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

=== 5.3.1 Paleoclimate Context ===

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==== 5.3.1.1 Paleocene–Eocene Thermal Maximum ====

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The Paleocene–Eocene Thermal Maximum (PETM) was an episode of global warming exceeding pre-industrial temperatures by 4°C–8°C ( [[#McInerney--2011|McInerney and Wing, 2011]] ; [[#Dunkley%20Jones--2013|Dunkley]] [[#Jones--2013|]] [[#Jones--2013|Jones et al., 2013]] ) that occurred 55.9–55.7 Ma. The PETM involved a large pulse of geologic CO <sub>2</sub> released into the ocean–atmosphere system in 3–20 kyr ( [[#Zeebe--2016|Zeebe et al., 2016]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Kirtland%20Turner--2017|Kirtland]] [[#Turner--2017|Turner et al., 2017]] ; [[#Kirtland%20Turner--2018|Kirtland Turner, 2018]] ; [[#Gingerich--2019|Gingerich, 2019]] ; [[#5.2.1.1|Section 5.2.1.1]] ). In response, observationally constrained model simulations report an increase in atmospheric CO <sub>2</sub> concentrations ranging from about 900 ppm to &gt;2000 ppm (Chapter 2; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Cui--2018|Cui and Schubert, 2018]] ; [[#Anagnostou--2020|Anagnostou et al., 2020]] ). The PETM thus provides a test for our understanding of the ocean’s response to the increase in carbon (and heat) emissions over geologically short time scales.

A limited number of independent proxy records indicate that the PETM was associated with a surface ocean pH decline ranging from 0.15 to 0.30 units ( [[#Penman--2014|Penman et al., 2014]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Babila--2018|Babila et al., 2018]] ) ''.'' It was also accompanied by a rapid (&lt;10 ka) shallowing of the carbonate saturation horizon, resulting in the widespread dissolution of sedimentary carbonate, followed by a gradual (100 kyr) recovery ( [[#Zachos--2005|Zachos et al., 2005]] ; [[#Bralower--2018|Bralower et al., 2018]] ). The remarkable similarity among sedimentary records spanning a wide range of ecosystems suggests with ''medium confidence'' that the perturbation in the ocean carbonate saturation was global ( [[#Babila--2018|Babila et al., 2018]] ) and directly resulted from elevated atmospheric CO <sub>2</sub> levels. The degree of acidification is similar to the 0.4 pH unit decrease projected for the end of the 21st century under RCP8.5 ( [[#Gattuso--2015|Gattuso et al., 2015]] ) and is estimated to have occurred at a rate about one order of magnitude slower than the current rate of ocean acidification ( [[#Zeebe--2016|Zeebe et al., 2016]] ). There is ''low confidence'' in the inferred rates of ocean acidification inherent to the range of uncertainties affecting rates estimates based on marine sediments ( [[#5.1.2.1|Section 5.1.2.1]] ).

Recent model outputs and globally distributed geochemical data reveal with ''medium confidence'' widespread ocean deoxygenation during the PETM ( [[#Dickson--2012|Dickson et al., 2012]] , 2014; [[#Winguth--2012|Winguth et al., 2012]] ; [[#Chang--2018|Chang et al., 2018]] ; [[#Remmelzwaal--2019|Remmelzwaal et al., 2019]] ), with parts of the ocean potentially becoming drastically oxygen-depleted (anoxic; [[#Yao--2018|Yao et al., 2018]] ; [[#Clarkson--2021|Clarkson et al., 2021]] ). Deoxygenation affected the surface ocean globally (including the Arctic Ocean; [[#Sluijs--2006|Sluijs et al., 2006]] ), due to vertical and lateral expansion of OMZs ( [[#Zhou--2014|Zhou et al., 2014]] ) that resulted from warming and related changes in ocean stratification ''.'' Expansion of OMZs may have stimulated N <sub>2</sub> O production through water-column (de)nitrification ( [[#Junium--2018|Junium et al., 2018]] ). The degree to which N <sub>2</sub> O production impacted PETM warming, however, has not yet been established.

The feedbacks associated with recovery from the PETM are uncertain, yet could include drawdown associated with silicate weathering ( [[#Zachos--2005|Zachos et al., 2005]] ) and regrowth of terrestrial and marine organic carbon stocks ( [[#Bowen--2010|Bowen and Zachos, 2010]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ).

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==== 5.3.1.2 Last Deglacial Transition ====

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The Last Deglacial Transition (LDT) is the best documented climatic transition in the past associated with a substantial atmospheric CO <sub>2</sub> rise ranging from 190 to 265 ppm between 18–11 ka ( [[#Marcott--2014|Marcott et al., 2014]] ). The amplitude of the deglacial CO <sub>2</sub> rise is thus on the same order of magnitude as the increase since the Industrial Revolution.

