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

== 5.1 Introduction ==

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The physical and biogeochemical controls of greenhouse gases (GHGs) is a central motivation for this chapter, which identifies biogeochemical feedbacks that have led or could lead to a future acceleration, slowdown or abrupt transitions in the rate of GHG accumulation in the atmosphere, and therefore of climate change. A characterization of the trends and feedbacks lead to improved quantification for the remaining carbon budgets for climate stabilization, and the responses of the carbon cycle to atmospheric carbon dioxide removal (CDR), which is embedded in many of the mitigation scenarios, to achieve the goals of the Paris Agreement.

Changes in the abundance of well-mixed GHGs – carbon dioxide (CO <sub>2</sub> ), methane (CH <sub>4</sub> ) and nitrous oxide (N <sub>2</sub> O) – in the atmosphere play a large role in determining the Earth’s radiative properties and its climate in the past, the present and the future (Chapters 2, 4, 6 and 7). Since 1950, the increase in atmospheric GHGs has been the dominant cause of the human-induced climate change ( [[IPCC:Wg1:Chapter:Chapter-3#3.3|Section 3.3]] ). While the main driver of changes in atmospheric GHGs over the past 200 years relates to the direct emissions from human activities, the net accumulation of GHGs in the atmosphere is controlled by biogeochemical source-sink dynamics of carbon that exchange between multiple reservoirs on land, oceans and atmosphere. The combustion of fossil fuels and land-use change for the period 1750–2019 released an estimated 700 ± 75 PgC (1 PgC = 10 <sup>15</sup> g of carbon) into the atmosphere, of which less than half remains in the atmosphere today (Sections 5.2.1.2; 5.2.1.5) ( [[#Friedlingstein--2020|Friedlingstein et al., 2020]] ). This emphasizes the central role of terrestrial and ocean CO <sub>2</sub> sinks in regulating its atmospheric concentration ( [[#Ballantyne--2012|Ballantyne et al., 2012]] ; [[#Li--2016|]] [[#Li--2016|]] [[#Li--2016|W. Li et al., 2016]] ; [[#Le%20Quéré--2018a|Le Quéré et al., 2018a]] ; [[#Ciais--2019|Ciais et al., 2019]] ; [[#Gruber--2019b|Gruber et al., 2019b]] ; [[#Friedlingstein--2020|Friedlingstein et al., 2020]] ).

The chapter covers three dominant GHGs in the human perturbation of the Earth’s radiation budget for which high-quality records exist: CO <sub>2</sub> , CH <sub>4</sub> and N <sub>2</sub> O (Figure 5.1).

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[[File:06e0852e552964cbebb793bb360bf7a3 IPCC_AR6_WGI_Figure_5_1.png]]

'''Figure 5.1 |''' '''Visual guid''' '''e to Chapter 5.'''

( [[#5.1|Section 5.1]] (this section) provides the time context on how unique current and future scenarios of GHGs atmospheric concentrations and growth rates are in the Earth’s history. It also introduces the main processes involved in carbon–climate feedbacks, followed by an assessment of what can be learned from the paleo record towards a better understanding of contemporary and future GHGs–climate dynamics and their response to different mitigation trajectories.

( [[#5.2|Section 5.2]] covers the state of the carbon cycle and other biogeochemical cycles, and global budgets of CO <sub>2</sub> , CH <sub>4</sub> and N <sub>2</sub> O for the industrial era (since 1750). The section emphasizes the last 60-year period for which high-resolution observations are available and the most recent decade for comprehensive GHG budgets. Significant advances have taken place since the IPCC Fifth Assessment Report (AR5), particularly in constraining the annual-to-decadal variability of the ocean and land carbon sources and sinks, and in revealing the sensitivity of carbon pools to current and future climate changes. There has been an important increase in modelling capability of the three GHGs, for land and oceans, atmospheric and ocean observations, and remote sensing products that have enabled researchers to constrain the causes of the observed trends and variability.

