Editing IPCC:AR6/WGI/Chapter-9 (section)

=== 9.6.2 Paleo Context of Global and Regional Sea Level Change ===

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As SROCC ( [[#Oppenheimer--2019|Oppenheimer et al., 2019]] ) noted, paleo sea level records provide information on past ice-sheet changes, and process-based ice-sheet models of past warm periods inform equilibrium responses. However, given uncertainties in paleo sea level and polar paleoclimate, and limited temporal resolution of paleo sea level records, there is ''low confidence'' in the utility of paleo sea level records for quantitatively informing near-term GMSL change. Nonetheless, the paleorecord does contextualize sea level and can test projection models (see also FAQ 1.3). Proxy constraints on GMSL and global ice volume are assessed in Sections 2.3.2.4. and 2.3.3.3 (see also FAQ 9.1). This section updates prior assessments of drivers of past GMSL changes and climatically coherent areas of relative sea level (RSL) variability. GMSL changes are framed in terms of global mean surface temperature (GMST) but noting that amplified high-latitude warming is a robust equilibrium response to elevated CO <sub>2</sub> ( [[#Masson-Delmotte--2013|Masson-Delmotte et al., 2013]] ): polar air temperatures during past warm periods were up to twice the GMST changes shown in Table 9.6. The SROCC assessment that past multi-metre sea level changes have resulted from significant ice-sheet changes beyond those presently observed is confirmed ( ''very high confidence'' ).

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'''Table 9.6''' '''|''' '''Reference ranges of age, global mean surface temperature, atmospheric carbon dioxide (CO''' <sub>2</sub> ''') concentration, and global mean sea level (GMSL) for the paleo periods discussed in this chapter.'''

{| class="wikitable"
|-
| '''Paleo Period'''

| '''Years'''

'''Cross-Chapter Box 2.1'''

| '''GMST relative to 1850–1900'''

'''Section 2.3.1.1'''

| '''CO''' <sub>2</sub>

'''Sections 2.2.3.1 and 2.2.3.2'''

| '''Global Mean Sea Level (GMSL)'''

'''Section 2.3.3.3'''

|-
| '''Early Eocene Climatic Optimum (EECO)'''

| 53–49 Ma

| +10°C to +18°C

| 1150 to 2500 ppm

| +70 to +76 m

|-
| '''Mid-Pliocene Warm Period (MPWP)'''

| 3.3–3.0 Ma

| +2.5°C to +4°C

| 360 to 420 ppm

| +5 to +25 m

|-
| '''Marine Isotope Stage (MIS) 11'''

| about 424–395 ka

| 0.5°C ± 1.6°C <sup>a</sup>

| 265 to 286 ppm

| +6 to +13 m

|-
| '''Last Interglacial (LIG)'''

| about 129–116 ka

| +0.5°C to +1.5°C

| 266 to 282 ppm

| +5 to +10 m

|-
| '''Last Glacial Maximum (LGM)'''

| 21–19 ka

| –5°C to –7°C

| 188 to 194 ppm

| –125 to –134 m

|-
| '''Last Deglacial Transition'''

| 18–11 ka

| n/a

| 193 to 271 ppm

| –120 to –50 m

|-
| '''Early Holocene'''

| 11.65–6.5 ka

| n/a

| 250 to 268 ppm

| –50 to –3.5 m

|-
| '''Mid-Holocene'''

| 6.5–5.5 ka

| +0.2°C to +1.0°C

| 260 to 268 ppm

| –3.5 to +0.5 m

|-
| '''Last Millennium'''

| 850–1850 CE

| –0.14°C to +0.24°C

| 278 to 285 ppm

| –0.05 to +0.03 m

|}

<sup>a</sup> Based on one study ( [[#Irvalı--2020|Irvalı et al., 2020]] ) relative to SST values around year 2000.

