Editing IPCC:AR6/WGII/Chapter-3 (section)

== 3.2 Observed Trends and Projections of Climatic Impact-Drivers in the Global Ocean ==

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=== 3.2.1 Introduction ===

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Climate change exposes ocean and coastal ecosystems to changing environmental conditions, including ocean warming, SLR, acidification, deoxygenation and other climatic impact-drivers (CIDs), which have distinct regional and temporal characteristics ( [[#Gruber--2011|Gruber, 2011]] ; [[#IPCC--2018|IPCC, 2018]] ). This section aims to build on the WGI AR6 assessment (Table 3.2) to provide an ecosystem-oriented framing of CIDs. Updating SROCC, projected trends assessed here are based on a new range of scenarios (Shared Socioeconomic Pathways, SSPs), as used in the Coupled Model Intercomparison Project Phase 6 (CMIP6; [[IPCC:Wg2:Chapter:Chapter-1#1.2.2|Section 1.2.2]] ).

'''Table 3.2 |''' Overview of the main global ocean climatic impact-drivers and their observed and projected trends from WGI AR6, with corresponding confidence levels and links to WGI chapters where these trends are assessed in detail

{| class="wikitable"
|-
! Climatic impact-drivers (hazards)

! Observed trends over the historical period

! WGI section

! Projected trends over the 21st century

! WGI section

|-
| ''Ocean temperature''

|
|-
| Ocean warming

| ‘At the ocean surface, temperature has on average increased by 0.88 [0.68–1.01] °C from 1850–1900 to 2011–2020.’

| 2.3.3.1, 9.2.1

( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| Ocean warming will continue over the 21st century ( ''virtually certain'' ), with the rate of global ocean warming starting to be scenario-dependent from about the mid-21st century ( ''medium confidence'' ).

| 9.2.1 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

|-
| Marine heatwaves (MHWs)

| MHWs became more frequent ( ''high confidence'' ), more intense and longer ( ''medium confidence'' ) over the 20th and early 21st centuries.

| Box 9.2 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

| MHWs will become ‘4 [2–9, ''likely'' range] times more frequent in 2081–2100 compared with 1995–2014 under SSP1-2.6, and 8 [3–15, ''likely'' range] times more frequent under SSP5-8.5.’

| Box 9.2 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

|-
| Climate velocities

| Not assessed in WGI

|
| Not assessed in WGI

|
|-
| ''Sea level''

|
|-
| Global mean sea level (GMSL)

| ‘Since 1901, GMSL has risen by 0.20 [0.15–0.25] m’, and the rate of rise is accelerating.

| 2.3.3, 9.6.1 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| There will be continued rise in GMSL throughout the 21st century under all assessed SSPs ( ''virtually certain'' ) ''.''

| 4.3.2.2, 9.6.3 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Lee--2021|Lee et al., 2021]] )

|-
| Extreme sea levels

| Relative sea level rise is driving a global increase in the frequency of extreme sea levels ( ''high confidence'' ).

| 9.6.4 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

| Rising mean relative sea level will continue to drive an increase in the frequency of extreme sea levels ( ''high confidence'' ).

| 9.6.4 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

|-
| ''Ocean circulation''

|
|-
| Ocean stratification

| ‘The upper ocean has become more stably stratified since at least 1970 […] ( ''virtually certain'' ) ''.'' ’

| 9.2.1.3 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

| ‘Upper-ocean stratification will continue to increase throughout the 21st century ( ''virtually certain'' ) ''.'' ’

| 9.2.1.3 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

|-
| Eastern boundary upwelling systems

| ‘Only the California current system has experienced some large-scale upwelling-favourable wind intensification since the 1980s ( ''medium confidence'' ) ''.'' ’

| 9.2.5 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

| ‘Eastern boundary upwelling systems will change, with a dipole spatial pattern within each system of reduction at low latitude and enhancement at high latitude ( ''high confidence'' ) ''.'' ’

| 9.2.5 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

|-
| Atlantic overturning circulation (AMOC)

| There is ''low confidence'' in reconstructed and modelled AMOC changes for the 20 th century.

| 2.3.3.4, 9.2.3 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| The AMOC will decline over the 21st century ( ''high confidence'' , but ''low confidence'' for quantitative projections).

| 4.3.2.3, 9.2.3 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Lee--2021|Lee et al., 2021]] )

|-
| ''Sea ice''

|
|-
| Arctic sea ice changes

| ‘Current Arctic sea ice coverage levels are the lowest since at least 1850 for both annual mean and late-summer values ( ''high confidence'' ).’

| 2.3.2.1, 9.3.1 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| ‘The Arctic will become practically ice-free in September by the end of the 21st century under SSP2-4.5, SSP3-7.0 and SSP5-8.5[…]( ''high confidence'' ).’

| 4.3.2.1, 9.3.1 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Lee--2021|Lee et al., 2021]] )

|-
| Antarctic sea ice changes

| There is no global significant trend in Antarctic sea ice area from 1979 to 2020 ( ''high confidence'' ).

| 2.3.2.1, 9.3.2 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| There is ''low confidence'' in model simulations of future Antarctic sea ice.

| 9.3.2 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

|-
| ''Ocean chemistry''

|
|-
| Changes in salinity

| The ‘large-scale, near-surface salinity contrasts have intensified since at least 1950 […] ( ''virtually certain'' ).’

| 2.3.3.2, 9.2.2.2 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| ‘Fresh ocean regions will continue to get fresher and salty ocean regions will continue to get saltier in the 21st century ( ''medium confidence'' ).’

| 9.2.2.2 ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] )

|-
| Ocean acidification

| Ocean surface pH has declined globally over the past four decades ( ''virtually certain'' ).

| 2.3.3.5, 5.3.2.2 ( [[#Canadell--2021|Canadell et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| Ocean surface pH will continue to decrease ‘through the 21st century, except for the lower-emission scenarios SSP1-1.9 and SSP1-2.6 […] ( ''high confidence'' ).’

| 4.3.2.5, 4.5.2.2, 5.3.4.1 ( [[#Lee--2021|Lee et al., 2021]] ; [[#Canadell--2021|Canadell et al., 2021]] )

|-
| Ocean deoxygenation

| Deoxygenation has occurred in most open ocean regions since the mid-20th century ( ''high confidence'' ).

