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

==== 5.1.2.2 Glacial–Interglacial Greenhouse Gas Records ====

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The Antarctic ice core record covering the past 800 kyr provides an important archive to explore the carbon–climate feedbacks prior to anthropogenic perturbations ( [[#Brovkin--2016|Brovkin et al., 2016]] ). Polar ice cores represent the only climatic archive from which past GHG concentrations can be directly measured. Major GHGs, CH <sub>4</sub> , N <sub>2</sub> O and CO <sub>2</sub> generally co-vary on orbital time scales ( [[#Loulergue--2008|Loulergue et al., 2008]] ; [[#Lüthi--2008|Lüthi et al., 2008]] ; [[#Schilt--2010b|Schilt et al., 2010b]] ; Chapter 2), with consistently higher atmospheric concentrations during warm intervals of the past, pointing to a strong sensitivity to climate (Figure 5.4). Modelling work suggests that the carbon cycle contributed to globalise and amplify changes in orbital forcing, which are pacing glacial–interglacial climate oscillations ( [[#Ganopolski--2017|Ganopolski and Brovkin, 2017]] ), with ocean biogeochemistry and physics, terrestrial vegetation, peatland, permafrost and exchanges with the lithosphere including chemical weathering, volcanic activity, sediment burial and marine calcium carbonate compensation all playing a role in modulating the concentration of atmospheric GHGs.

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'''Figure 5.4 |''' '''Atmospheric concentrations of CO''' <sub>2</sub> ''', CH''' <sub>4</sub> '''and N''' <sub>2</sub> '''O in air bubbles and clathrate crystals in ice cores (800,000 BCE to 1990 CE).''' Note the variable x-axis range and tick mark intervals for the three columns. Ice core data is over-plotted by atmospheric observations from 1958 to present for CO <sub>2</sub> , from 1984 for CH <sub>4</sub> and from 1994 for N <sub>2</sub> O. The time-integrated, millennial-scale linear growth rates for different time periods (800,000–0 BCE, 0–1900 CE and 1900–2017 CE) are given in each panel. For the BCE period, mean rise and fall rates are calculated for the individual slopes between the peaks (interglacials) and troughs (glacial periods), which are given in the panels in left column. The data for BCE period are used from the Vostok, EPICA, Dome C and WAIS ice cores ( [[#Petit--1999|Petit et al., 1999]] ; [[#Monnin--2001|Monnin, 2001]] ; [[#Pépin--2001|Pépin et al., 2001]] ; [[#Raynaud--2005|Raynaud et al., 2005]] ; [[#Siegenthaler--2005|Siegenthaler et al., 2005]] ; [[#Loulergue--2008|Loulergue et al., 2008]] ; [[#Lüthi--2008|Lüthi et al., 2008]] ; [[#Schilt--2010a|Schilt et al., 2010a]] ). The data after 0–yr CE are taken mainly from Law Dome ice core analysis ( [[#MacFarling%20Meure--2006|MacFarling Meure et al., 2006]] ). The surface observations for all species are taken from NOAA cooperative research network ( [[#Dlugokencky--2019|Dlugokencky and Tans, 2019]] ), where ALT, MLO and SPO stand for Alert (Canada), Mauna Loa Observatory, and South Pole Observatory, respectively. BCE = before current era, CE = current era. Further details on data sources and processing are available in the chapter data table (Table 5.SM.6).

Since AR5, the number of ice core records and the temporal resolution of their data for the last 800 kyr have improved, in particular for the last 60 kyr. Additionally, the advent of isotopic measurements on GHGs extracted from air trapped in ice, allows for more robust source apportionments and inventory assessments. Therefore, the ensuing discussion focuses on these two specific aspects.

Major pre-industrial sources of CH <sub>4</sub> comprise wetlands (including subglacial environments) and biomass burning ( [[#Bock--2010|Bock et al., 2010]] , [[#Bock--2017|2017]] ; [[#Lamarche-Gagnon--2019|Lamarche-Gagnon et al., 2019]] ; [[#Kleinen--2020|Kleinen et al., 2020]] ). Pre-industrial atmospheric N <sub>2</sub> O concentrations were regulated by microbial production in marine and terrestrial environments and by photochemical removal in the stratosphere (Schilt et al., 2014; [[#Battaglia--2018b|Battaglia and Joos, 2018b]] ; [[#Fischer--2019|H. Fischer et al., 2019]] ). Pre-industrial atmospheric CO <sub>2</sub> concentrations were largely regulated by exchange with exogenic terrestrial and ocean carbon reservoirs. The imbalance between geological sources and sinks in the ocean–atmosphere–land biosphere system additionally plays an important role in modulating the air–sea partitioning of the active carbon inventory on multi-millennial time scales ( [[#Cartapanis--2018|Cartapanis et al., 2018]] ).