Boron isotope ( δ <sup>11</sup> B) data suggest a 0.15–0.05 unit decrease in sea surface pH ( [[#Hönisch--2005|Hönisch and Hemming, 2005]] ; [[#Henehan--2013|Henehan et al., 2013]] ) across the LDT, an average rate of decline of about 0.002 units per century compared with the current rate of more than 0.1 units per century ( [[#Bopp--2013|Bopp et al., 2013]] ; [[#Gattuso--2015|Gattuso et al., 2015]] ). Planktonic foraminiferal shell weights decreased by 40% to 50% ( [[#Barker--2002|Barker and Elderfield, 2002]] ), and coccolith mass decreased by about 25% ( [[#Beaufort--2011|Beaufort et al., 2011]] ) across the LDT. Independent proxy reconstructions thus highlight with ''high confidence'' that pH values decreased as atmospheric CO <sub>2</sub> concentrations increased across the LDT. There is, however, ''low confidence'' in the inferred rate of ocean acidification owing to multiple sources of uncertainties affecting rates estimates based on marine sediments ( [[#5.1.2.1|Section 5.1.2.1]] ).

Geochemical and micropaleontological evidence suggest that intermediate-depth OMZs almost vanished during the Last Glacial Maximum (LGM) ( [[#Jaccard--2014|Jaccard et al., 2014]] ). However, multiple lines of evidence suggest with ''medium confidence'' that the deep (&gt;1500 m) ocean became depleted in O <sub>2</sub> (concentrations were possibly lower than 50 μmol kg <sup>–1</sup> ) globally (Jaccard and Galbraith, 2012; [[#Hoogakker--2015|Hoogakker et al., 2015]] , 2018; [[#Gottschalk--2016|Gottschalk et al., 2016]] , 2020a; [[#Anderson--2019|Anderson et al., 2019]] ) as a combined result of sluggish ventilation of the ocean subsurface ( [[#Gottschalk--2016|Gottschalk et al., 2016]] , 2020a; [[#Skinner--2017|Skinner et al., 2017]] ) and a generally more efficient marine biological carbon pump ( [[#Buchanan--2016|Buchanan et al., 2016]] ; [[#Yamamoto--2019|Yamamoto et al., 2019]] ; [[#Galbraith--2020|Galbraith and Skinner, 2020]] ).

During the LDT, deep ocean ventilation increased as Antarctic Bottom Water (AABW) ( [[#Skinner--2010|Skinner et al., 2010]] ; [[#Gottschalk--2016|Gottschalk et al., 2016]] ; [[#Jaccard--2016|Jaccard et al., 2016]] ) and subsequently the Atlantic meridional overturning circulation ( [[#McManus--2004|McManus et al., 2004]] ; [[#Lippold--2016|Lippold et al., 2016]] ) resumed, transferring previously sequestered remineralized carbon from the ocean interior to the upper ocean, and eventually the atmosphere ( [[#Skinner--2010|Skinner et al., 2010]] ; [[#Galbraith--2015|Galbraith and Jaccard, 2015]] ; [[#Gottschalk--2016|Gottschalk et al., 2016]] ; [[#Ronge--2016|Ronge et al., 2016]] , 2020; [[#Sikes--2016|Sikes et al., 2016]] ; [[#Rae--2018|Rae et al., 2018]] ), contributing to the deglacial CO <sub>2</sub> rise. Intermediate depths lost oxygen as a result of sluggish ventilation and increasing temperatures (decreasing saturation). As the world emerged from the last Glacial period, OMZs underwent a large volumetric increase at the beginning of the Bølling-Allerød (B/A), a northern-hemisphere wide warming event, 14.7 ka ( [[#Jaccard--2012|Jaccard and Galbraith, 2012]] ; [[#Praetorius--2015|Praetorius et al., 2015]] ) with deleterious consequences for benthic ecosystems (e.g., [[#Moffitt--2015|Moffitt et al., 2015]] ). These observations indicate with ''high confidence'' that the rate of warming, affecting the solubility of oxygen and upper water column stratification, coupled with changes in subsurface ocean ventilation, impose a direct control on the degree of ocean deoxygenation, implying a high sensitivity of ocean oxygen loss to warming. The expansion of OMZs contributed to a widespread increase in water column (de)nitrification (Galbraith and Kienast, 2013), which contributed substantially to enhanced marine N <sub>2</sub> O emissions. Nitrogen stable isotope measurements on N <sub>2</sub> O extracted from ice cores suggest that approximately one-third (of the order of 0.7 ± 0.3TgN yr <sup>–1</sup> ) of thedeglacial increase in N <sub>2</sub> O emissions relates to oceanic sources (Schilt et al., 2014; H. [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ).

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