| 2.3.3.6, 5.3.3.2 ( [[#Canadell--2021|Canadell et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] )

| Subsurface oxygen content ‘is projected to transition to historically unprecedented condition with decline over the 21st century ( ''medium confidence'' ).’

| 5.3.3.2 ( [[#Canadell--2021|Canadell et al., 2021]] )

|-
| Changes in nutrient concentrations

| Not assessed in WGI

|
| Not assessed in WGI

|
|}

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<span id="physical-changes"></span>
=== 3.2.2 Physical Changes ===

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==== 3.2.2.1 Ocean Warming, Climate Velocities and Marine Heatwaves ====

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Global mean SST has increased since the beginning of the 20th century by 0.88°C ( ''very likely'' range: 0.68–1.01°C), and it is ''virtually certain'' that the global ocean has warmed since at least 1971 (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). A key characteristic of ocean temperature change relevant for ecosystems is climate velocity, a measure of the speed and direction at which isotherms move under climate change ( [[#Burrows--2011|Burrows et al., 2011]] ), which gives the rate at which species must migrate to maintain constant climate conditions. It has been shown to be a useful and simple predictor of species distribution shifts in marine ecosystems ( [[#Chen--2011|Chen et al., 2011]] ; [[#Pinsky--2013|Pinsky et al., 2013]] ; [[#Lenoir--2020|Lenoir et al., 2020]] ). Median climate velocity in the surface ocean has been 21.7 km per decade since 1960, with higher values in the Arctic/sub-Arctic and within 15° of the Equator (Figure 3.3; [[#Burrows--2011|Burrows et al., 2011]] ). While climate velocity has been slower in the mesopelagic layer (200–1000 m) than in the epipelagic layer (0–200 m) over the past 50 years, it has been shown to be faster in the bathypelagic (1000–4000 m) and abyssopelagic (&gt;4000 m) layers (Figure 3.4; [[#Brito-Morales--2020|Brito-Morales et al., 2020]] ), suggesting that deep-ocean species could be as exposed to effects of warming as species in the surface ocean ( [[#Brito-Morales--2020|Brito-Morales et al., 2020]] ).

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[[File:708c67e23788e6c29827206356733693 IPCC_AR6_WGII_Figure_3_003.png]]

'''Figure 3.3 |''' '''Observed surface ocean warming, surface climate velocity and reconstructed changes in marine heatwaves (MHWs) over the past 100 years.''' (a) Sea surface temperature trend (degrees Celsius per century) over 1925–2016 from Hadley Centre Sea Ice and Sea Surface Temperature 1.1 (HadISST1.1; (b) surface climate velocity (kilometres per decade) over 1925–2016 computed from HadISST1.1 and (c) change in total MHW days for the surface ocean over 1925–1954 to 1987–2016 based on monthly proxies. (Data from [[#Oliver--2018|Oliver et al., 2018]] ).

Marine heatwaves (MHWs) are periods of extreme seawater temperature relative to the long-term mean seasonal cycle, that persist for days to months, and that may carry severe consequences for marine ecosystems and their services (WGI AR6 Box 9.2; [[#Hobday--2016a|Hobday et al., 2016a]] ; [[#Smale--2019|Smale et al., 2019]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). MHWs became more frequent over the 20th century ( ''high confidence'' ) and into the beginning of the 21st century, approximately doubling in frequency ( ''high confidence'' ) and becoming more intense and longer since the 1980s ( ''medium confidence'' ) (WGI AR6 Box 9.2; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). These trends in MHWs are explained by an increase in ocean mean temperatures ( [[#Oliver--2018|Oliver et al., 2018]] ), and human influence has ''very likely'' contributed to 84–90% of them since at least 2006 (WGI AR6 Box 9.2; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). The probability of occurrence (as well as duration and intensity) of the largest and most impactful MHWs that have occurred in the past 30 years has increased more than 20-fold due to anthropogenic climate change ( [[#Laufkötter--2020|Laufkötter et al., 2020]] ).

Ocean warming will continue over the 21st century ( ''virtually certain'' ), with the rate of global ocean warming starting to be scenario-dependent from about the mid-21st century ( ''medium confidence'' ). At the ocean surface, it is ''virtually certain'' that SST will continue to increase throughout the 21st century, with increasing hazards to many marine ecosystems (WGI AR6 Box 9.2; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ) ''.'' The future global mean SST increase projected by CMIP6 models for the period 1995–2014 to 2081–2100 is 0.86°C ( ''very likely'' range: 0.43–1.47°C) under SSP1-2.6, 1.51°C (1.02–2.19°C) under SSP2-4.5, 2.19°C (1.56–3.30°C) under SSP3-7.0 and 2.89°C (2.01–4.07°C) under SSP5-8.5 (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2.1]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). Stronger surface warming occurs in parts of the tropics, in the North Pacific, and in the Arctic Ocean, where SST increases by &gt;4°C in 2080–2099 under SSP5-8.5 ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). The CMIP6 climate models also project ocean warming at the seafloor, with the magnitude of projected changes being less than that of surface waters but having larger uncertainties ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). The projected end-of-the-century warming in CMIP6 as reported here is greater than assessed with Coupled Model Intercomparison Project 5 (CMIP5) models in AR5 and in SROCC for similar radiative forcing scenarios (Figure 3.5; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ), because of greater climate sensitivity in the CMIP6 model ensemble than in CMIP5 (WGI AR6 Chapter 4; [[#Forster--2020|Forster et al., 2020]] ; [[#Lee--2021|Lee et al., 2021]] ).

