Editing IPCC:AR6/WGI/Chapter-2 (section)

== 2.2 Changes in Climate Drivers ==

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This section assesses the magnitude and rates of changes in both natural and anthropogenically mediated climate drivers over a range of time scales. First, changes in insolation (orbital and solar; [[#2.2.1|Section 2.2.1]] ), and volcanic stratospheric aerosol ( [[#2.2.2|Section 2.2.2]] ) are assessed. Next, well-mixed greenhouse gases (GHGs; CO <sub>2</sub> , N <sub>2</sub> O and CH <sub>4</sub> ) are covered in [[#2.2.3|Section 2.2.3]] , with climate feedbacks and other processes involved in the carbon cycle assessed in Chapter 5. The section continues with the assessment of changes in halogenated GHGs ( [[#2.2.4|Section 2.2.4]] ), stratospheric water vapour, stratospheric and tropospheric ozone ( [[#2.2.5|Section 2.2.5]] ), and tropospheric aerosols ( [[#2.2.6|Section 2.2.6]] ). Short-lived climate forcers (SLCFs), their precursor emissions and key processes are assessed in more detail in Chapter 6. [[#2.2.7|Section 2.2.7]] assesses the effect of historical land cover change on climate, including biophysical and biogeochemical processes. [[#2.2.8|Section 2.2.8]] summarizes the changes in the Earth’s energy balance since 1750 using the comprehensive assessment of effective radiative forcing (ERF) performed in Section 7.3. For some SLCFs with insufficient spatial or temporal observational coverage, ERFs are based on model estimates, but also reported here for completeness and context. Tabulated global mixing ratios of all well-mixed GHGs and ERFs from 1750–2019 are provided in Annex III.

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=== 2.2.1 Solar and Orbital Forcing ===

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The AR5 assessed solar variability over multiple time scales, concluding that total solar irradiance (TSI) multi-millennial fluctuations over the past 9 kyr were &lt;1 W m <sup>–2</sup> , but with no assessment of confidence provided. For multi-decadal to centennial variability over the last millennium, AR5 emphasized reconstructions of TSI that show little change (&lt;0.1%) since the Maunder Minimum (1645–1715) when solar activity was particularly low, again without providing a confidence level. The AR5 further concluded that the best estimate of radiative forcing due to TSI changes for the period 1750–2011 was 0.05–0.10 W m <sup>–2</sup> ( ''medium confidence'' ), and that TSI ''very likely'' changed by –0.04 [–0.08 to 0.00] W m <sup>–2</sup> between 1986 and 2008. Potential solar influences on climate due to feedbacks arising from interactions with galactic cosmic rays are assessed in Section 7.3.4.5.

Slow periodic changes in the Earth’s orbit around the Sun mainly cause variations in seasonal and latitudinal receipt of incoming solar radiation. Precise calculations of orbital variations are available for tens of millions of years ( [[#Berger--1991|Berger and Loutre, 1991]] ; [[#Laskar--2011|Laskar et al., 2011]] ). The range of insolation averaged over boreal summer at 65°N was about 83 W m <sup>−2</sup> during the past million years, and 3.2 W m <sup>−2</sup> during the past millennium, but there was no substantial effect upon global average radiative forcing (0.02 W m <sup>–2</sup> during the past millennium).

A new reconstruction of solar irradiance extends back 9 kyr based upon updated cosmogenic isotope datasets and improved models for production and deposition of cosmogenic nuclides ( [[#Poluianov--2016|Poluianov et al., 2016]] ), and shows that solar activity during the second half of the 20th century was in the upper decile of the range. TSI features millennial-scale changes with typical magnitudes of 1.5 [1.4 to 2.1] W m <sup>–2</sup> (C.-J. [[#Wu--2018|]] [[#Wu--2018|Wu et al., 2018]] ). Although stronger variations in the deeper past cannot be ruled out completely ( [[#Egorova--2018|Egorova et al., 2018]] ; [[#Reinhold--2019|Reinhold et al., 2019]] ), there is no indication of such changes having happened over the last 9 kyr.

Recent estimates of TSI and spectral solar irradiance (SSI) for the past millennium are based upon updated irradiance models (e.g., [[#Egorova--2018|Egorova et al., 2018]] ; C.-J. [[#Wu--2018|]] [[#Wu--2018|Wu et al., 2018]] ) and employ updated and revised direct sunspot observations over the last three centuries ( [[#Clette--2014|Clette et al., 2014]] ; [[#Chatzistergos--2017|Chatzistergos et al., 2017]] ) as well as records of sunspot numbers reconstructed from cosmogenic isotope data prior to this ( [[#Usoskin--2016|Usoskin et al., 2016]] ). These reconstructed TSI time series (Figure 2.2a) feature little variation in TSI averaged over the past millennium. The TSI between the Maunder Minimum (1645–1715) and second half of the 20th century increased by 0.7–2.7 W m <sup>–2</sup> ( [[#Jungclaus--2017|Jungclaus et al., 2017]] ; [[#Egorova--2018|Egorova et al., 2018]] ; [[#Lean--2018|Lean, 2018]] ; C.-J. [[#Wu--2018|]] [[#Wu--2018|Wu et al., 2018]] ; [[#Lockwood--2020|Lockwood and Ball, 2020]] ; [[#Yeo--2020|Yeo et al., 2020]] ). This TSI increase implies a change in ERF of 0.09–0.35 W m <sup>–2</sup> (Section 7.3.4.4).