( [[#5.3|Section 5.3]] builds on the Special Report on the Ocean and Cryosphere (SROCC) covering the change in ocean acidification due to oceanic CO <sub>2</sub> uptake across the paleo, historical periods and future projections using Coupled Model Intercomparison Project Phase 6 (CMIP6), with consequences for marine life (assessed in the Sixth Assessment Report Working Group II, AR6 WGII) and biogeochemical cycles. The section also assesses changes in deoxygenation of the oceans due to warming, increased stratification of the surface ocean, and slowing of the meridional overturning circulation.

( [[#5.4|Section 5.4]] covers the future projections of biogeochemical cycles and their feedbacks to the climate system fully utilizing the database of the concentration-driven CMIP6. Since AR5, Earth system models (ESMs) have made progress towards including more complex carbon cycle and associated biogeochemical processes that enable exploring a range of possible future carbon–climate feedbacks and their influences on the climate system. The section addresses uncertainties and limits of our models to predict future dynamics for GHG emissions trajectories, as well as new understanding on processes involved in carbon–climate feedbacks and the possibility for rapid and abrupt changes brought by non-linear dynamics.

( [[#5.5|Section 5.5]] covers the development of the total and remaining carbon budgets to climate stabilization targets and the associated transient climate response to cumulative CO <sub>2</sub> emissions. The section shows the progress made since AR5 ( [[#IPCC--2013a|IPCC, 2013a]] ) and the Special Report on Global Warming of 1.5°C ( [[#IPCC--2018|IPCC, 2018]] ), particularly on key components required to estimate the remaining carbon budget, including the transient response to cumulative emissions of CO <sub>2</sub> , the zero emissions commitment, the projected non-CO <sub>2</sub> warming, and the unrepresented Earth system feedbacks.

( [[#5.6|Section 5.6]] assesses the impacts of CDR and solar radiation modification for the purpose of climate change mitigation on the global carbon cycle, building from the assessment in the IPCC Special Report on Climate Change and Land (SRCCL). It includes an overview of the major CDR options and potential collateral biogeochemical effects beyond the intended climate change mitigation strategies. The potential capacity to deliver atmospheric reductions and the socio-economic feasibility of such options are assessed in detail in AR6 working group III (WGIII).

Finally, [[#5.7|Section 5.7]] highlights the knowledge gaps as limits to the assessment. The assessment would have been strengthened had those gaps not existed.

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=== 5.1.1 The Physical and Biogeochemical Processes in Carbon–Climate Feedbacks ===

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The influence of anthropogenic CO <sub>2</sub> emissions and emissions scenarios on the carbon–climate system is the primary driver of ocean and terrestrial sinks as the major negative feedbacks that determine the atmospheric CO <sub>2</sub> levels, which then drive climate feedbacks through radiative forcing (Figure 5.2) ( [[#Friedlingstein--2006|Friedlingstein et al., 2006]] ; [[#Jones--2013|]] [[#Jones--2013|Jones et al., 2013]] ; [[#Jones--2020|Jones and Friedlingstein, 2020]] ). Biogeochemical feedbacks follow as an outcome of both carbon and climate forcing on the physics and the biogeochemical processes of the ocean and terrestrial carbon cycles (Figure 5.2) ( [[#Katavouta--2018|Katavouta et al., 2018]] ; [[#Williams--2019|Williams et al., 2019]] ; [[#Jones--2020|Jones and Friedlingstein, 2020]] ). Together, these carbon–climate feedbacks can amplify or suppress climate change by altering the rate at which CO <sub>2</sub> builds up in the atmosphere through changes in the land and ocean sources and sinks (Figure 5.2; C.D. [[#Jones--2013|]] [[#Jones--2013|Jones et al., 2013]] ; [[#Raupach--2014|Raupach et al., 2014]] ; [[#Williams--2019|Williams et al., 2019]] ). These changes depend on the, often non-linear, interaction of the drivers (CO <sub>2</sub> and climate) and processes in the ocean and land as well as the emissions scenarios (Figure 5.2; Sections 5.4 and 5.6) ( [[#Raupach--2014|Raupach et al., 2014]] ; [[#Schwinger--2014|Schwinger et al., 2014]] ; [[#Williams--2019|Williams et al., 2019]] ). There is ''high'' ''confidence'' that carbon–climate feedbacks and their century scale evolution play a critical role in two linked climate metrics that have significant climate and policy implications: (i) the fraction of anthropogenic CO <sub>2</sub> emissions that remains in the atmosphere, the so-called airborne fraction of CO <sub>2</sub> (AF; [[#5.2.1.2|Section 5.2.1.2]] , Figures 5.2 and 5.7, and FAQ 5.1); and (ii) the quasi-linear trend characteristic of the transient temperature response to cumulative CO <sub>2</sub> emissions (TCRE; [[#5.5|Section 5.5]] ; [[#MacDougall--2016|MacDougall, 2016]] ; [[#Williams--2016|Williams et al., 2016]] ; [[#Jones--2020|Jones and Friedlingstein, 2020]] ) and other GHGs (CH <sub>4</sub> and N <sub>2</sub> O). This chapter assesses the implications of these issues from the perspective of carbon cycle processes (Figure 5.2) in [[#5.2|Section 5.2]] (historical and contemporary), [[#5.3|Section 5.3]] (changing carbonate chemistry), [[#5.4|Section 5.4]] (future projections), [[#5.5|Section 5.5]] (remaining carbon budget) and [[#5.6|Section 5.6]] (response to carbon dioxide removal and solar radiation modification).