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<span id="mid-pliocene-warm-period"></span>
==== 9.6.2.1 Mid-Pliocene Warm Period ====

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During the mid-Pliocene Warm Period (MPWP), GMST was 2.5°C–4°C warmer than 1850–1900 ( ''medium confidence'' ) and GMSL was between 5 and 25 m higher than today ( ''medium confidence'' ) (Table 9.6 and [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] ). The AR5 ( [[#Masson-Delmotte--2013|Masson-Delmotte et al., 2013]] ) concluded that ice-sheet models consistently produce near-complete deglaciation of the Greenland and West Antarctic ice sheets, and multi-meter loss of the East Antarctic Ice Sheet (EAIS) in response to MPWP climate conditions. Studies since AR5 have yielded a consistent but broader range, due in part to larger ensembles exploring more parameters ( [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#Yan--2016|Yan et al., 2016]] ; [[#DeConto--2021|DeConto et al., 2021]] ). Partly on the basis of these studies, SROCC proposed a ‘plausible’ upper bound on GMSL of 25 m ( ''low confidence'' ) with evidence suggesting an Antarctic contribution of anywhere between 5.4 and 17.8 m.

The MPWP climate had substantial polar amplification, up to 8°C above pre-industrial levels in Arctic Russia ( [[IPCC:Wg1:Chapter:Chapter-7#7.4.4.1|Section 7.4.4.1]] ; [[#Fischer--2018|Fischer et al., 2018]] ). Ice-sheet model simulations indicate that Northern Hemisphere glaciation was limited to high-elevation regions in eastern and southern Greenland ( ''medium confidence'' ) (Figure 9.17; [[#De%20Schepper--2014|De Schepper et al., 2014]] ; [[#Yan--2014|Yan et al., 2014]] ; [[#Koenig--2015|Koenig et al., 2015]] ; [[#Dowsett--2016|Dowsett et al., 2016]] ; [[#Berends--2019|Berends et al., 2019]] ) with Northern Hemisphere glaciation only becoming more widespread from the (cooler) late Pliocene ( [[#Bachem--2017|Bachem et al., 2017]] ; [[#Blake-Mizen--2019|Blake-Mizen et al., 2019]] ; [[#Knutz--2019|Knutz et al., 2019]] ; [[#Sánchez-Montes--2020|Sánchez-Montes et al., 2020]] ). Southern Hemisphere glaciation was characterized by an Antarctic Ice Sheet (AIS) reduced in volume from the present ( ''medium confidence'' ) (Figure 9.18; [[#Dowsett--2016|Dowsett et al., 2016]] ; [[#Berends--2019|Berends et al., 2019]] ; [[#Grant--2019|Grant et al., 2019]] ; [[#Miller--2020|Miller et al., 2020]] ) with mountain ice fields in the Andes of South America ( [[#De%20Schepper--2014|De Schepper et al., 2014]] ). Ice-sheet models are inconsistent in the magnitude of the sea level contribution from Antarctica ( [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#Yan--2016|Yan et al., 2016]] ; [[#Golledge--2017b|Golledge et al., 2017b]] ; [[#Berends--2019|Berends et al., 2019]] ; [[#DeConto--2021|DeConto et al., 2021]] ) but near-field sedimentological reconstructions support precessionally modulated and eccentricity-paced multi-metre sea level contributions from the Wilkes Subglacial Basin over 3–5 kyr ( [[#Patterson--2014|Patterson et al., 2014]] ; [[#Bertram--2018|Bertram et al., 2018]] ). Insummary, under a past warming level of around 2.5°C–4°C, ice sheets in both hemispheres were reduced in extent compared to present ( ''high confidence'' ). Proxy-based evidence ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] ) combined with numerical modelling indicates that, on millennial time scales, the GMSL contribution arising from ice sheets was &gt;5 m ( ''high confidence'' ) or &gt;10 m ( ''medium confidence'' ) (Figures 9.17 and 9.18; [[#Moucha--2017|Moucha and Ruetenik, 2017]] ; [[#Berends--2019|Berends et al., 2019]] ; [[#Dumitru--2019|Dumitru et al., 2019]] ).