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[[File:9c8adae7a832062e6d12e3e88dba7c14 IPCC_AR6_WGI_Figure_2_2.png]]

'''Figure''' '''2.2 |''' '''Time series of solar and volcanic forcing for the past 2500 years (a, c) and since 1850 (b, d). (a)''' Total solar irradiance (TSI) reconstruction (10-year running averages) recommended for CMIP6/PMIP4 millennial experiments based on the radiocarbon dataset before 1850 (blue) scaled to the CMIP6 historical forcing after 1850 (purple). '''(b)''' TSI time series (six-month running averages) from CMIP6 historical forcing as inferred from sunspot numbers (blue), compared to CMIP5 forcing based on (red) and an update to CMIP6 by a TSI composite (orange). '''(c)''' Volcanic forcing represented as reconstructed stratospheric aerosol optical depth (SAOD; as presented in Section 7.3.4.6) at 550 nm. Estimates covering 500 BCE to 1900 CE (green) and 1850–2015 (blue). '''(d)''' SAOD reconstruction from CMIP6 (v 4) (blue), compared to CMIP5 forcing (red). Note the change in y-axis range between panels (c) and (d). Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Estimation of TSI changes since 1900 (Figure 2.2b) has further strengthened, and confirms a small (less than about 0.1 W m <sup>–2</sup> ) contribution to global climate forcing (Section 7.3.4.4). New reconstructions of TSI over the 20th century ( [[#Lean--2018|Lean, 2018]] ; [[#Wu--2018|]] [[#Wu--2018|C.-J. Wu et al., 2018]] ) support previous results that the TSI averaged over the solar cycle ''very likely'' increased during the first seven decades of the 20th century and decreased thereafter (Figure 2.2b). TSI did not change significantly between 1986 and 2019. Improved insights ( [[#Krivova--2006|Krivova et al., 2006]] ; [[#Yeo--2015|Yeo et al., 2015]] , 2017; [[#Coddington--2016|Coddington et al., 2016]] ) show that variability in the 200–400 nm UV range was greater than previously assumed. Building on these results, the forcing proposed by [[#Matthes--2017|Matthes et al. (2017)]] has a 16% stronger contribution to TSI variability in this wavelength range compared to the forcing used in the 5th Phase of the Coupled Model Intercomparison Project (CMIP5).

To conclude, solar activity since the late 19th century was relatively high but not exceptional in the context of the past 9 kyr ( ''high confidence'' ). The associated global mean ERF is in the range of –0.06 to +0.08 W m <sup>–2</sup> (Section 7.3.4.4).

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=== 2.2.2 Volcanic Aerosol Forcing ===

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The AR5 concluded that, on interannual time scales, the radiative effects of volcanic aerosols are a dominant natural driver of climate variability, with the greatest effects occurring within the first 2–5 years following a strong eruption. Reconstructions of radiative forcing by volcanic aerosols used in the Paleoclimate Modelling Intercomparison Project Phase III (PMIP3) simulations and in AR5 featured short-lived perturbations of a range of magnitudes, with events of greater magnitude than –1 W m <sup>–2</sup> (annual mean) occurring on average every 35–40 years, although no associated assessment of confidence was given. This section focuses on advances in reconstructions of stratospheric aerosol optical depth (SAOD), whereas ( [[IPCC:Wg1:Chapter:Chapter-7|Chapter 7]] focuses on the ERF of volcanic aerosols, and [[IPCC:Wg1:Chapter:Chapter-5|Chapter 5]] assesses volcanic emissions of CO <sub>2</sub> and CH <sub>4</sub> ; tropospheric aerosols are discussed in [[#2.2.6|Section 2.2.6]] . Cro ss-Ch apter Box 4.1 undertakes an integrative assessment of volcanic effects including potential for 21st century effects.