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[[File:439c36b87d834f98da0d29ab09480871 IPCC_AR6_WGI_Figure_5_2.png]]

'''Figure 5.2 |''' '''Key compartments, processes and pathways that govern historical and future CO''' <sub>2</sub> '''concentrations and carbon–climate feedbacks through the coupled Earth system.''' The anthropogenic CO <sub>2</sub> emissions, including land-use change, are partitioned via negative feedbacks (turquoise dotted arrows) between the ocean (23%), the land (31%) and the airborne fraction (46%) of anthropogenic CO <sub>2</sub> that sets the changing CO <sub>2</sub> concentration in the atmosphere (2010–2019; Table 5.1). This regulates most of the radiative forcing that creates the heat imbalance that drives the climate feedbacks to the ocean (blue) and land (green). Positive feedbacks (red arrows) result from processes in the ocean and on land (red text). Positive feedbacks are influenced by both carbon-concentration and carbon–climate feedbacks simultaneously. Additional biosphere processes have been included, but these have an as-yet-uncertain feedback impact (blue-dotted arrows). CO <sub>2</sub> removal from the atmosphere into the ocean, land and geological reservoirs, necessary for negative emissions, has been included (grey arrows). Although this schematic is built around CO <sub>2</sub> (the dominant greenhouse gas), some of the same processes also influence the fluxes of CH <sub>4</sub> and N <sub>2</sub> O and the strength of the positive feedbacks from the terrestrial and ocean systems.

The airborne fraction is an important constraint for adjustments in carbon–climate feedbacks and reflects the partitioning of CO <sub>2</sub> emissions between reservoirs by the negative feedbacks, which were 31% on land and 23% in the ocean for the decade 2010–2019 and also dominated the historical period (Figure 5.2; Table 5.1) ( [[#Friedlingstein--2020|Friedlingstein et al., 2020]] ). During the period 1959–2019, the airborne fraction has largely followed the growth in anthropogenic CO <sub>2</sub> emissions with a mean of 44% and a large interannual variability ( [[#Ballantyne--2012|Ballantyne et al., 2012]] ; [[#Ciais--2019|Ciais et al., 2019]] ; [[#Friedlingstein--2020|Friedlingstein et al., 2020]] , [[#5.2.1.2|Section 5.2.1.2]] ; Table 5.1). The negative feedback to CO <sub>2</sub> concentrations is associated with its impact on the air–sea and air–land CO <sub>2</sub> exchange through strengthening of partial pressure of CO <sub>2</sub> ( ''p'' CO <sub>2</sub> ) gradients as well as the internal processes that enhance uptake. Two of these key processes are the buffering capacity of the ocean and the CO <sub>2</sub> fertilization effect on gross primary production (Sections 5.4.1–5.4.4).