Model-based estimates indicate that wetland CH <sub>4</sub> emissions were reduced by 24–40% during the Last Glacial Maximum (LGM) when compared to pre-industrial, while CH <sub>4</sub> emissions related to biomass burning (wildfires) decreased by 35–75% ( [[#Valdes--2005|Valdes et al., 2005]] ; [[#Hopcroft--2017|Hopcroft et al., 2017]] ; [[#Kleinen--2020|Kleinen et al., 2020]] ). N <sub>2</sub> O emissions decreased by about 30% during the LGM based on data-constrained model estimates ( [[#Schilt--2014|Schilt et al., 2014]] ; [[#Fischer--2019|H. Fischer et al., 2019]] ) owing to a combination of a weaker hydrological cycle and a generally better ventilated intermediate depth ocean relative to present, reducing (de)nitrification processes ( [[#Galbraith--2013|Galbraith et al., 2013]] ; [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ).

During past ice ages, generally colder and drier climate conditions contributed to a substantial decline of the land biosphere carbon inventory, in particular in boreal peatlands (–300 PgC; [[#Treat--2019|Treat et al., 2019]] ). Estimates assessing the glacial decrease in the global terrestrial biosphere carbon stock vary between –300 and –600 PgC ( [[#Ciais--2012|Ciais et al., 2012]] ; [[#Peterson--2014|Peterson et al., 2014]] ; Menviel et al., 2017; [[#Kleinen--2020|Kleinen et al., 2020]] ), possibly –850 PgC when accounting for ocean-sediment interactions and burial ( [[#Jeltsch-Thömmes--2019|Jeltsch-Thömmes et al., 2019]] ), a considerable contraction when compared to the modern land biosphere stock. The large range of estimates reflects a yet limited understanding of how carbon cycle dynamics were altered by glacially perturbed nutrient fluxes and soil dynamics, as well as largely exposed shelf areas in the tropics as a result of lowered sea level. Recent estimates suggest deep-sea CO <sub>2</sub> storage during the last ice age exceeded modern values by as much as 750 – 950 PgC ( [[#Skinner--2015|Skinner et al., 2015]] , [[#Skinner--2017|2017]] ; [[#Buchanan--2016|Buchanan et al., 2016]] ; [[#Anderson--2019|Anderson et al., 2019]] ; [[#Gottschalk--2020b|Gottschalk et al., 2020b]] ). A combination of increased CO <sub>2</sub> solubility associated with 2–3°C lower mean oceanic temperatures ( [[#Bereiter--2018|Bereiter et al., 2018]] ), increased the oceanic residence time of CO <sub>2</sub> ( [[#Skinner--2017|Skinner et al., 2017]] ), altered oceanic alkalinity ( [[#Yu--2010|Yu et al., 2010]] ; [[#Cartapanis--2018|Cartapanis et al., 2018]] ). A generally more efficient marine biological carbon pump (BCP; [[#Galbraith--2015|Galbraith and Jaccard, 2015]] ; [[#Yu--2019|]] [[#Yu--2019|Yu et al., 2019]] ; [[#Galbraith--2020|Galbraith and Skinner, 2020]] ) enhanced the partition CO <sub>2</sub> into the ocean interior, (although the relative contribution of each mechanism remains a matter of debate). Recent observationally constrained ESM results highlight that air–sea disequilibrium amplifies the effect of cooling and iron fertilization on glacial carbon storage ( [[#Khatiwala--2019|Khatiwala et al., 2019]] ).

Ice core observations combined with model-based estimates thus reveal with ''high confidence'' that both terrestrial and marine CH <sub>4</sub> and N <sub>2</sub> O emissions were reduced under glacial climate conditions. Multiple lines of evidence indicate with ''high confidence'' that enhanced storage of remineralized CO <sub>2</sub> in the ocean interior, owing to a combination of synergistic mechanisms, was sufficient to balance the removal of carbon from the atmosphere and the terrestrial biosphere reservoirs combined during the last ice age.