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==== 9.6.2.2 Marine Isotope Stage 11 ====

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The SROCC ( [[#Meredith--2019|Meredith et al., 2019]] ) noted that Greenland may have been ice-free for extensive periods during Pleistocene interglaciations, implying a high sensitivity of the Greenland Ice Sheet to warming levels close to present day. The AR5 ( [[#Church--2013b|Church et al., 2013b]] ) assigned ''medium confidence'' to a Marine Isotope Stage 11 (MIS 11) GMSL of 6–15 m above present, requiring a loss of much of the Greenland and West Antarctic ice sheets, and a possible contribution from East Antarctica. High-resolution multi-proxy sea surface temperature reconstructions and climate model simulations concur that MIS 11 was an extremely long interglacial that exhibited positive annual at 0.5°C ± 1.6 °C ( [[#Irvalı--2020|Irvalı et al., 2020]] ) and summer at 2.1°C–3.4 °C ( [[#Robinson--2017|Robinson et al., 2017]] ) temperature anomalies ( [[#de%20Wet--2016|de Wet et al., 2016]] ). The GMSL was 6–13 m above present ( ''medium confidence'' ) ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] ). The Greenland Ice Sheet lost 4.5–6 m ( [[#Reyes--2014|Reyes et al., 2014]] ) or about 6.1 m (3.9–7 m, 95% confidence) sea level equivalent (SLE) by about 7 kyr after peak summer warmth ( [[#Robinson--2017|Robinson et al., 2017]] ), with marine-based ice from AIS ( [[#Blackburn--2020|Blackburn et al., 2020]] ) contributing 6.4–8.8 m SLE at this time ( [[#Mas%20e%20Braga--2021|Mas e Braga et al., 2021]] ). Agreement between GMSL and ice-sheet reconstructions gives ''high confidence'' in identifying a high sensitivity of both ice sheets to the protracted duration of thermal forcing, even at low warming levels ( [[#Reyes--2014|Reyes et al., 2014]] ; [[#Robinson--2017|Robinson et al., 2017]] ; [[#Irvalı--2020|Irvalı et al., 2020]] ; [[#Mas%20e%20Braga--2021|Mas e Braga et al., 2021]] ). Modelled mean mass loss rates for the Greenland Ice Sheet of 0.4 m kyr <sup>–1</sup> during MIS 11 ( [[#Robinson--2017|Robinson et al., 2017]] ) are indistinguishable from recent mass loss rates averaged over 1992–2018 ( [[#9.4.1.1|Section 9.4.1.1]] ). In summary, geological reconstructions and numerical simulations consistently show that past warming levels of &lt;2°C (GMST) are sufficient to trigger multi-metre mass loss from both the Greenland and Antarctic ice sheets if maintained for millennia ( ''high confidence'' ), in agreement with SROCC findings for comparable warming levels during MIS 5e, the Last Interglacial.

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<span id="last-interglacial"></span>
==== 9.6.2.3 Last Interglacial ====

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The AR5 found that the Last Interglacial (LIG) GMSL was &gt;5 m ( ''very high confidence'' ) but &lt;10 m ( ''high confidence'' ). Their best estimate of 6 m was based on two studies ( [[#Kopp--2009|Kopp et al., 2009]] ; [[#Dutton--2012|Dutton and Lambeck, 2012]] ). The SROCC concluded that, during the LIG, Greenland’s contribution to the GMSL highstand (the highest sea levels during the LIG) of 6–9 m increased gradually, whereas the Antarctic contribution occurred early, from about 129 ka. Due to widely varying reconstructions from model studies (Greenland) and the paucity of direct evidence of ice-sheet change (Antarctic), the magnitude of sea level contributions from both ice sheets was assigned ''low confidence.''

