONE 6: e22588.DOI: http://dx.doi.org/10.1371/journal.pone.0022588
Rasmussen TL, Thomsen E, Troelstra SR, Kuijpers A and Prins MA 2003. Millennial-scale glacial variability versus Holocene stability: Changes in planktic and benthic foraminifera faunas and ocean circulation in the North Atlantic during the last 60,000 years. Mar Micropaleontol 47: 143–176, DOI: http://dx.doi.org/10.1016/S0377-8398(02)00115-9
Reid PC Fischer AC Lewis-Brown E Meredith MP Sparrow M et al. 2009. Impacts of the oceans on climate change. Adv Mar Biol 56: 1–150, DOI: http://dx.doi.org/10.1016/S0065-2881(09)560014
Renaud PE Riedel A Michel C Morata N Gosselin M et al. 2007. Seasonal variation in benthic community oxygen demand: A response to an ice algal bloom in the Beaufort Sea, Canadian Arctic?. J Mar Sys 67: 1–12, DOI: http://dx.doi.org/10.1016/j.jmarsys.2006.07.006
Roberts JM Davies AJ Henry L-A Dodds LA Duineveld GCA et al. 2009a. Mingulay reef complex: An interdisciplinary study of cold-water coral habitat, hydrography and biodiversity. Mar Ecol Prog Ser 397: 139–151, DOI: http://dx.doi.org/10.3354/meps08112
Roberts JM, Wheeler AJ and Freiwald A 2006. Reefs of the deep: the biology and geology of cold water coral ecosystems. Science 312: 543–547, DOI: http://dx.doi.org/10.1126/science.1119861
Roberts JM, Wheeler AJ, Freiwald A and Cairns SD 2009b. Cold-water corals: The biology and geology of deep-sea coral habitats In: Cambridge: Cambridge University Press, DOI: http://dx.doi.org/10.1017/CBO9780511581588
Roder C, Berumen ML, Bouwmeester J, Al-Suwailem AEP and Voolstra CR 2013. First biological measurements of deep-sea corals from the Red Sea. Sci Rep 3: 2802.DOI: http://dx.doi.org/10.1038/srep02802
Roemmich D Church J Gilson J Monselesan D Sutton P et al. 2015. Unabated planetary warming and its ocean structure since 2006. Nature Clim Change 5: 240–245, DOI: http://dx.doi.org/10.1038/nclimate2513
Rosa R and Seibel BA 2008. Synergistic effects of climate-related variables suggest future physiological impairment in a top oceanic predator. Proc Nat Acad Sci USA 105: 20776–20780, DOI: http://dx.doi.org/10.1073/pnas.0806886105
Ruhl HA Bett BJ Hughes SJM Alt CHS Ross EJ et al. 2014. Links between deep-sea respiration and community dynamics. Ecology 95: 1651–1662, DOI: http://dx.doi.org/10.1890/13-0675.1
Ruhl HA and Smith KL 2004. Shifts in deep-sea community structure linked to climate and food supply. Science 305: 513–515, DOI: http://dx.doi.org/10.1126/science.1099759
Sahade R Lagger C Torre L Momo F Monien P et al. 2015. Climate change and glacier retreat drive shifts in an Antarctic benthic ecosystem. Sci Adv 1: e1500050.DOI: http://dx.doi.org/10.1126/sciadv.1500050
Saidova HM 1965. Sediment stratigraphy and paleogeography of the Pacific Ocean by benthonic foraminifera during the Quaternary. Prog Oceanogr 4: 143–151, DOI: http://dx.doi.org/10.1016/0079-6611(65)90047-9
Sato KN, Levin LA and Schiff K 2016. Habitat compression and expansion of sea urchins in response to changing climate conditions on the California continental shelf and slope (1994–2013). Deep-Sea Res II, DOI: http://dx.doi.org/10.1016/j.dsr2.2016.08.012 in press.