Advances in analysis of sulphate records from the Greenland Ice Sheet (GrIS) and AIS have resulted in improved dating and completeness of SAOD reconstructions over the past 2.5 kyr ( [[#Sigl--2015|Sigl et al., 2015]] ), a more uncertain extension back to 10 ka ( [[#Kobashi--2017|Kobashi et al., 2017]] ; [[#Toohey--2017|Toohey and Sigl, 2017]] ), and a better differentiation of sulphates that reach high latitudes via stratospheric (strong eruptions) versus tropospheric pathways ( [[#Burke--2019|]] [[#Burke--2019|A. Burke et al., 2019]] ; [[#Gautier--2019|Gautier et al., 2019]] ). The PMIP4 volcanic reconstruction extends the period analysed in AR5 by 1 kyr (Figure 2.2c; [[#Jungclaus--2017|Jungclaus et al., 2017]] ) and features multiple strong events that were previously misdated, underestimated or not detected, particularly before about 1500 CE. The period between successive large volcanic eruptions (Negative ERF greater than –1 W m <sup>–2</sup> ), ranges from 3–130 years, with an average of 43 ± 7.5 years between such eruptions over the past 2.5 kyr (data from [[#Toohey--2017|Toohey and Sigl, 2017]] ). The most recent such eruption was that of Mt Pinatubo in 1991. Century-long periods that lack such large eruptions occurred once every 400 years on average. Systematic uncertainties related to the scaling of sulphate abundance in glacier ice to radiative forcing have been estimated to be about 60% ( [[#Hegerl--2006|Hegerl et al., 2006]] ). Uncertainty in the timing of eruptions in the proxy record is ± 2 years (95% confidence interval) back to 1.5 ka and ± 4 years before ( [[#Toohey--2017|Toohey and Sigl, 2017]] ).

SAOD averaged over the period 950–1250 CE (0.012) was lower than for the period 1450–1850 CE (0.017) and similar to the period 1850–1900 (0.011). Uncertainties associated with these inter-period differences are not well quantified but have little effect because the uncertainties are mainly systematic throughout the record. Over the past 100 years, SAOD averaged 14% lower than the mean of the previous 24 centuries (back to 2.5 ka), and well within the range of centennial-scale variability ( [[#Toohey--2017|Toohey and Sigl, 2017]] ).

Direct observations of volcanic gas-phase sulphur emissions (mostly SO <sub>2</sub> ), sulphate aerosols, and their radiative effects are available from a variety of sources ( [[#Kremser--2016|Kremser et al., 2016]] ). New estimates of SO <sub>2</sub> emissions from explosive eruptions have been derived from satellite (beginning in 1979) and in situ measurements ( [[#Höpfner--2015|Höpfner et al., 2015]] ; [[#Carn--2016|Carn et al., 2016]] ; [[#Neely%20III--2016|Neely III and Schmidt, 2016]] ; [[#Brühl--2018|Brühl, 2018]] ). Satellite observations of aerosol extinction after recent eruptions have uncertainties of about 15–25% ( [[#Vernier--2011|Vernier et al., 2011]] ; [[#Bourassa--2012|Bourassa et al., 2012]] ). Additional uncertainties occur when gaps in the satellite records are filled by complementary observations or using statistical methods ( [[#Thomason--2018|Thomason et al., 2018]] ). Merged datasets ( [[#Thomason--2018|Thomason et al., 2018]] ) and sparse ground-based measurements ( [[#Stothers--1997|Stothers, 1997]] ) allow for volcanic forcing estimates back to 1850. In contrast to the CMIP5 historical volcanic forcing datasets ( [[#Ammann--2003|Ammann et al., 2003]] ), updated time series (Figure 2.2d; [[#Luo--2018|Luo, 2018]] ) feature a more comprehensive set of optical properties including latitude-, height- and wavelength-dependent aerosol extinction, single scattering albedo and asymmetry parameters. A series of small-to-moderate eruptions since 2000 resulted in perturbations in SAOD of 0.004–0.006 ( [[#Andersson--2015|Andersson et al., 2015]] ; [[#Schmidt--2018|Schmidt et al., 2018]] ).

To conclude, strong individual volcanic eruptions cause multi-annual variations in radiative forcing. However, the average magnitude and variability of SAOD and its associated volcanic aerosol forcing since 1900 are not unusual in the context of at least the past 2.5 kyr ( ''medium confidence'' ).

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<span id="well-mixed-greenhouse-gases-wmghgs"></span>
=== 2.2.3 Well-mixed Greenhouse Gases (WMGHGs) ===

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Well-mixed greenhouse gases generally have lifetimes of more than several years. The AR5 assigned ''medium confidence'' to the values of atmospheric CO <sub>2</sub> concentrations (mixing ratios) during the warm geological periods of the early Eocene and Pliocene. It concluded with ''very high confidence'' that, by 2011, the mixing ratios of CO <sub>2</sub> , CH <sub>4</sub> , and N <sub>2</sub> O in the atmosphere exceeded the range derived from ice cores for the previous 800 kyr, and that the observed rates of increase of the greenhouse gases were unprecedented on centennial timescales over at least the past 22 kyr. It reported that over 2005–2011 atmospheric burdens of CO <sub>2</sub> , CH <sub>4</sub> , and N <sub>2</sub> O increased, with 2011 levels of 390.5 parts per million (ppm), 1803.2 parts per billion (ppb) and 324.2 ppb, respectively. Increases of CO <sub>2</sub> and N <sub>2</sub> O over 2005–2011 were comparable to those over 1996–2005, while CH <sub>4</sub> resumed increasing in 2007, after remaining nearly constant over 1999–2006. A comprehensive process-based assessment of changes in CO <sub>2</sub> , CH <sub>4</sub> , and N <sub>2</sub> O is undertaken in Chapter 5.