Since AR5, information has improved about the LIG, when GMST was about 0.5°C–1.5°C above 1850–1900 ( ''medium confidence'' ) ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1.1|Section 2.3.1.1]] ). The LIG had higher summer insolation than present and polar amplified sea surface and surface air temperatures that reached &gt;1°C–4°C and &gt;3°C–11 °C in the Arctic respectively ( [[#Landais--2016|Landais et al., 2016]] ; [[#Capron--2017|Capron et al., 2017]] ; [[#Fischer--2018|Fischer et al., 2018]] ). Mean annual and maximum summer ocean temperatures peaked early (129–125 ka) in the interglacial period, reaching 1.1 ± 0.3 °C above the modern global mean ( [[#Shackleton--2020|Shackleton et al., 2020]] ) with summer anomalies of 2.5°C–3.5 °C in the Southern Ocean ( [[#Bianchi--2002|Bianchi and Gersonde, 2002]] ) and spatially variable timing ( [[#Chadwick--2020|Chadwick et al., 2020]] ). It is ''virtually certain'' that GMSL was higher than today, ''likely'' by 5–10 m ( ''medium confidence'' ) ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] ). Global mean thermal expansion peaked at about 0.9 ± 0.3 m early in the LIG (about 129 ka), declining to modern levels by about 127 ka ( [[#Shackleton--2020|Shackleton et al., 2020]] ). With no more than 0.3 ± 0.1 m of GMSL rise from glaciers ( [[#9.5.1|Section 9.5.1]] ), at most 1.0 ± 0.3 m of the GMSL rise originated from sources other than the polar ice sheets.

Recent LIG ice-sheet simulations agree that peak loss from the Greenland Ice Sheet occurred late (125–120 ka; [[#Goelzer--2016|Goelzer et al., 2016]] ; [[#Tabone--2018|Tabone et al., 2018]] ; [[#Plach--2019|Plach et al., 2019]] ) when Northern Hemisphere insolation was greater than at present ( ''medium confidence'' ) ( [[#Capron--2017|Capron et al., 2017]] ), consistent with inferences from marine sediment records ( [[#Hatfield--2016|Hatfield et al., 2016]] ; [[#Irvalı--2020|Irvalı et al., 2020]] ) and far-field GMSL indicators ( [[#Rohling--2019|Rohling et al., 2019]] ). Best estimates of the GMSL contribution from Greenland (Figure 9.17) differ between models: ≤1 m ( [[#Albrecht--2020|Albrecht et al., 2020]] ; [[#Clark--2020|Clark et al., 2020]] ), 1–2 m ( [[#Calov--2015|Calov et al., 2015]] ; [[#Goelzer--2016|Goelzer et al., 2016]] ; [[#Bradley--2018|Bradley et al., 2018]] ), up to 3 m ( [[#Tabone--2018|Tabone et al., 2018]] ; [[#Plach--2019|Plach et al., 2019]] ), and &gt;5 m ( [[#Yau--2016|Yau et al., 2016]] ). There is ''high confidence'' that the response time of the Greenland Ice Sheet to LIG warming was multi-millennial, and ''high confidence'' that it contributed to LIG GMSL change, but ''low agreement'' in the contribution magnitude.