Schmittner A, Galbraith ED, Hostetler SW, Pedersen TF and Zhang R 2007. Large fluctuations of dissolved oxygen in the Indian and Pacific oceans during Dansgaard-Oeschger oscillations caused by variations of North Atlantic Deep Water Subduction. Palaeoceanography 22: PA3207.DOI: http://dx.doi.org/10.1029/2006PA001384
Schulz M, Bergmann M, von Juterzenka K and Soltwedel T 2010. Colonisation of hard substrata along a channel system in the deep Greenland Sea. Polar Biol 33: 1359–1369, DOI: http://dx.doi.org/10.1007/s00300-010-0825-9
Smale DA and Barnes DKA 2008. Likely responses of the Antarctic benthos to climate-related changes in physical disturbance during the 21st century, based primarily on evidence from the West Antarctic Peninsula region. Ecography 31: 289–305, DOI: http://dx.doi.org/10.1111/j.0906-7590.2008.05456.x
Smith CR Berelson W Demaster DJ Dobbs FC Hammond D et al. 1997. Latitudinal variations in benthic processes in the abyssal equatorial Pacific: control by biogenic particle flux. Deep-Sea Res 44: 2295–2317, DOI: http://dx.doi.org/10.1016/S0967-0645(97)00022-2
Smith CR, De Leo FC, Bernardino AF, Sweetman AK and Arbizu PM 2008. Abyssal food limitation, ecosystem structure and climate change. Trends Ecol Evol 23: 518–528, DOI: http://dx.doi.org/10.1016/j.tree.2008.05.002
Smith CR and DeMaster DJ 2008. Preface and brief synthesis for the FOODBANCS volume. Deep-Sea Res II 55: 2399–2403, DOI: http://dx.doi.org/10.1016/j.dsr2.2008.08.001
Smith CR, Grange LJ, Honig DL, Naudts L, Huber B, Guidi L and Domack E 2011. A large population of king crabs in Palmer Deep on the west Antarctic Peninsula shelf and potential invasive impacts. Proc Roy Soc B, DOI: http://dx.doi.org/10.1098/rspb.2011.1496
Smith CR, Levin LA, Hoover DJ, McMurtry G and Gage JD 2000. Variations in bioturbation across the oxygen minimum zone in the northwest Arabian Sea. Deep-Sea Res II 47: 227–258, DOI: http://dx.doi.org/10.1016/S0967-0645(99)00108-3
Smith KL Baldwin RJ Ruhl HA Kahru M Mitchell BG et al. 2006. Climate effect on food supply to depths greater than 4,000 meters in the northeast Pacific. Limnol Oceanogr 51: 166–176, DOI: http://dx.doi.org/10.4319/lo.2006.51.1.0166
Smith KL Robison BH Helly JJ Kaufmann RS Ruhl HA et al. 2007. Free-drifting icebergs: Hotspots of chemical and biological enrichment in the Weddell Sea. Science 317: 478–482, DOI: http://dx.doi.org/10.1126/science.1142834
Smith KL Ruhl HA Bett BJ Billett DSM Lampitt RS et al. 2009. Climate, carbon cycling, and deep-ocean ecosystems. Proc Nat Acad S USA 106: 19211–19218, DOI: http://dx.doi.org/10.1073/pnas.0908322106
Smith KL, Ruhl HA, Kahru M, Huffard CL and Sherman AD 2013. Deep ocean communities impacted by changing climate over 24 y in the abyssal northeast Pacific Ocean. Proc Nat Acad Sci USA 110: 19838–19841, DOI: http://dx.doi.org/10.1073/pnas.1315447110
Smith KL Sherman AD Huffard CL McGill PR Henthorn R et al. 2014. Large salp bloom export from the upper ocean and benthic community response in the abyssal northeast Pacific: day to week resolution. Limnol Oceanogr 59: 745–757, DOI: http://dx.doi.org/10.4319/lo.2014.59.3.0745
Soetaert K, Mohn C, Rengstorf A, Grehan A and van Oevelen D 2016. Ecosystem engineering creates a direct nutritional link between 600-m deep cold-water coral mounds and surface productivity. Sci Rep 6: 35057.DOI: http://dx.doi.org/10.1038/srep35057
Solomon S, Portmann RW and Thompson DW 2007. Contrasts between Antarctic and Arctic ozone depletion. Proc Nat Acad Sci USA 104: 445–449, DOI: http://dx.doi.org/10.1073/pnas.0604895104