Positive and negative climate and carbon feedbacks involve: (i) fast processes on land and oceans at time scales from minutes to years, such as photosynthesis, soil respiration, net primary production, shallow ocean physics and air–sea fluxes; and (ii) slower processes taking from decades to millennia, such as changing ocean buffering capacity, ocean ventilation, vegetation dynamics, permafrost changes, peat formation and decomposition (Figure 5.2; [[#Ciais--2013|Ciais et al., 2013]] ; [[#Forzieri--2017|Forzieri et al., 2017]] ; [[#Williams--2019|Williams et al., 2019]] ). Depending on the particular combination of driver process and response dynamics, they behave as positive or negative feedbacks that amplify or dampen the magnitude and rates of climate change, respectively ( [[#Cox--2000|Cox et al., 2000]] ; [[#Friedlingstein--2003|Friedlingstein et al., 2003]] , [[#Friedlingstein--2006|2006]] ; [[#Hauck--2015|Hauck and Völker, 2015]] ; [[#Williams--2019|Williams et al., 2019]] ); red and turquoise arrows in Figure 5.2 and Table 5.1).

Carbon cycle feedbacks co-exist with climate (heat and moisture) feedbacks (Cross-Chapter Boxes 5.1 and 5.3), which together drive contemporary ( [[#5.2|Section 5.2]] ) and future ( [[#5.4|Section 5.4]] ) carbon–climate feedbacks ( [[#Williams--2019|Williams et al., 2019]] ). The excess heat generated by radiative forcing from increasing concentration of atmospheric CO <sub>2</sub> and other GHGs is mostly taken up by the ocean (&gt;90%) and the residual balance partitioned between atmospheric, terrestrial and ice melting (Cross-Chapter Box 9.2; [[#Frölicher--2015|Frölicher et al., 2015]] ). The combined effect of these two large-scale negative feedbacks of CO <sub>2</sub> and heat are reflected in the TCRE ( [[#5.5|Section 5.5]] and Cross-Chapter Box 5.3), which points to a quasi-linear and quasi-emission-path independent relationship between cumulative emissions of CO <sub>2</sub> and global warming, which is used as the basis to estimate the remaining carbon budget ( [[#5.5|Section 5.5]] ; [[#MacDougall--2015|MacDougall and Friedlingstein, 2015]] ; [[#MacDougall--2017|MacDougall, 2017]] ; [[#Bronselaer--2020|Bronselaer and Zanna, 2020]] ; [[#Jones--2020|Jones and Friedlingstein, 2020]] ). There is still ''low confidence'' on the relative roles and importance of the ocean and terrestrial carbon processes on TCRE variability and uncertainty on centennial time scales ( [[#MacDougall--2016|MacDougall, 2016]] ; [[#MacDougall--2017|MacDougall et al., 2017]] ; [[#Williams--2017|Williams et al., 2017]] ; [[#Katavouta--2018|Katavouta et al., 2018]] , [[#Katavouta--2019|2019]] ; [[#Jones--2020|Jones and Friedlingstein, 2020]] ) (Sections 5.5.1.1, 5.5.1.2).

The combined effects of climate and CO <sub>2</sub> concentration feedbacks on the global carbon cycle are projected by ESMs to modify both the processes and natural reservoirs of carbon on a regional and global scale that may result in positive feedbacks (red arrows in Figure 5.2), which could weaken the major terrestrial and ocean sinks and disrupt the airborne fraction and TCRE under medium- to high-emissions scenarios ( [[#5.4.5|Section 5.4.5]] and Figure 5.25).

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=== 5.1.2 Paleo Trends and Feedbacks ===

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Paleoclimatic proxy records extend beyond the variability of recent decadal climate oscillations and thus provide an independent perspective on feedbacks between climate and carbon cycle dynamics. According to reconstructions, these past changes were slower than the current anthropogenic ones, so they cannot provide an unequivocal comparison. Nonetheless, they can help appraise sensitivities and point towards potentially dominant mechanisms of change ( [[#Tierney--2020|Tierney et al., 2020]] ) on (sub)centennial to (multi)millennial time scales.