Far-field GMSL records suggest that the AIS contributed to LIG sea level from 129.5–125 ka (Figure 9.18) but direct evidence is sparse. Thinning of part of the WAIS is interpreted from a 130–80 ka hiatus in the Patriot Hills horizontal ice core record ( [[#Turney--2020|Turney et al., 2020]] ). Marine sediment records suggest a dynamic response of the Wilkes Subglacial Basin (WSB) of the EAIS during this period, indicating a response time scale of 1000–2500 yr ( [[#Wilson--2018|Wilson et al., 2018]] ), consistent with modelling studies ( [[#Mengel--2014|Mengel and Levermann, 2014]] ; [[#Golledge--2017b|Golledge et al., 2017b]] ; [[#Sutter--2020|Sutter et al., 2020]] ). Isotopic changes in the Talos Dome ice core are inconsistent with local surface lowering, limiting retreat to 0.4–0.8 m SLE from this sector ( [[#Sutter--2020|Sutter et al., 2020]] ). Ice-sheet models forced with unmodified atmosphere–ocean models ( [[#Goelzer--2016|Goelzer et al., 2016]] ; [[#Clark--2020|Clark et al., 2020]] ) simulate 3–4.4 m SLE mass loss, primarily from the WAIS, with no retreat in WSB (e.g., Figure 9.18). Models forced with proxy-based or ad hoc LIG ocean temperature anomalies ( [[#DeConto--2016|DeConto and Pollard, 2016]] ; [[#Sutter--2016|Sutter et al., 2016]] ) indicate collapse of West Antarctica under 2°C–3°C ocean forcing yielding 3–7.5 m sea level contribution, but modest or no retreat in the WSB. Based on ''limited evidence'' and ''limited agreement'' between models, there is ''low confidence'' in both the magnitude and timing of LIG mass loss from the AIS.

In summary, paleo-environmental and modelling studies indicate that, under past warming of the level achieved during the LIG (ca. 0.5°C–1.5°C), it is ''likely'' that both the Greenland and Antarctic ice sheets responded dynamically over multiple millennia ( ''high confidence'' ).

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==== 9.6.2.4 Last Glacial Maximum ====

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At the Last Glacial Maximum (LGM) geological proxies and GIA models indicate that GMSL was 125–134 m below present ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] and Figures 9.17 and 9.18). New studies have not changed AR5’s conclusions regarding the size or timing of the LGM and last glacial termination, but have further examined the LGM sea level budget. Based on a synthesis of multiple prior studies, ( [[#Simms--2019|Simms et al., 2019]] ) estimated central 67% probability contributions to the LGM lowstand (i.e., lowest levels during the LGM) of 76 ± 7 m from the North American Laurentide Ice Sheet, 18 ± 5 m from the Eurasian Ice Sheet, 10 ± 2 m from Antarctica, 4 ± 1 m from Greenland, 5.5 ± 0.5 m from glaciers, and 2.4 ± 0.3 m due to an increase in ocean density. Of the residual, up to about 1.4 m may be ascribed to groundwater, leaving a shortfall of 16 ± 10 m yet to be allocated among land ice reservoirs or lakes.

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==== 9.6.2.5 Last Deglacial Transition: Meltwater pulse 1A ====

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During Meltwater pulse 1A (MWP-1A), GMSL ''very likely'' ( ''medium confidence'' ) rose by 8–15 m ( [[#Liu--2016|Liu et al., 2016]] ). Consistent with AR5, the drivers of this rapid rise remain ambiguous. The spatial patterns of RSL change over this interval are inadequately observed to constrain the relative contributions of the North American and Antarctic ice sheets ( [[#Liu--2016|Liu et al., 2016]] ). Modelling studies of the North American Ice Sheet permit a 3–6 m ( [[#Gregoire--2016|Gregoire et al., 2016]] ) or 6–9 m contribution over the duration of MWP-1A ( [[#Tarasov--2012|Tarasov et al., 2012]] ). Sedimentological evidence ( [[#Weber--2014|Weber et al., 2014]] ; [[#Bart--2018|Bart et al., 2018]] ) provides near-field evidence for an Antarctic contribution, consistent with modelling studies ( [[#Golledge--2014|Golledge et al., 2014]] ; [[#Stuhne--2015|Stuhne and Peltier, 2015]] ), but does not constrain the magnitude of the contribution. A recent statistical analysis of Norwegian Sea and Arctic Ocean sediments suggests a 3–7 m contribution from the Eurasian Ice Sheet ( [[#Brendryen--2020|Brendryen et al., 2020]] ), a possibility not considered in AR5 or the meta-analysis of [[#Liu--2016|Liu et al. (2016)]] . In summary, MWP-1A appears to have been driven by a combination of melt in North America ( ''high confidence'' ), Eurasia ( ''low confidence'' ), and Antarctica ( ''low confidence'' ), but the budget is not closed.