Soltwedel T Bauerfeind E Bergmann M Budaeva N Hoste E et al. 2005. HAUSGARTEN: multidisciplinary investigations at a deep-sea, long-term observatory in the Arctic Ocean. Oceanography 18: 46–61, DOI: http://dx.doi.org/10.5670/oceanog.2005.24
Somavilla R, Schauer U and Budéus G 2013. Increasing amount of Arctic Ocean deep waters in the Greenland Sea. Geophys Res Lett 40: 4361–4366, DOI: http://dx.doi.org/10.1002/grl.50775
Sosdian S and Rosenthal Y 2009. Deep-sea temperature and ice volume changes across the Pliocene-Pleistocene climate transitions. Science 325: 306–310, DOI: http://dx.doi.org/10.1126/science.1169938
Sperling EA, Frieder CA and Levin LA 2016. Biodiversity response to natural gradients of multiple stressors on continental margins. Proc Roy Soc B 283: 20160637.DOI: http://dx.doi.org/10.1098/rspb.2016.063
Spielhagen RF Werner K Sørensen SA Zamelczyk K Kandiano E et al. 2011. Enhanced modern heat transfer to the Arctic by warm Atlantic water. Science 331: 450–453, DOI: http://dx.doi.org/10.1126/science.1197397
Steinacher M Joos F Frölicher TL Bopp L Cadule P et al. 2010. Projected 21st century decrease in marine productivity: A multi-model analysis. Biogeosciences 7: 979–1005, DOI: http://dx.doi.org/10.5194/bg-7-979-2010
Stendardo I, Kieke D, Rhein M, Gruber N and Steinfeldt R 2015. Interannual to decadal oxygen variability in the mid-depth water masses of the eastern North Atlantic. Deep-Sea Res I 95: 85–98, DOI: http://dx.doi.org/10.1016/j.dsr.2014.10.009
Stramma L, Johnson GC, Sprintall J and Mohrholz V 2008. Expanding oxygen-minimum zones in the tropical oceans. Science 320: 655–658, DOI: http://dx.doi.org/10.1126/science.1153847
Stramma L Prince ED Schmidtko S Luo J Hoolihan JP et al. 2012. Expansion of oxygen minimum zones may reduce available habitat for tropical pelagic fishes. Nature Clim Change 2: 33–37, DOI: http://dx.doi.org/10.1038/nclimate1304
Stramma L, Schmidt S, Levin LA and Johnson GC 2010. Ocean oxygen minima expansions and their biological impacts. Deep-Sea Res 210: 587–595, DOI: http://dx.doi.org/10.1016/j.dsr.2010.01.005
Sun X, Corliss BH, Brown CW and Showers WJ 2006. The effect of primary productivity and seasonality on the distribution of deep-sea benthic foraminifera in the North Atlantic. Deep-Sea Res 53: 28–47, DOI: http://dx.doi.org/10.1016/j.dsr.2005.07.003
Sweetman AK Chelsky A Pitt KA Andrade H van Oevelen D et al. 2016. Jellyfish decomposition at the seafloor rapidly alter biogeochemical cycling and carbon flow through benthic food-webs. Limnol Oceanogr 61: 1449–1461, DOI: http://dx.doi.org/10.1002/lno.10310
Sweetman AK, Norling K, Gunderstad C, Haugland B and Dale T 2014b. Benthic ecosystem functioning beneath fish farms in different hydrodynamic environments. Limnol Oceanogr 59: 1139–1151, DOI: http://dx.doi.org/10.4319/lo.2014.59.4.1139
Sweetman AK, Smith CR, Dale T and Jones DOB 2014a. Rapid scavenging of jellyfish carcasses reveals the importance of gelatinous material to deep-sea food webs. Proc Roy Soc B 281: 20142210.DOI: http://dx.doi.org/10.1098/rspb.2014.2210
Sydeman WJ García-Reyes M Schoeman DS Rykaczewski RR Thompson SA et al. 2014. Climate change and wind intensification in coastal upwelling ecosystems. Science 345: 77–80, DOI: http://dx.doi.org/10.1126/science.1251635
Syvitski JPM 1989. On the deposition of sediment with glacier-influenced fjords: oceanographic controls. Mar Geol 85: 301–329, DOI: http://dx.doi.org/10.1016/0025-3227(89)90158-8
Syvitski JP, Farrow GE, Atkinson RJA, Moore PG and Andrews JT 1989. Baffin Island fjord macrobenthos: bottom communities and environmental significance. Arctic 42: 232–247, DOI: http://dx.doi.org/10.14430/arctic1662