The AR5 (WGI, Chapter 5) concluded with ''medium'' ''confidence'' that atmospheric CO <sub>2</sub> concentrations reached 350–450 ppm during the mid-Pliocene (3.3–3.0 million years ago (Ma)), and possibly 1000 ppm during the Early Eocene (52–48 Ma). The AR5 (WGI, Chapter 5) also concluded with ''very high confidence'' that the current rates of increases in CO <sub>2</sub> , CH <sub>4</sub> and N <sub>2</sub> O atmospheric concentrations were unprecedented with respect to the ice core record covering the last deglacial transition (LDT, 18–11 ka) and with ''medium confidence'' that the rate of change of the reconstructed GHG rise was also unprecedented compared to the lower resolution of the records of the past 800 kyr.

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==== 5.1.2.1 Cenozoic Proxy CO <sub>2</sub> Record ====

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Quantifying past changes in the rate of CO <sub>2</sub> accumulation in the atmosphere based on reconstructions using marine sediment proxies is complex as age model uncertainties, assumptions and shortcomings underlying proxy applications and sedimentary processes conspire to alter and confound rate estimates ( [[#Ajayi--2020|Ajayi et al., 2020]] ). Differential sediment mixing and bioturbation contribute to smooth and attenuate proxy records ( [[#Hupp--2020|Hupp and Kelly, 2020]] ), thereby tending to underestimate maximum rates of change ( [[#Kemp--2015|Kemp et al., 2015]] ). Considering the extent to which uncertainties can affect sediment-based rate estimates, and notwithstanding recent effort in minimizing their inherent contribution, there is generally ''low to medium confidence'' in quantifying rates of change on a time scale less than a decade back thousands of years, and less than a millennium back millions of years in the past based on marine sediments.

In the past, atmospheric CO <sub>2</sub> concentrations reached much higher levels than present day (Cross-Chapter Box 2.1 and Figure 5.3). In particular, the Paleocene–Eocene thermal maximum (PETM), 55.9–55.7 Ma (Figure 5.3), provides some level of comparison with the current and projected anthropogenic increase in CO <sub>2</sub> emissions (Chapter 2). Atmospheric CO <sub>2</sub> concentrations increased from about 900 to around 2000 ppm in 3–20 kyr as a result of geological carbon release to the ocean–atmosphere system ( [[#Zeebe--2016|Zeebe et al., 2016]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Cui--2018|Cui and Schubert, 2018]] ; [[#Kirtland%20Turner--2018|Kirtland Turner, 2018]] ). There is ''low'' to ''medium confidence'' in evaluations of the total amount of carbon released during the PETM, as proxy data constrained estimates vary from around 3000 to more than 7000 PgC, with methane hydrates, volcanic emissions, terrestrial and/or marine organic carbon, or some combination thereof, as the probable sources of carbon ( [[#Zeebe--2009|Zeebe et al., 2009]] ; [[#Cui--2011|Cui et al., 2011]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Elling--2019|Elling et al., 2019]] ; [[#Jones--2019|]] [[#Jones--2019|Jones et al., 2019]] ; [[#Haynes--2020|Haynes and Hönisch, 2020]] ). Methane emissions related to hydrate/permafrost thawing and fossil carbon oxidation may have acted as positive feedbacks ( [[#Lunt--2011|Lunt et al., 2011]] ; [[#Armstrong%20McKay--2018|Armstrong McKay and Lenton, 2018]] ; [[#Lyons--2019|Lyons et al., 2019]] ), as the inferred increase in atmospheric CO <sub>2</sub> can only account for approximately half of the reported warming ( [[#Zeebe--2009|Zeebe et al., 2009]] ). The estimated, time-integrated carbon input is broadly similar to the RCP8.5 extension scenario, although CO <sub>2</sub> emissions rates (0.3–1.5 Pg yr <sup>–1</sup> ) and by inference the rate of CO <sub>2</sub> accumulation in the atmosphere (4–42 ppm per century) during the PETM were at least 4–5 lower than during the modern era (from 1995 to 2014; Table 2.1; [[#Zeebe--2016|Zeebe et al., 2016]] ; [[#Gingerich--2019|Gingerich, 2019]] ).