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==== 9.6.2.6 Holocene ====

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Around half (50–60 m) of the GMSL rise since the LGM occurred during the early Holocene at a sustained rate of about 15 m kyr <sup>–1</sup> from around 11.4–8.2 ka ( [[#Lambeck--2014|Lambeck et al., 2014]] ), possibly punctuated by abrupt meltwater pulses ( [[#Smith--2011|Smith et al., 2011]] ; [[#Carlson--2012|Carlson and Clark, 2012]] ; [[#Törnqvist--2012|Törnqvist and Hijma, 2012]] ; [[#Harrison--2019|Harrison et al., 2019]] ). An abrupt (about 1.1 m) sea level rise around 8.2 ka was associated with drainage of the pro-glacial Agassiz and Ojibway lakes, attributed to accelerated melt from collapsing Laurentide Ice Sheet ice saddles ( [[#Matero--2017|Matero et al., 2017]] ). The Laurentide Ice Sheet provided the greatest contribution (27 m) to early Holocene GMSL ( [[#Peltier--2015|Peltier et al., 2015]] ; [[#Roy--2017|Roy and Peltier, 2017]] ), the Scandinavian Ice Sheet contributed about 2 m from the beginning of the Holocene until its demise by around 10.5 ka, ( [[#Cuzzone--2016|Cuzzone et al., 2016]] ), while the Barents Sea Ice Sheet contributed a small but unknown amount ( [[#Patton--2015|Patton et al., 2015]] , 2017; [[#Auriac--2016|Auriac et al., 2016]] ). The Greenland Ice Sheet contributed about 4 m, consistent with ice thinning rates inferred from the Camp Century ice core ( [[#Lecavalier--2017|Lecavalier et al., 2017]] ; [[#McFarlin--2018|McFarlin et al., 2018]] ). Recent estimates of Antarctic contributions during the early Holocene vary considerably from about 1.2 m to 8.5 m ( [[#Whitehouse--2012|Whitehouse et al., 2012]] ; [[#Ivins--2013|Ivins et al., 2013]] ; [[#Argus--2014|Argus et al., 2014]] ; [[#Briggs--2014|Briggs et al., 2014]] ; [[#Golledge--2014|Golledge et al., 2014]] ; [[#Pollard--2016|Pollard et al., 2016]] ; [[#Roy--2017|Roy and Peltier, 2017]] ; [[#Albrecht--2020|Albrecht et al., 2020]] ). In summary, the early Holocene was characterized by steadily rising GMSL as global ice sheets continued to retreat from their LGM extents. This steady rise was punctuated by abrupt pulses during episodes of rapid meltwater discharge.