Vegetation regrowth and increased precipitation in wetland regions associated with the mid-deglacial Northern Hemisphere warming (referred to as the Bølling/Allerød (B/A) warm interval, 14.7–12.7 ka), in particular in the (sub)tropics, accounts for large increases in both CH <sub>4</sub> and N <sub>2</sub> O emissions to the atmosphere ( [[#Baumgartner--2014|Baumgartner et al., 2014]] ; [[#Schilt--2014|Schilt et al., 2014]] ; [[#Bock--2017|Bock et al., 2017]] ; H. [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ). Specifically, changes in CH <sub>4</sub> sources were steered by variations in vegetation productivity, source size area, temperatures and precipitation as modulated by insolation, local sea level changes and monsoon intensity ( [[#Bock--2017|Bock et al., 2017]] ; [[#Kleinen--2020|Kleinen et al., 2020]] ). Changes in the CH <sub>4</sub> atmospheric sink term probably only played a secondary role in modulating atmospheric CH <sub>4</sub> inventories across the LDT ( [[#Hopcroft--2017|Hopcroft et al., 2017]] ; [[#Kleinen--2020|Kleinen et al., 2020]] ) Geological emissions, related to the destabilization of fossil (radiocarbon-dead) CH <sub>4</sub> sources buried in continental margins as a result of sudden warming, appear small ( [[#Bock--2017|Bock et al., 2017]] ; [[#Petrenko--2017|Petrenko et al., 2017]] ; [[#Dyonisius--2020|Dyonisius et al., 2020]] ). Stable isotope analysis on N <sub>2</sub> O extracted from Antarctic and Greenland ice reveal that marine and terrestrial emissions increased by 0.7 ± 0.3 and 1.7 ± 0.3 TgN, respectively, across the LDT ( [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ). During abrupt Northern Hemisphere warmings, terrestrial emissions responded rapidly to the northward displacement of the Intertropical Convergence Zone (ITCZ) associated with the resumption of the Atlantic meridional overturning circulation (AMOC; H. [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ). About 90% of these step increases occurred rapidly, possibly in less than 200 years ( [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ). In contrast, marine emissions increased more gradually, modulated by global ocean circulation reorganization.

The gradual increase in atmospheric CO <sub>2</sub> across the LDT was punctuated by three centennial 10–13 ppm increments, coeval with 100–200 ppb increases in CH <sub>4</sub> ( [[#Marcott--2014|Marcott et al., 2014]] ), reminiscent of similar oscillations reported for the last ice age associated with transient warming events (Dansgaard/Oeschger (DO) events; [[#Ahn--2014|Ahn and Brook, 2014]] ; [[#Rhodes--2017|Rhodes et al., 2017]] ; [[#Bauska--2018|Bauska et al., 2018]] ) as well as previous deglacial transitions ( [[#Nehrbass-Ahles--2020|Nehrbass-Ahles et al., 2020]] ). The rate of change in atmospheric CO <sub>2</sub> accumulation during these transient events exceeds the averaged deglacial growth rates by at least 50% (Table 2.1, Figure 5.4). The early deglacial release of remineralized carbon from the ocean abyss coincided with the resumption of Southern Ocean overturning circulation ( [[#Skinner--2010|Skinner et al.,2010]] ; [[#Schmitt--2012|Schmitt et al., 2012]] ; [[#Ferrari--2014|Ferrari et al., 2014]] ; [[#Gottschalk--2016|Gottschalk et al., 2016]] , 2020a; [[#Jaccard--2016|Jaccard et al., 2016]] ; [[#Rae--2018|Rae et al., 2018]] ; [[#Moy--2019|Moy et al., 2019]] ) and the concomitant reduction in the global efficiency of the marine BCP, associated, in part, with dwindling iron fertilization ( [[#Hain--2010|Hain et al., 2010]] ; [[#Martínez-García--2014|Martínez-García et al., 2014]] ; [[#Jaccard--2016|Jaccard et al., 2016]] ) The two subsequent pulses, centred 14.8 and 12.9 ka, are associated with enhanced air–sea gas exchange in the Southern Ocean (T. [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ), iron fertilization in the South Atlantic and North Pacific ( [[#Lambert--2021|Lambert et al., 2021]] ) and rapid increase in soil respiration owing to the resumption of AMOC and associated southward migration of the ITCZ ( [[#Marcottet--2014|Marcottet al., 2014]] ; [[#Bauska--2018|Bauska et al., 2018]] ). Rapid warming of high northern latitudes contributed to thaw permafrost, possibly liberating labile organic carbon to the atmosphere (Köhler et al.,2014; [[#Crichton--2016|Crichton et al., 2016]] ; [[#Winterfeld--2018|Winterfeld et al., 2018]] ; [[#Meyer--2019|Meyer et al., 2019]] ). Ocean surface pH reconstructions indicate that the ocean was oversaturated with respect to the atmosphere during the early, mid-LDT ( [[#Martínez-Botí--2015b|Martínez-Botí et al., 2015b]] ; [[#Shao--2019|Shao et al., 2019]] ; [[#Shuttleworth--2021|Shuttleworth et al., 2021]] ), suggesting that ocean sources at that time may have been larger than terrestrial sources. Over the course of the LDT, the decrease in Northern Hemisphere permafrost carbon stocks has been more than compensated by an increase in the carbon stocks of mineral soils, peatland and vegetation ( [[#Lindgren--2018|Lindgren et al., 2018]] ; [[#Jeltsch-Thömmes--2019|Jeltsch-Thömmes et al., 2019]] ). The land biosphere was, on average, a net sink for atmospheric carbon and accumulated several hundred Gt of carbon over the LDT. Detailed investigations reveal that Antarctic air temperatures, and more generally Southern Hemisphere (30°S–60°S) proxy temperature reconstructions, led the rise in ''p'' CO <sub>2</sub> at the onset of the LDT, 18 ka ago, by several hundred years ( [[#Shakun--2012|Shakun et al., 2012]] ; [[#Chowdhry%20Beeman--2019|Chowdhry Beeman et al., 2019]] ). Atmospheric CO <sub>2</sub> led reconstructed global average temperature by several centuries ( [[#Shakun--2012|Shakun et al., 2012]] ), corroborating the importance of CO <sub>2</sub> as an amplifier of orbitally driven warming. During the LDT, the phasing between Antarctic air temperature and atmospheric GHG concentration changes was nearly synchronous, yet variable, owing to the complex nature of the mechanisms modulating the global carbon cycle ( [[#Chowdhry%20Beeman--2019|Chowdhry Beeman et al., 2019]] ). Mean ocean temperature reconstructions, based on noble gas extracted from Antarctic ice are closely correlated with Antarctic air temperature and ''p'' CO <sub>2</sub> records, emphasizing the role the Southern Ocean is playing in modulating global climate variability ( [[#Bereiter--2018|Bereiter et al., 2018]] ; [[#Baggenstos--2019|Baggenstos et al., 2019]] ).