Marine heatwaves will continue to increase in frequency, with a ''likely'' global increase of 2–9 times in 2081–2100 compared with 1995–2014 under SSP1-2.6, and 3–15 times under SSP5-8.5, with the largest increases in tropical and Arctic oceans (WGI AR6 Box 9.2; [[#Frölicher--2018|Frölicher et al., 2018]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

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[[File:47f328cb87e020e334c82d4c176308d0 IPCC_AR6_WGII_Figure_3_004.png]]

'''Figure 3.4 |''' '''Historical and projected climate velocity.''' Climate velocities (in kilometres per decade) are shown for the '''(a,d,g)''' historical period (1965–2014), and the last 50 years of the 21st century (2051–2100), under '''(b,e,h)''' SSP1-2.6 and '''(c,f,i)''' SSP5-8.5. Also shown are the epipelagic (0–200 m), mesopelagic (200–1000 m) and bathypelagic (1000–4000 m) domains. Updated figure from [[#Brito-Morales--2020|Brito-Morales et al. (2020)]] , with Coupled Model Intercomparison Project 6 models used in [[#Kwiatkowski--2020|Kwiatkowski et al. (2020)]] .

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<span id="sea-level-rise-and-extreme-sea-levels"></span>
==== 3.2.2.2 Sea Level Rise and Extreme Sea Levels ====

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Global mean sea level (GMSL) (Cross-Chapter Box SLR in Chapter 3) has risen by about 0.20 m since 1901 and continues to accelerate (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] ; [[#Church--2011|Church and White, 2011]] ; [[#Jevrejeva--2014|Jevrejeva et al., 2014]] ; [[#Hay--2015|Hay et al., 2015]] ; [[#Kopp--2016|Kopp et al., 2016]] ; [[#Dangendorf--2017|Dangendorf et al., 2017]] ; WCRP Global Sea Level Budget Group, 2018; [[#Kemp--2018|Kemp et al., 2018]] ; [[#Ablain--2019|Ablain et al., 2019]] ; [[#Gulev--2021|Gulev et al., 2021]] ).

Most coastal ecosystems (mangroves, seagrasses, salt marshes, shallow coral reefs, rocky shores and sandy beaches) are affected by changes in relative sea level (RSL, the change in the mean sea level relative to the land; [[#3.4.2|Section 3.4.2]] ). Regional rates of RSL rise differ from the global mean due to a range of factors, including local subsidence driven by anthropogenic activities such as groundwater and hydrocarbon extraction (WGI AR6 Box 9.1; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). In many deltaic regions, anthropogenic subsidence is currently the dominant driver of RSL rise (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.6.3|Section 9.6.3.2]] ; [[#Tessler--2018|Tessler et al., 2018]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). RSL rise is driving a global increase in the frequency of extreme sea levels ( ''high confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.6.4.1|Section 9.6.4.1]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

GMSL rise through the middle of the 21st century exhibits limited dependence on emissions scenario; between 1995–2014 and 2050, GMSL is ''likely'' to rise by 0.15–0.23 m under SSP1-1.9 and 0.20–0.30 m under SSP5-8.5 (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.6.3|Section 9.6.3]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). Beyond 2050, GMSL and RSL projections are increasingly sensitive to the differences among emission scenarios. Considering only processes in which there is at least ''medium confidence'' (e.g., thermal expansion, land-water storage, land-ice surface mass balance and some ice-sheet dynamic processes), GMSL between 1995–2014 and 2100 is ''likely'' to rise by 0.28–0.55 m under SSP1-1.9, 0.33–0.61 m under SSP1-2.6, 0.44–0.76 m under SSP2-4.5, 0.55–0.90 m under SSP3-7.0 and 0.63–1.02 m under SSP5-8.5 (Figure 3.5). Under high-emission scenarios, ice-sheet processes in which there is ''low confidence'' and ''deep uncertainty'' might contribute more than one additional metre to GMSL rise by 2100 (WGI AR6 Chapter 9; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

Rising mean RSL will continue to drive an increase in the frequency of extreme sea levels ( ''high confidence'' ). The expected frequency of the current 1-in-100-year extreme sea level is projected to increase by a median of 20–30 times across tide-gauge sites by 2050, regardless of emission scenario ( ''medium confidence'' ). In addition, extreme-sea-level frequency may be affected by changes in tropical cyclone climatology ( ''low confidence'' ), wave climatology ( ''low confidence'' ) and tides ( ''high confidence'' ) associated with climate change and sea level change (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.6.4.2|Section 9.6.4.2]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

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==== 3.2.2.3 Changes in Ocean Circulation, Stratification and Coastal Upwelling ====

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Ocean circulation and its variations are key to the evolution of the physical, chemical and biological properties of the ocean. Vertical mixing and upwelling are critical factors affecting the supply of nutrients to the sunlit ocean and hence the magnitude of primary productivity. Ocean currents not only transport heat, salt, carbon and nutrients, but they also control the dispersion of many organisms and the connectivity between distant populations.

Ocean stratification is an important factor controlling biogeochemical cycles and affecting marine ecosystems. WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2.1.3]] ( [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ) assessed that it is ''virtually certain'' that stratification in the upper 200 m of the ocean has been increasing since 1970. Recent evidence has strengthened estimates of the rate of change ( [[#Yamaguchi--2019|Yamaguchi and Suga, 2019]] ; [[#Li--2020a|Li et al., 2020a]] ; [[#Sallée--2021|Sallée et al., 2021]] ), with an estimated increase of 1.0 ± 0.3% ( ''very likely'' range) per decade over the period 1970–2018 ( ''high confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2.1.3]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ), higher than assessed in SROCC. It is ''very likely'' that stratification in the upper few hundred metres of the ocean will increase substantially in the 21st century in all ocean basins, driven by intensified surface warming and near-surface freshening at high latitudes (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2.1.3]] ; [[#Capotondi--2012|Capotondi et al., 2012]] ; [[#Fu--2016|Fu et al., 2016]] ; [[#Bindoff--2019a|Bindoff et al., 2019a]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

Contrasting changes among the major eastern boundary coastal upwelling systems (EBUS) were identified in AR5 ( [[#Hoegh-Guldberg--2014|Hoegh-Guldberg et al., 2014]] ). While SROCC assessed with ''high confidence'' that three (Benguela, Peru-Humboldt, California) out of the four major EBUS have experienced upwelling-favourable wind intensification in the past 60 years ( [[#Sydeman--2014|Sydeman et al., 2014]] ; [[#Bindoff--2019a|Bindoff et al., 2019a]] ), WGI AR6 revisited this assessment based on evidence showing ''low agreement'' between studies that have investigated trends over past decades ( [[#Varela--2015|Varela et al., 2015]] ). WGI AR6 assessed that only the California Current system has undergone large-scale upwelling-favourable wind intensification since the 1980s ( ''medium confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2.1.5]] ; [[#García-Reyes--2010|García-Reyes and Largier, 2010]] ; [[#Seo--2012|Seo et al., 2012]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