Takatani Y, Sasano D, Nakano T and Midorikawa T 2012. Decrease of dissolved oxygen after the mid-1980s in the western North Pacific subtropical gyre along the 137° E repeat section. Global Biogeochem Cycles 26: GB2013.DOI: http://dx.doi.org/10.1029/2011GB004227
Taylor KE, Stouffer RJ and Meehl GA 2012. An overview of the CMIP5 and the experiment design. Bull Amer Meteorol Soc 93: 485–498, DOI: http://dx.doi.org/10.1175/BAMS-D-11-00094.1
Thomas E and Gooday AJ 1996. Cenozoic deep-sea benthic foraminifers: Tracers for changes in oceanic productivity?. Geology 24: 355–358, DOI: http://dx.doi.org/10.1130/0091-7613(1996)024<0355:CDSBFT>2.3.CO;2
Thurber AR, Sweetman AK, Narayanaswamy BE, Jones DOB, Ingels J and Hansman RL 2014. Ecosystem functions and services in the deep sea. Biogeosci 11: 3941–3963, DOI: http://dx.doi.org/10.5194/bg-11-3941-2014
Tittensor DP, Baco AR, Hall-Spenser JM, Orr JC and Rogers AD 2010. Seamounts as refuge from ocean acidification for cold-water corals. Mar Ecol 31: 1212–1225, DOI: http://dx.doi.org/10.1111/j.14390485.2010.00393.x
Tittensor DP, Rex MA, Stuart CT, McClain CR and Smith CR 2011. Species-energy relationships in deep-sea molluscs. Biol Lett, DOI: http://dx.doi.org/10.1098/rsbl.2010.1174
Tremblay JÉ, Robert D, Varela DE, Lovejoy C, Darnis G, Nelson RJ and Sastri AR 2012. Current state and trends in Canadian Arctic marine ecosystems: I. Primary production. Climatic Change 115(1): 161–178, DOI: http://dx.doi.org/10.1007/s10584-012-0496-3
Treude T, Kiel S, Linke P, Peckmann J and Goedert JL 2011. Elasmobranch egg capsules associated with modern and ancient cold seeps: A nursery for marine deep-water predators. Mar Ecol Prog Ser 437: 175–181, DOI: http://dx.doi.org/10.3354/meps09305
van Nugteren P, Herman PMJ, Moodley L, Middelburg JJ, Vos M and Heip CHR 2009. Spatial distribution of detrital resources determines the outcome of competition between bacteria and a facultative detritivorous worm. Limnol Oceanogr 54: 1413–1419, DOI: http://dx.doi.org/10.4319/lo.2009.54.5.1413
van Oevelen D Duineveld G Lavaleye M Mienis F Soeteart K et al. 2009. The cold-water coral community as a hot spot for carbon cycling on continental margins: A food-web analysis from Rockall Bank (northeast Atlantic). Limnol Oceanogr 54: 1829–1844, DOI: http://dx.doi.org/10.4319/lo.2009.54.6.1829
Vanreusel A, Hilario A, Ribeiro PA, Menot L and Arbizu PM 2016. Threatened by mining, polymetallic nodules are required to preserve abyssal epifauna. Sci Rep 6: 26808.DOI: http://dx.doi.org/10.1038/srep26808
Wang D, Gouhier TC, Menge BA and Ganguly AR 2015. Intensification and spatial homogenization of coastal upwelling under climate change. Nature 518: 390–394, DOI: http://dx.doi.org/10.1038/nature14235
Wang S, Bailey D, Lindsay K, Moore JK and Holland M 2014. Impact of sea ice on the marine iron cycle and phytoplankton productivity. Biogeosci 11: 4713–4731, DOI: http://dx.doi.org/10.5194/bg-11-4713-2014
Watling L, Guinotte J, Clark MR and Smith CR 2013. A proposed biogeography of the deep ocean floor. Progr Oceanogr 111: 91–112, DOI: http://dx.doi.org/10.1016/j.pocean.2012.11.003
Wedding LM Reiter SM Smith CR Gjerde KM Kittinger JN et al. 2015. Managing mining of the deep seabed. Science 349: 144–145, DOI: http://dx.doi.org/10.1126/science.aac6647
Wenzhöfer F Adler M Kohls O Hensen C Strotmann B et al. 2001. Calcite dissolution driven by benthic mineralization in the deep-sea: in situ measurements of Ca2+, pH, pCO2 and O2. Geochim Cosmochim Acta 65: 2677–2690, DOI: http://dx.doi.org/10.1016/S0016-7037(01)00620-2