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[[File:216de2b6c3c945cc85550154b813b7ff IPCC_AR6_WGI_Figure_5_3.png]]

'''Figure 5.3 |''' '''Atmospheric CO''' <sub>2</sub> '''concentrations and growth rates for the past 60 million years (Myr) and projections to 2100.''' '''(a)''' CO <sub>2</sub> concentrations data for the period 60 Myr to the time prior to 800 kyr (left column) are shown as the LOESS Fit and 68% range (data from Chapter 2) ( [[#Foster--2017|Foster et al., 2017]] ). Concentrations from 1750 and projections through 2100 are taken from Shared Socio-economic Pathways of IPCC AR6 ( [[#Meinshausen--2017|Meinshausen et al., 2017]] ). '''(b)''' Growth rates are shown as the time derivative of the concentration time series. Inserts in (b) show growth rates at the scale of the sampling resolution. Further details on data sources and processing are available in the chapter data table (Table 5.SM.6).

The last 50 Myr (50 million years) have been characterized by a gradual decline in atmospheric CO <sub>2</sub> levels at a rate of about 16 ppm Myr <sup>–1</sup> (Figure 5.3; [[#Foster--2017|Foster et al., 2017]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ). The exact cause of this long-term change in CO <sub>2</sub> remains uncertain, but may be related to an imbalance between long-term sources of CO <sub>2</sub> (volcanic outgassing) and long-term sinks (organic carbon burial and silicate weathering).

The most recent time interval when atmospheric CO <sub>2</sub> concentration was as high as 1000 ppm (i.e., similar to the end of 21st century projection for the high-end emissions scenario RCP8.5) was around 33.5 Ma, prior to the Eocene-Oligocene transition ( [[#Zhang--2013|]] [[#Zhang--2013|Zhang et al., 2013]] ; [[#Anagnostou--2016|Anagnostou et al., 2016]] ). Atmospheric CO <sub>2</sub> levels then reached a critical threshold (1000–750 ppm; [[#DeConto--2008|DeConto et al., 2008]] ) to allow for the development of permanent regional ice sheets on Antarctica, associated with changes in Southern Ocean hydrography, which would have increased deep ocean CO <sub>2</sub> storage ( [[#Leutert--2020|Leutert et al., 2020]] ).

The most recent interval characterized by atmospheric CO <sub>2</sub> levels similar to modern (i.e., 360–420 ppm) was the mid-Pliocene Warm Period (MPWP, 3.3–3.0 Ma; Martínez-Botí et al., 2015; [[#de%20la%20Vega--2020|de la Vega et al., 2020]] ) (Chapter 2). The relatively high atmospheric CO <sub>2</sub> concentration during the MPWP are related to vigorous ocean circulation and a rather inefficient marine biological carbon pump ( [[#Burls--2017|Burls et al., 2017]] ), which would have reduced deep ocean carbon storage. After the MPWP, atmospheric CO <sub>2</sub> concentrations declined gradually at a rate of 30 ppm Myr <sup>–1</sup> (Figure 5.3; [[#de%20la%20Vega--2020|de la Vega et al., 2020]] ), as an increase in ocean stratification led to enhanced ocean carbon storage, allowing for major, sustained advances in Northern Hemisphere ice sheets, 2.7 Ma ( [[#Sigman--2004|Sigman et al., 2004]] ; [[#DeConto--2008|DeConto et al., 2008]] ).