While no consistent pattern of contemporary changes in upwelling-favourable winds emerges from observation-based studies, numerical and theoretical work projects that summertime winds near poleward boundaries of upwelling zones will intensify, while winds near equatorward boundaries will weaken ( ''high confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2.3.5]] ; [[#García-Reyes--2015|García-Reyes et al., 2015]] ; [[#Rykaczewski--2015|Rykaczewski et al., 2015]] ; [[#Wang--2015|Wang et al., 2015]] ; [[#Aguirre--2019|Aguirre et al., 2019]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). Nevertheless, projected future annual cumulative upwelling wind changes at most locations and seasons remain within ±10–20% of present-day values ( ''medium confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.2|Section 9.2.3.5]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

Continuous observation of the Atlantic meridional overturning circulation (AMOC) has improved the understanding of its variability ( [[#Frajka-Williams--2019|Frajka-Williams et al., 2019]] ), but there is ''low confidence'' in the quantification of AMOC changes in the 20th century because of ''low agreement'' in quantitative reconstructed and simulated trends (WGI AR6 Sections 2.3.3, 9.2.3.1; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] ). Direct observational records since the mid-2000s remain too short to determine the relative contributions of internal variability, natural forcing and anthropogenic forcing to AMOC change ( ''high confidence'' ) (WGI AR6 Sections 2.3.3, 9.2.3.1; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] ). Over the 21st century, AMOC will ''very likely'' decline for all SSP scenarios but will not involve an abrupt collapse before 2100 (WGI AR6 Sections 4.3.2, 9.2.3.1; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Lee--2021|Lee et al., 2021]] ).

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<span id="sea-ice-changes"></span>
==== 3.2.2.4 Sea Ice Changes ====

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Sea ice is a key driver of polar marine life, hosting unique ecosystems and affecting diverse marine organisms and food webs through its impact on light penetration and supplies of nutrients and organic matter ( [[#Arrigo--2014|Arrigo, 2014]] ). Since the late 1970s, Arctic sea ice area has decreased for all months, with an estimated decrease of 2 million km 2 (or 25%) for summer sea ice (averaged for August, September and October) in 2010–2019 as compared with 1979–1988 (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.3.1|Section 9.3.1.1]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). For Antarctic sea ice there is no significant global trend in satellite-observed sea ice area from 1979 to 2020 in either winter or summer, due to regionally opposing trends and large internal variability (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.3.2|Section 9.3.2.1]] ; [[#Maksym--2019|Maksym, 2019]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

CMIP6 simulations project that the Arctic Ocean will ''likely'' become practically sea ice free (area below 1 million km 2 ) for the first time before 2050 and in the seasonal sea ice minimum in each of the four emission scenarios SSP1-1.9, SSP1-2.6, SSP2-4.5 and SSP5-8.5 (Figure 3.7; WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.3.2|Section 9.3.2.2]] ; Notze and SIMIP Community, 2020; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ). Antarctic sea ice area is also projected to decrease during the 21st century, but due to mismatches between model simulations and observations, combined with a lack of understanding of reasons for substantial inter-model spread, there is ''low confidence'' in model projections of future Antarctic sea ice changes, particularly at the regional level (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-9#9.3.2|Section 9.3.2.2]] ; [[#Roach--2020|Roach et al., 2020]] ; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ).

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<span id="chemical-changes"></span>
=== 3.2.3 Chemical Changes ===

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<span id="ocean-acidification"></span>
==== 3.2.3.1 Ocean Acidification ====

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The ocean’s uptake of anthropogenic carbon affects its chemistry in a process referred to as ocean acidification, which increases the concentrations of aqueous CO 2 , bicarbonate and hydrogen ions, and decreases pH, carbonate ion concentrations and calcium carbonate mineral saturation states ( [[#Doney--2009|Doney et al., 2009]] ). Ocean acidification affects a variety of biological processes with, for example, lower calcium carbonate saturation states reducing net calcification rates for some shell-forming organisms and higher CO 2 concentrations increasing photosynthesis for some phytoplankton and macroalgal species ( [[#3.3.2|Section 3.3.2]] ).

Direct measurements of ocean acidity from ocean time series, as well as pH changes determined from other shipboard studies, show consistent decreases in ocean surface pH over the past few decades ( ''virtually certain'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3.2|Section 5.3.2.2]] ; [[#Takahashi--2014|Takahashi et al., 2014]] ; [[#Bindoff--2019a|Bindoff et al., 2019a]] ; [[#Sutton--2019|Sutton et al., 2019]] ; [[#Canadell--2021|Canadell et al., 2021]] ).

Since the 1980s, surface ocean pH has declined by a ''very likely'' rate of 0.016–0.020 per decade in the subtropics and 0.002–0.026 per decade in the subpolar and polar zones (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3.2|Section 5.3.2.2]] ; [[#Canadell--2021|Canadell et al., 2021]] ). Typically, the pH of global surface waters has decreased from 8.2 to 8.1 since the pre-industrial era (1750 CE), a trend attributable to rising atmospheric CO 2 ( ''virtually certain'' ) ( [[#Orr--2005|Orr et al., 2005]] ; [[#Jiang--2019|Jiang et al., 2019]] ).