Węsławski JM Kendall MA Włodarska-Kowalczuk M Iken K Kędra M et al. 2011. Climate change effects on Arctic fjord and coastal macrobenthic diversity—observations and predictions. Mar Biodiv 41: 71–85, DOI: http://dx.doi.org/10.1007/s12526-010-0073-9
Włodarska-Kowalczuk M, Pearson TH and Kendall MA 2005. Benthic response to chronic natural physical disturbance by glacial sedimentation in an Arctic fjord. Mar Ecol Prog Ser 303: 31–41, DOI: http://dx.doi.org/10.3354/meps303031
Wood HL, Spicer JI and Widdicombe S 2008. Ocean acidification may increase calcification rates, but at a cost. Proc Roy Soc B 275: 1767–1773, DOI: http://dx.doi.org/10.1098/rspb.2008.0343
Woulds C, Cowie G, Andersson JH, Middelburg JJ and Levin LA 2009. 13C tracer studies on the short-term fate of organic carbon in marine sediments: comparing the Pakistan margin to other regions. Deep-Sea Res 56: 393–402, DOI: http://dx.doi.org/10.1016/j.dsr2.2008.10.008
Yamamoto A Abe-Ouchi A Shigemitsu M Oka A Takahashi K et al. 2015. Global deep ocean oxygenation by enhanced ventilation in the Southern Ocean under long-term global warming. Global Biogeochem Cycles 29: 1801–1815, DOI: http://dx.doi.org/10.1002/2015GB005181
Yasuhara M and Cronin TM 2008. Climatic influences on deep-sea ostracode (Crustacea) diversity for the last three million years. Ecology 89: S52–S65, DOI: http://dx.doi.org/10.1890/07-1021.1
Yasuhara M, Cronin TM, deMenocal PB, Okahashi H and Linsley BK 2008. Abrupt climate change and collapse of deep-sea ecosystems. Proc Nat Acad Sci USA 105: 1556–1560, DOI: http://dx.doi.org/10.1073/pnas.0705486105
Yasuhara M and Danovaro R 2016. Temperature impacts on deep-sea biodiversity. Biol Rev 91: 275–287, DOI: http://dx.doi.org/10.1111/brv.12169
Yasuhara M, Doi H, Wei CL, Danovaro R and Myhre SE 2016. Biodiversity-ecosystem functioning relationships in long-term time series and palaeoecological records: deep sea as a test bed. Phil Trans Roy Soc B 371: 20150282.DOI: http://dx.doi.org/10.1098/rstb.2015.0282
Yasuhara M Hunt G Cronin TM Hokanishi N Kawahata H et al. 2012a. Climatic forcing of Quaternary deep-sea benthic communities in the North Pacific Ocean. Paleobiology 38: 162–179, DOI: http://dx.doi.org/10.1666/10068.1
Yasuhara M, Hunt G, Cronin TM and Okahashi H 2009. Temporal latitudinal-gradient dynamics and tropical instability of deep-sea species diversity. Proc Nat Acad Sci USA 106: 21717–21720, DOI: http://dx.doi.org/10.1073/pnas.0910935106
Yasuhara M Hunt G van Dijken G Arrigo KR Cronin TM et al. 2012b. Patterns and controlling factors of species diversity in the Arctic Ocean. J Biogeogr 39: 2081–2088, DOI: http://dx.doi.org/10.1111/j.1365-2699.2012.02758.x
Yasuhara M, Okahashi H, Cronin TM, Rasmussen TL and Hunt G 2014. Response of deep-sea biodiversity to deglacial and Holocene abrupt climate changes in the North Atlantic Ocean. Global Ecol Biogeogr 23: 957–967, DOI: http://dx.doi.org/10.1111/geb.12178
Yasuhara M, Tittensor DP, Hillebrand H and Worm B 2015. Combining marine macroecology and palaeoecology in understanding biodiversity: microfossils as a model. Biol Rev, DOI: http://dx.doi.org/10.1111/brv.12223
Yesson C Taylor ML Tittensor DP Davies AJ Guinotte J et al. 2012. Global habitat suitability of cold-water octocorals. J Biogeogr 39: 1278–1292, DOI: http://dx.doi.org/10.1111/j.13652699.2011.02681.x
Yool A, Martin AP, Fernandez C and Clark DR 2007. The significance of nitrification for oceanic new production. Nature 447: 999–1002, DOI: http://dx.doi.org/10.1038/nature05885
Zeppilli D, Pusceddu A, Trincardi F and Danovaro R 2016. Seafloor heterogeneity influences the biodiversity–ecosystem functioning relationships in the deep sea. Sci Rep 6: 26352.DOI: http://dx.doi.org/10.1038/srep26352