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==== 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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[[File:edf7550e1705ff332a99148a9bb8a71e IPCC_AR6_WGI_Figure_5_4.png]]

'''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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==== 5.1.2.3 Holocene Changes ====

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Atmospheric GHG concentrations were much less variable during the pre-industrial Holocene (from 11.7 ka to 1750 CE). Atmospheric CH <sub>4</sub> concentrations decreased at the beginning of the Holocene, consistent with a general weakening of boreal sources ( [[#Yang--2017|Yang et al., 2017]] ; [[#Beck--2018|Beck et al., 2018]] ) and further decline during the mid-Holocene owing to a reduction in Southern Hemisphere emissions concomitant with a southward shift of the ITCZ ( [[#Singarayer--2011|Singarayer et al., 2011]] ; [[#Beck--2018|Beck et al., 2018]] ). Atmospheric CH <sub>4</sub> concentrations increased about 5 ka, which prompted the hypothesis of an early anthropogenic influence related to land-use changes in South East Asia ( [[#Ruddiman--2016|Ruddiman et al., 2016]] ). However, stable isotope compositions on CH <sub>4</sub> extracted from Greenland and Antarctic ice ( [[#Beck--2018|Beck et al., 2018]] ) reveal that natural emissions located in the southern tropics were responsible for the rise in atmospheric CH <sub>4</sub> concentrations, in line with model simulations ( [[#Singarayer--2011|Singarayer et al., 2011]] ) thus disputing the early anthropogenic influence on the global CH <sub>4</sub> budget. Atmospheric N <sub>2</sub> O concentrations increased slightly (20 ppb) across the Holocene, associated with a gradual decline in its nitrogen stable isotope composition (H. [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ). The combined signal is consistent with a small increase in terrestrial emissions, offset by a reduction in marine emissions ( [[#Schilt--2010b|Schilt et al., 2010b]] ; [[#Fischer--2019|]] [[#Fischer--2019|Fischer et al., 2019]] ).

The early Holocene decrease in CO <sub>2</sub> concentration by about 5 ppm ( [[#Schmitt--2012|Schmitt et al., 2012]] ) has been attributed to post-glacial regrowth in terrestrial biomass and a gradual increase in peat reservoirs over the Holocene, resulting in the sequestration of several hundred PgC ( [[#Yu--2010|Yu et al., 2010]] ; [[#Nichols--2019|Nichols and Peteet, 2019]] ). Peat accumulation rates in boreal and temperate regions were higher under warmer summer conditions in the early to mid-Holocene ( [[#Loisel--2014|Loisel et al., 2014]] ; [[#Stocker--2017|Stocker et al., 2017]] ). The 20 ppm gradual increase of atmospheric CO <sub>2</sub> starting 7 ka has been attributed to a decrease in natural terrestrial biomass due to climate change, carbonate compensation and enhanced shallow water carbonate deposition ( [[#Menviel--2012|Menviel and Joos, 2012]] ; [[#Brovkin--2016|Brovkin et al., 2016]] ), consistent with stable carbon isotope measurements on CO <sub>2</sub> extracted from Antarctic ice ( [[#Elsig--2009|Elsig et al., 2009]] ; [[#Schmitt--2012|Schmitt et al., 2012]] ). These isotopic measurements do not support an early anthropogenic influence on atmospheric CO <sub>2</sub> due to land-use change and forest clearing ( [[#Ruddiman--2016|Ruddiman et al., 2016]] ). Recent paleoceanographic evidence suggests that remineralized carbon outgassing associated with increased Southern Ocean circulation and upwelling ( [[#Studer--2018|Studer et al., 2018]] ), possibly promoted by stronger Southern Hemisphere westerly winds ( [[#Saunders--2018|Saunders et al., 2018]] ), could have additionally contributed to the late Holocene increase in atmospheric CO <sub>2</sub> concentrations. However, the role of these mechanisms remained insignificant in transient Holocene ESM simulations ( [[#Brovkin--2019|Brovkin et al., 2019]] ). Overall, as in AR5 (WGI, Chapter 5), there is ''medium confidence'' in the key drivers of the CO <sub>2</sub> increase between the early Holocene and the beginning of the industrial era, yet there is ''low confidence'' in the relative contributions of these drivers due to insufficient quantitative constraints on particular processes.

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