Ocean acidification is also developing in the ocean interior ( ''very high confidence'' ) due to the transport of anthropogenic CO 2 to depth by ocean currents and mixing (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3|Section 5.3.3.1]] ; [[#Canadell--2021|Canadell et al., 2021]] ). There, it leads to the shoaling of saturation horizons of aragonite and calcite ( ''high confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3|Section 5.3.3.1]] ; [[#Canadell--2021|Canadell et al., 2021]] ), below which dissolution of these calcium carbonate minerals is thermodynamically favoured. The calcite or aragonite saturation horizons have migrated upwards in the North Pacific (1–2 m yr –1 over 1991–2006) ( [[#Feely--2012|Feely et al., 2012]] ) and in the Irminger Sea (10–15 m yr –1 for the aragonite saturation horizon over 1991–2016) ( [[#Perez--2018|Perez et al., 2018]] ). In some locations of the western Atlantic Ocean, calcite saturation depth has risen by ~300 m since the pre-industrial era due to increasing concentrations of deep-ocean dissolved inorganic carbon ( [[#Sulpis--2018|Sulpis et al., 2018]] ). In the Arctic, where some coastal surface waters are already undersaturated with respect to aragonite due to the degradation of terrestrial organic matter ( [[#Mathis--2015|Mathis et al., 2015]] ; [[#Semiletov--2016|Semiletov et al., 2016]] ), the deep aragonite saturation horizon shoaled on average 270 ± 60 m during 1765–2005 ( [[#Terhaar--2020|Terhaar et al., 2020]] ).

Detection and attribution of ocean acidification in coastal environments are more difficult than in the open ocean due to larger spatio-temporal variability of carbonate chemistry ( [[#Duarte--2013|Duarte et al., 2013]] ; [[#Laruelle--2017|Laruelle et al., 2017]] ; [[#Torres--2021|Torres et al., 2021]] ) and to the influence of other natural acidification drivers such as freshwater and high-nutrient riverine inputs ( [[#Cai--2011|Cai et al., 2011]] ; [[#Laurent--2017|Laurent et al., 2017]] ; [[#Fennel--2019|Fennel et al., 2019]] ; [[#Cai--2020|Cai et al., 2020]] ) or anthropogenic acidification drivers ( [[#3.1|Section 3.1]] ) like atmospherically deposited nitrogen and sulphur ( [[#Doney--2007|Doney et al., 2007]] ; [[#Hagens--2014|Hagens et al., 2014]] ). Since AR5, the observing network in coastal oceans has expanded substantially, improving understanding of both the drivers and amplitude of observed variability ( [[#Sutton--2016|Sutton et al., 2016]] ). Recent studies indicate that two more decades of observations may be required before anthropogenic ocean acidification emerges over natural variability in some coastal sites and regions (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3|Section 5.3.5.2]] ; [[#Sutton--2019|Sutton et al., 2019]] ; [[#Turk--2019|Turk et al., 2019]] ; [[#Canadell--2021|Canadell et al., 2021]] ).

Mean open-ocean surface pH is projected to decline by 0.08 ± 0.003 ( ''very likely range'' ), 0.17 ± 0.003, 0.27 ± 0.005 and 0.37 ± 0.007 pH units in 2081–2100 relative to 1995–2014, for SSP1-2.6, SSP2-4.5, SSP3-7.0 and SSP5-8.5, respectively (Figure 3.5; WGI AR6 [[IPCC:Wg2:Chapter:Chapter-4#4.3.2|Section 4.3.2]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ; [[#Lee--2021|Lee et al., 2021]] ). Projected changes in surface pH are relatively uniform in contrast with those of other surface-ocean variables, but they are largest in the Arctic Ocean (Figure 3.6; WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3|Section 5.3.4.1]] ; [[#Canadell--2021|Canadell et al., 2021]] ). Similar declines in the concentration of carbonate ions are projected by Earth system models (ESMs; [[#Bopp--2013|Bopp et al., 2013]] ; [[#Gattuso--2015|Gattuso et al., 2015]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). The North Pacific, the Southern Ocean and Arctic Ocean regions will become undersaturated for calcium carbonate minerals first ( [[#Orr--2005|Orr et al., 2005]] ; [[#Pörtner--2014|Pörtner et al., 2014]] ). Concurrent impacts on the seasonal amplitude of carbonate chemistry variables are anticipated (i.e., increased amplitude for ''p'' CO 2 and hydrogen ions, decreased amplitude for carbonate ions; [[#McNeil--2016|McNeil and Sasse, 2016]] ; [[#Kwiatkowski--2018|Kwiatkowski and Orr, 2018]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ).

<div id="_idContainer017" class="Figure"></div>

[[File:9bdda09a395d9b52ee9d20810614cfc3 IPCC_AR6_WGII_Figure_3_005.png]]

'''Figure 3.5 |''' '''Projected trends in climatic impact-drivers for ocean ecosystems.''' Panels (a,b,c,d) represent Coupled Model Intercomparison Project 5 (CMIP5) Representative Concentration Pathway (RCP) and CMIP6 Shared Socioeconomic Pathway (SSP) end-of-century changes in '''(a)''' global sea level; '''(b)''' average surface pH, '''(c)''' subsurface (100–600 m) dissolved oxygen concentration and '''(d)''' euphotic-zone (0–100 m) nitrate (NO 3 ) concentration against anomalies in sea surface temperature. All anomalies are model-ensemble averages over 2080–2099 relative to the 1870–1899 baseline period (from [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ), except for sea level, which shows model-ensemble median in 2100 relative to 1901 (from AR6 WGI Chapter 9). Error bars represent ''very likely'' ranges, except for SLR where they represent ''likely'' ranges. ''Very likely'' ranges for pH changes are too narrow to appear in the figure (see text). Panels (e,f,g,h) show regions where end-of-century projected CMIP6 surface warming exceeds 2°C, where surface ocean pH decline exceeds 0.3, where subsurface dissolved oxygen decline exceeds 30 mmol m -3 and where euphotic-zone (0–100 m) nitrate decline exceeds 1 mmol m -3 in '''(e)''' SSP1-2.6, '''(f)''' SSP2-4.5, '''(g)''' SSP3-7.0 and '''(h)''' SSP5-8.5. All anomalies are 2080–2099 relative to the 1870–1899 baseline period. (Modified from [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ).