Enhanced mid-ocean ridge magmatism and/or hydrothermal activity modulated by sea level rise has recently been hypothesized to have contributed to the deglacial CO <sub>2</sub> rise ( [[#Crowley--2015|Crowley et al., 2015]] ; [[#Lund--2016|Lund et al., 2016]] ; [[#Huybers--2017|Huybers and Langmuir, 2017]] ; [[#Stott--2019b|Stott et al., 2019b]] ). While geological carbon release may have affected the ocean’s radiocarbon budget ( [[#Ronge--2016|Ronge et al., 2016]] ; [[#Rafter--2019|Rafter et al., 2019]] ; [[#Stott--2019a|Stott et al., 2019a]] ), model results suggest that the potential contribution of geological carbon sources to the atmosphere remained small (Roth and Joos, 2012; [[#Hasenclever--2017|Hasenclever et al., 2017]] ).

Simulations of Earth models of intermediate complexity (EMIC) with coupled glacial–interglacial climate and the carbon cycle were able to reproduce first-order changes in the atmospheric CO <sub>2</sub> content for the first time in recent years (Ganopolski and Brovkin, 2017; [[#Khatiwala--2019|Khatiwala et al., 2019]] ). The most important processes accounting for the full deglacial CO <sub>2</sub> amplitude in the models include solubility changes, changes in oceanic circulation and marine carbonate chemistry. The effect of the terrestrial carbon cycle, variable volcanic outgassing and the temperature dependence on the oceanic remineralization length scale contribute less than 15 ppm CO <sub>2</sub> between the glacial and interglacial intervals of the cycles. However, details in the simulated response of the marine carbon cycle and atmospheric CO <sub>2</sub> concentrations to changes in ocean circulation depend to a large degree on model parametrization ( [[#Gottschalk--2019|Gottschalk et al., 2019]] ).

Independent paleoclimatic evidence suggests with ''high confidence'' that marine and terrestrial CH <sub>4</sub> and N <sub>2</sub> O emissions are highly sensitive to climate on (sub)centennial time scales. Limited, yet internally consistent ice core measurements indicate with ''medium confidence'' that pulsed geologic CH <sub>4</sub> release from continental margins associated with warming remained negligible across the LDT. Multiple lines of evidence suggest with ''high confidence'' that CO <sub>2</sub> was released from the ocean interior on centennial time scales during the LDT in response to, or associated with warming, contributing to the transition out of the last glacial stage to the current interglacial period.

Multiple lines of evidence inferred from marine sediment proxies indicate with ''low to medium confidence'' that the millennial rates of CO <sub>2</sub> concentration change in the atmosphere during the last 56 Myr were at least four to five times lower than during the last century (Figure 5.3). In spite of uncertainties in ice core reconstructions related to delayed enclosure of air bubbles, which tend to smooth the records, there is ''high confidence'' that the rates of atmospheric CO <sub>2</sub> and CH <sub>4</sub> change during the last century were at least 10 and 5 times faster, respectively, than the maximum centennial growth rate averages of those gases during the last 800 kyr (Fig. 5.4).

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