In the middle Holocene, GMST peaked at 0.2°C–1.0°C higher than 1850–1900 temperature between 7 and 6 ka ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.1.1.2|Section 2.3.1.1.2]] ). GMSL rise slowed coincidently with final melting of the Laurentide ice sheet by 6.7 ± 0.4 ka ( [[#Ullman--2016|Ullman et al., 2016]] ), after which only Greenland and Antarctic ice sheets could have contributed significantly. At 6 ka, GMSL was –3.5 to +0.5 m ( ''medium confidence'' ) ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] ). Simulations of the Holocene Thermal Maximum give a Greenland Ice Sheet broadly consistent with geological reconstructions so, despite uncertainties regarding the timing of minimum ice-sheet volume and extent, there is ''medium confidence'' that minima were reached at different times in different areas during the period 8–3 ka BP ( [[#Larsen--2015|Larsen et al., 2015]] ; [[#Young--2015|Young and Briner, 2015]] ; [[#Briner--2016|Briner et al., 2016]] ). Geochronological and numerical modelling studies indicate that it is ''likely'' ( ''medium confidence'' ) that the period of smaller-than-present ice extent in all sectors of Greenland persisted for at least 2000 to 3000 years ( [[#Larsen--2015|Larsen et al., 2015]] ; [[#Young--2015|Young and Briner, 2015]] ; [[#Briner--2016|Briner et al., 2016]] ; [[#Nielsen--2018|Nielsen et al., 2018]] ). Based on ice-sheet modelling and carbon-14 ( <sup>14</sup> C) dating ( [[#Kingslake--2018|Kingslake et al., 2018]] ) suggested that West Antarctic grounding lines retreated prior to around 10 ka BP, followed by a readvance. Other studies from the same region conclude that retreat was fastest from 9–8 ka BP ( [[#Spector--2017|Spector et al., 2017]] ), or from 7.5–4.8 ka BP ( [[#Venturelli--2020|Venturelli et al., 2020]] ). Marine geological evidence indicates open marine conditions east of Ross Island by 8.6 ± 0.2 ka BP ( [[#McKay--2016|McKay et al., 2016]] ). In the western Weddell Sea, [[#Johnson--2019|Johnson et al. (2019)]] reported rapid glacier thinning from 7.5–6 ka BP. [[#Hein--2016|Hein et al. (2016)]] concluded that the fastest thinning further south took place from 6.5–3.5 ka BP, potentially contributing 1.4–2 m to GMSL. Geophysical data indicate stabilization or readvance in this area around 6 ± 2 ka BP ( [[#Wearing--2019|Wearing and Kingslake, 2019]] ). In coastal Dronning Maud Land (East Antarctica) rapid thinning occurred 9–5 ka BP ( [[#Kawamata--2020|Kawamata et al., 2020]] ), whereas glaciers in the Northern Antarctic Peninsula receded during the period 11–8 ka BP and readvanced to their maximal extents by 7–4 ka BP ( [[#Kaplan--2020|Kaplan et al., 2020]] ). In summary, higher-than-pre-industrial GMST during the mid-Holocene coincided with recession of the Greenland Ice Sheet to a smaller-than-present extent ( ''high confidence'' ). Multiple lines of evidence give ''high confidence'' that thinning or retreat in parts of Antarctica during the Holocene took place at different times in different places. However, limited data means there is only ''low confidence'' in whether or not the ice sheet as a whole was smaller than present during the mid-Holocene.

In summary, both proxies and model simulations indicate that GMSL changes during the early to mid-Holocene were the result of episodic pulses, due to drainage of meltwater lakes, superimposed on a trend of steady rise due to continued ice-sheet retreat ( ''high confidence'' ).

The combination of tide gauge observations and geological reconstructions indicates that a sustained increase of GMSL began between 1820–1860 and led to a 20th-century GMSL rise that was ''very likely'' ( ''high confidence'' ) faster than in any preceding century in the last 3000 years ( [[IPCC:Wg1:Chapter:Chapter-2#2.3.3.3|Section 2.3.3.3]] ). At a regional level, tide gauge and geological data from the North Atlantic and Australasia show inflections in RSL trends between 1895–1935, with an increase of 0.8 to 2.5 mm yr <sup>–1</sup> across the inflection ( [[#Gehrels--2013|Gehrels and Woodworth, 2013]] ). A statistical meta-analysis of globally distributed geological and tide gauge data ( [[#Kopp--2016|Kopp et al., 2016]] ) found that, in all 20 examined regions with geological records stretching back at least 2000 years, the rate of RSL rise in the 20th century was greater than the local average over 0–1700 CE. In four of the 20 regions, all in the North Atlantic (Connecticut, New Jersey, North Carolina, and Iceland), the 19th century rate was also greater than the 0–1700 CE average (90% confidence interval). In summary, rates of RSL rise exceeding the pre-industrial background rate of rise are apparent in parts of the North Atlantic in the 19th century ( ''medium confidence'' ), and in most of the world in the 20th century ( ''high confidence'' ).

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<span id="future-sea-level-changes"></span>