Future declines in subsurface pH (Figure 3.6) will be modulated by changes in ocean overturning and water-mass subduction ( [[#Resplandy--2013|Resplandy et al., 2013]] ), and in organic matter remineralisation ( [[#Chen--2017|Chen et al., 2017]] ). In particular, decreases in pH will be less consistent at the seafloor than at the surface and will be linked to the transport of surface anomalies to depth. For example, &gt;20% of the North Atlantic seafloor deeper than 500 m, including canyons and seamounts designated as marine protected areas (MPAs), will experience pH reductions &gt;0.2 by 2100 under RCP8.5 ( [[#Gehlen--2014|Gehlen et al., 2014]] ). Changes in pH in the abyssal ocean (&gt;3000 m deep) are greatest in the Atlantic and Arctic Oceans, with lesser impacts in the Southern and Pacific Oceans by 2100, mainly due to ventilation time scales ( [[#Sweetman--2017|Sweetman et al., 2017]] ).

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<span id="ocean-deoxygenation"></span>
==== 3.2.3.2 Ocean Deoxygenation ====

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Ocean deoxygenation, the loss of oxygen in the ocean, results from ocean warming, through a reduction in oxygen saturation, increased oxygen consumption, increased ocean stratification and ventilation changes ( [[#Keeling--2010|Keeling et al., 2010]] ; [[#IPCC--2019a|IPCC, 2019a]] ). In recent decades, anthropogenic inputs of nutrients and organic matter ( [[#3.1|Section 3.1]] ) have increased the extent, duration and intensity of coastal hypoxia events worldwide ( [[#Diaz--2008|Diaz and Rosenberg, 2008]] ; [[#Rabalais--2010|Rabalais et al., 2010]] ; [[#Breitburg--2018|Breitburg et al., 2018]] ), while pollution-induced atmospheric deposition of soluble iron over the ocean has accelerated open-ocean deoxygenation ( [[#Ito--2016|Ito et al., 2016]] ). Deoxygenation and acidification often coincide because biological consumption of oxygen produces CO 2 . Deoxygenation can have a range of detrimental effects on marine organisms and reduce the extent of marine habitats (Sections 3.3.2, 3.4.3.1; [[#Vaquer-Sunyer--2008|Vaquer-Sunyer and Duarte, 2008]] ; [[#Chu--2015|Chu and Tunnicliffe, 2015]] ).

Changes in ocean oxygen concentrations have been analysed from compilations of ''in situ'' data dating back to the 1960s ( [[#Helm--2011|Helm et al., 2011]] ; [[#Ito--2017|Ito et al., 2017]] ; [[#Schmidtko--2017|Schmidtko et al., 2017]] ). SROCC concluded that a loss of oxygen had occurred in the upper 1000 m of the ocean ( ''medium confidence'' ), with a global mean decrease of 0.5–3.3% ( ''very likely range'' ) over 1970–2010 ( [[#Bindoff--2019a|Bindoff et al., 2019a]] ). Based on new regional assessments ( [[#Queste--2018|Queste et al., 2018]] ; [[#Bronselaer--2020|Bronselaer et al., 2020]] ; [[#Cummins--2020|Cummins and Ross, 2020]] ; [[#Stramma--2020|Stramma et al., 2020]] ), WGI AR6 assesses that ocean deoxygenation has occurred in most regions of the open ocean since the mid-20th century ( ''high confidence'' ), but it is modified by climate variability on interannual and inter-decadal time scales ( ''medium confidence'' ) (WGI AR6 Sections 2.3.3.6, 5.3.3.2; [[#Canadell--2021|Canadell et al., 2021]] ; [[#Gulev--2021|Gulev et al., 2021]] ). New findings since SROCC also confirm that the volume of oxygen minimum zones (OMZs) are expanding at many locations ( ''high confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3|Section 5.3.3.2]] ; [[#Canadell--2021|Canadell et al., 2021]] ).

The most recent estimates of future oxygen loss in the subsurface ocean (100–600 m), using CMIP6 models, amount to −4.1 ± 4.2 ( ''very likely range'' ), −6.6 ± 5.7, −10.1 ± 6.7 and −11.2 ± 7.7% in 2081–2100 relative to 1995–2014 for SSP1-2.6, SSP2-4.5, SSP3-7.0 and SSP5-8.5, respectively (Figure 3.5; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Based on these CMIP6 projections, WGI AR6 concludes that the oxygen content of the subsurface ocean is projected to decline to historically unprecedented conditions over the 21st century ( ''medium confidence'' ) (WGI AR6 [[IPCC:Wg2:Chapter:Chapter-5#5.3|Section 5.3.3.2]] ; [[#Canadell--2021|Canadell et al., 2021]] ). These declines are greater (by 31–72%) than simulated by the CMIP5 models in their Representative Concentration Pathway (RCP) analogues, a ''likely'' consequence of enhanced surface warming and stratification in CMIP6 models (Figure 3.5; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). At the regional scale and for subsurface waters, projected changes are not spatially uniform, and there is ''lower agreement'' among models than they show for the global mean trend ( [[#Bopp--2013|Bopp et al., 2013]] ; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). In particular, large uncertainties remain for these future projections of ocean deoxygenation in the subsurface tropical oceans, where the major OMZs are located ( [[#Cabré--2015|Cabré et al., 2015]] ; [[#Bopp--2017|Bopp et al., 2017]] ).

<div id="3.2.3.3 " class="h3-container"></div>

<span id="changes-in-nutrient-availability"></span>
==== 3.2.3.3 Changes in Nutrient Availability ====

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The availability of nutrients in the surface ocean often limits primary productivity, with implications for marine food webs and the biological carbon pump. Nitrogen availability tends to limit phytoplankton productivity throughout most of the low-latitude ocean, whereas dissolved iron availability limits productivity in high-nutrient, low-chlorophyll regions, such as in the main upwelling region of the Southern Ocean and the Eastern Equatorial Pacific ( ''high confidence'' ) ( [[#Moore--2013|Moore et al., 2013]] ; [[#IPCC--2019b|IPCC, 2019b]] ). Phosphorus, silicon, other micronutrients such as zinc, and vitamins can also co-limit marine phytoplankton productivity in some ocean regions ( [[#Moore--2013|Moore et al., 2013]] ). Whereas some studies have shown coupling between climate variability and nutrient trends in specific regions, such as in the North Atlantic ( [[#Hátún--2016|Hátún et al., 2016]] ), North Pacific ( [[#Di%20Lorenzo--2009|Di Lorenzo et al., 2009]] ; [[#Yasunaka--2014|Yasunaka et al., 2014]] ) and tropical ( [[#Stramma--2021|Stramma and Schmidtko, 2021]] ) Oceans, very few studies have been able to detect long-term changes in ocean nutrient concentrations (but see [[#Yasunaka--2016|Yasunaka et al., 2016]] ).

Future changes in nutrient concentrations have been estimated using ESMs, with future increases in stratification generally leading to decreased nutrient levels in surface waters ( [[#IPCC--2019b|IPCC, 2019b]] ). CMIP6 models project a decline in the nitrate concentration of the upper 100 m in 2080–2099 relative to 1995–2014 of −0.46 ± 0.45 ( ''very likely range'' ), −0.60 ± 0.58, −0.80 ± 0.77 and −1.00 ± 0.78 mmol m –3 under SSP1-2.6, SSP2-4.5 and SSP5-8.5, respectively (Figure 3.5; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). These declines in nitrate concentration are greater than simulated by the CMIP5 models in their RCP analogues, a ''likely'' consequence of enhanced surface warming and stratification in CMIP6 models (Figure 3.5; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). It is concluded that the surface ocean will encounter reduced nitrate concentrations in the 21st century ( ''medium confidence'' ).

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<span id="global-synthesis-on-multiple-climate-induced-drivers"></span>
=== 3.2.4 Global Synthesis on Multiple Climate-induced Drivers ===

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In the 21st century, ocean and coastal ecosystems are projected to face conditions unprecedented over past centuries to millennia ( ''high confidence'' ) ( [[#3.2|Section 3.2]] ; WGI AR6 Chapters 4, 9; [[#Fox-Kemper--2021|Fox-Kemper et al., 2021]] ; [[#Lee--2021|Lee et al., 2021]] ), with increased temperatures ( ''virtually certain'' ) and frequency and severity of MHWs ( ''very high confidence'' ), stronger upper-ocean stratification ( ''high confidence'' ), continued rise in GMSL throughout the 21st century ( ''high confidence'' ) and increased frequency of extreme sea levels ( ''high confidence'' ), further acidification ( ''virtually certain'' ), oxygen decline ( ''high confidence'' ) and decreased surface nitrate inventories ( ''medium confidence'' ).

The rates and magnitudes of these changes largely depend on the extent of future emissions ( ''very high confidence'' ), with surface ocean warming and acidification ( ''very likely range'' ) at +3.47°C ± 1.28°C and −0.44 pH units ± 0.008 pH units in 2080–2099 (relative to 1870–1899) for SSP5-8.5 compared with +1.42°C ± 0.53°C and −0.16 pH units ± 0.003 pH units for SSP1-2.6 (Figure 3.5; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ).

<div id="3.2.4.1 " class="h3-container"></div>

<span id="compound-changes-in-the-21st-century"></span>
==== 3.2.4.1 Compound Changes in the 21st century ====

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Earth system models project distinct regional evolutions of the different CIDs over the 21st century ( ''very high confidence'' ) (Figures 3.5, 3.6, 3.7; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Tropical and subtropical oceans are characterised by projected warming and acidification, accompanied by declining nitrate concentrations in equatorial upwelling regions. The North Atlantic is characterised by a high exposure to acidification and declining nitrate concentrations. The North Pacific is characterised by high sensitivity to compound changes, with high rates of warming, acidification, deoxygenation and nutrient depletion. In contrast, the development of compound hazards is limited in the Southern Ocean, where rates of warming and nutrient depletion are lower. The Arctic Ocean is characterised by the highest rates of acidification and warming, strong nutrient depletion, and it will ''likely'' become practically sea ice free in the September mean for the first time before the year 2050 in all SSP scenarios ( ''high confidence'' ) (Figures 3.5, 3.6, 3.7; Sections 3.2.2, 3.2.3).

In general, the projected changes in climate-induced drivers are less in absolute terms in the deep-sea (mesopelagic and bathypelagic domains and deep-sea habitats) than in the surface ocean and in shallow-water habitats (e.g., kelp ecosystems, warm-water corals) ( ''very high confidence'' ) (Figures 3.6, 3.7; [[#Mora--2013|Mora et al., 2013]] ; [[#Sweetman--2017|Sweetman et al., 2017]] ). The mesopelagic domain will be nevertheless exposed to high rates of deoxygenation (Figure 3.6) and high climate velocities (Figure 3.4; [[#3.2.2.1|Section 3.2.2.1]] ), as well as impacted by the shoaling of aragonite or calcite saturation horizon ( [[#3.2.3|Section 3.2.3.2]] ). Significant differences in projected trends between the SSPs show that mitigation strategies will limit exposure of deep-sea ecosystems to potential warming, acidification and deoxygenation during the 21st century ( ''very high confidence'' ) (Figure 3.6; [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ).

<div id="_idContainer019" class="Figure"></div>

[[File:1e3c4d810aece0e691068cb80ead7195 IPCC_AR6_WGII_Figure_3_006.png]]

'''Figure 3.6 |''' '''Projected trends across open-ocean systems.''' Projected annual and global (a) average warming, (b) acidification, (c) changes in dissolved oxygen concentrations and (d) changes in nitrate (NO 3 ) concentrations for four open-ocean systems, including the epipelagic (0–200 m depth), mesopelagic (200–1000 m), bathypelagic (&gt;1000 m) domains and deep benthic waters (&gt;200 m). All projections are based on Coupled Model Intercomparison Project 6 models and for three Shared Socioeconomic Pathways (SSPs): SSP1-2.6, SSP2-4.5 and SSP5-8.5 ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Anomalies in the near-term (2020–2041), mid-term (2041–2060) and long-term (2081–2100) are all relative to 1985–2014. Error bars represent ''very likely'' ranges.

<div id="_idContainer021" class="Figure"></div>

[[File:32a5117ac077716b623928bcd121634b IPCC_AR6_WGII_Figure_3_007.png]]

'''Figure 3.7 |''' '''Projected trends across coastal-ocean ecosystems.''' Projected '''(a)''' warming, '''(b)''' acidification, '''(c)''' changes in dissolved oxygen concentrations, '''(d)''' changes in nitrate (NO 3 ) concentrations and '''(e)''' changes in summer sea ice cover fraction (September and north of 66°N for the Northern Polar Oceans, and March and south of 66°S for the Southern Polar Ocean) for five coastal-ocean ecosystems. All projected trends are for the surface ocean, except oxygen concentration changes that are computed for the subsurface ocean (100–600 m depth) for the upwelling ecosystems and the polar seas. All projections are based on Coupled Model Intercomparison Project 6 (CMIP6) models and for three Shared Socioeconomic Pathways (SSPs): SSP1-2.6, SSP2-4.5 and SSP5-8.5 ( [[#Kwiatkowski--2020|Kwiatkowski et al., 2020]] ). Anomalies in the near term (2020–2041), mid term (2041–2060) and long term (2081–2100) are all relative to 1985–2014. Error bars represent ''very likely'' ranges. Coastal seas are defined on a 1° × 1° grid when bathymetry is less than 200 m deep. Distribution of warm-water corals is from UNEP-WCMC et al. (2018). Distribution of kelp ecosystems is from [[#OBIS--2020|OBIS (2020)]] . Upwelling areas are defined according to [[#Rykaczewski--2015|Rykaczewski et al. (2015)]] .

<div id="3.2.4.2 " class="h3-container"></div>

<span id="time-of-emergence"></span>
==== 3.2.4.2 Time of Emergence ====

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Anthropogenic changes in climate-induced drivers assessed here exhibit vastly distinct times of emergence, which is the time scale over which an anthropogenic signal related to climate change is statistically detected to emerge from the background noise of natural climate for a specific region ( [[#Christensen--2007|Christensen et al., 2007]] ; [[#Hawkins--2012|Hawkins and Sutton, 2012]] ). SROCC concluded that for ocean properties, the time of emergence ranges from under a decade (e.g., surface ocean pH) to over a century (e.g., net primary production; see [[#3.4.3.3.4|Section 3.4.3.3.4]] for time of emergence of biological properties; [[#Bindoff--2019a|Bindoff et al., 2019a]] ).

The literature assessed in SROCC mainly focused on surface ocean properties and gradual mean changes. Since then, the time of emergence has also been investigated for subsurface properties, ocean extreme events and particularly vulnerable regions, such as the Arctic Ocean ( [[#Hameau--2019|Hameau et al., 2019]] ; [[#Oliver--2019|Oliver et al., 2019]] ; [[#Burger--2020|Burger et al., 2020]] ; [[#Landrum--2020|Landrum and Holland, 2020]] ; [[#Schlunegger--2020|Schlunegger et al., 2020]] ), but subsequent assessments are ''low confidence'' due to ''limited evidence'' . Below the surface, changes in temperature typically emerge from internal variability prior to changes in oxygen; however, in about a third of the global thermocline, deoxygenation emerges prior to warming ( [[#Hameau--2019|Hameau et al., 2019]] ). Permanent MHW states, defined as when SST exceeds the MHW threshold continuously over a full calendar year, will emerge during the 21st century in many parts of the surface ocean ( [[#Oliver--2019|Oliver et al., 2019]] ). Ocean acidification extremes have already emerged from background natural internal variability during the 20th century in most of the surface ocean ( [[#Burger--2020|Burger et al., 2020]] ). In the Arctic, anthropogenic sea ice changes have already emerged from the background internal variability, and anthropogenic alteration of air temperatures will emerge in the early- to mid-21st century ( [[#Landrum--2020|Landrum and Holland, 2020]] ).

<div id="3.2.4.4 " class="h3-container"></div>

<span id="perspectives-from-paleoclimatology-data"></span>
==== 3.2.4.4 Perspectives from Paleoclimatology Data ====

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Paleoclimatology observations are useful to assess multiple hazards of environmental change while excluding direct anthropogenic impacts ( [[#3.4.3.3|Section 3.4.3.3]] ). Ancient intervals of rapid climate warming that occurred between 300 and 50 million years ago (Ma) were triggered by the release of greenhouse gases ( ''high confidence'' ). The sources of greenhouse gases varied but include volcanic degassing from continental flood basalts and methane hydrates stored in marine sediments and soils ( [[#Foster--2018|Foster et al., 2018]] ). Six extreme ancient hyperthermal events are known from the last 300 Ma, when tropical SSTs reached 1.5°C–10°C warmer than pre-industrial conditions, and with substantial impacts on ancient life (Cross-Chapter Box PALEO in Chapter 1). Warming and deoxygenation in the oceans were closely associated in hyperthermal events ( ''high confidence'' ), with anoxia reaching the photic zone and abyssal depths ( [[#Kaiho--2014|Kaiho et al., 2014]] ; [[#Müller--2017|Müller et al., 2017]] ; [[#Penn--2018|Penn et al., 2018]] ; [[#Weissert--2019|Weissert, 2019]] ), whereas ocean acidification has not been demonstrated consistently ( ''medium confidence'' ) ( [[#Hönisch--2012|Hönisch et al., 2012]] ; [[#Penman--2014|Penman et al., 2014]] ; [[#Clarkson--2015|Clarkson et al., 2015]] ; [[#Harper--2020a|Harper et al., 2020a]] ; [[#Jurikova--2020|Jurikova et al., 2020]] ; [[#Müller--2020|Müller et al., 2020]] ).

Greenhouse gases also contributed substantially to shaping the longer-term climate trends over the past 50 million years, although changes in continental configuration and ocean circulation as well as planetary orbital cycles were equally important (WGI AR6 Cross-Chapter Box 2.1 in Chapter 2; [[#Westerhold--2020|Westerhold et al., 2020]] ; [[#Gulev--2021|Gulev et al., 2021]] ). There is little evidence for ocean acidification in the past 2.6 Ma ( ''low confidence'' ) ( [[#Hönisch--2012|Hönisch et al., 2012]] ), but ocean ventilation was highly sensitive to even modest warming such as observed in the past 10,000 years ( ''medium confidence'' ) ( [[#Jaccard--2012|Jaccard and Galbraith, 2012]] ; [[#Lembke-Jene--2018|Lembke-Jene et al., 2018]] ).

<div id="3.3" class="h1-container"></div>

<span id="linking-biological-responses-to-climate-induced-drivers"></span>