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==== 2.2.3.1 CO <sub>2</sub> During 450 Ma to 800 ka ====

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Isotopes from continental and marine sediments using improved analytical techniques and sampling resolution have reinforced the understanding of long-term changes in atmospheric CO <sub>2</sub> during the past 450 Myr (Table 2.1 and Figure 2.3). In particular, for the last 60 Myr, sampling resolution and accuracy of the boron isotope proxy in ocean sediments has improved ( [[#Penman--2014|Penman et al., 2014]] ; [[#Anagnostou--2016|Anagnostou et al., 2016]] , 2020; [[#Chalk--2017|Chalk et al., 2017]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Babila--2018|Babila et al., 2018]] ; [[#Dyez--2018|Dyez et al., 2018]] ; [[#Raitzsch--2018|Raitzsch et al., 2018]] ; [[#Sosdian--2018|Sosdian et al., 2018]] ; [[#Henehan--2019|Henehan et al., 2019]] , 2020; [[#de%20la%20Vega--2020|de la Vega et al., 2020]] ; [[#Harper--2020|Harper et al., 2020]] ), the understanding of the alkenone CO <sub>2</sub> proxy has increased (e.g., [[#Badger--2019|Badger et al., 2019]] ; [[#Stoll--2019|Stoll et al., 2019]] ; Y. [[#Zhang--2019|]] [[#Zhang--2019|Zhang et al., 2019]] ; [[#Zhang--2020|Zhang et al., 2020]] ; [[#Rae--2021|Rae et al., 2021]] ) and new phytoplankton proxies have been developed and applied (e.g., [[#Witkowski--2018|Witkowski et al., 2018]] ). Understanding of the boron isotope CO <sub>2</sub> proxy has improved since AR5 with studies showing very good agreement between boron-CO <sub>2</sub> estimates and co-existing ice core CO <sub>2</sub> ( [[#Hönisch--2005|Hönisch and Hemming, 2005]] ; [[#Foster--2008|Foster, 2008]] ; [[#Henehan--2013|Henehan et al., 2013]] ; [[#Chalk--2017|Chalk et al., 2017]] ; [[#Raitzsch--2018|Raitzsch et al., 2018]] ; see Figure 2.3c). Such independent validation has proven difficult to achieve with the other available CO <sub>2</sub> proxies (e.g., [[#Badger--2019|Badger et al., 2019]] ; [[#Da--2019|Da et al., 2019]] ; [[#Stoll--2019|Stoll et al., 2019]] ; Y. [[#Zhang--2019|]] [[#Zhang--2019|Zhang et al., 2019]] ). Remaining uncertainties in these ocean sediment based proxies ( [[#Hollis--2019|Hollis et al., 2019]] ) partly limit the applicability of the alkenone δ <sup>13</sup> C and boron δ <sup>11</sup> B proxies beyond the Cenozoic, although new records are emerging, for example, [[#Jurikova--2020|Jurikova et al. (2020)]] . CO <sub>2</sub> estimates from the terrestrial CO <sub>2</sub> proxies, such as stomatal density in fossil plants and δ <sup>13</sup> C of palaeosol carbonates, are available for much of the last 420 Myr. Given the low sampling density, relatively large CO <sub>2</sub> uncertainty, and high age uncertainty (relative to marine sediments) of the terrestrial proxies, preference here is given to the marine based proxies (and boron in particular) where possible.

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'''Table 2.1 |''' '''Concentration (mixing ratios) and, where applicable, century time-scale rate of change of atmospheric CO''' <sub>2</sub> '''based on multiple datasets for target paleoclimate reference (Cross-Chapter Box 2.1, and Figure 2.34) and selected other periods.''' Modern data are from [[#2.2.3.3|Section 2.2.3.3]] and Annex III. ‘AR6’ denotes best estimates assessed in this report and propagated to Figure 2.34. Units for the rate of change are given only for centennial periods characterized by rapid changes. ''confidence'' levels are ''very high'' for instrumentally derived concentrations, ''high'' for values derived from air in glacier ice (back to LIG), ''medium'' for values supported by multiple proxy types (MPWP, EECO), and ''low'' for values from a single sedimentary proxy type (PETM). ‘ ''→'' ’ indicates transition from the beginning to the end of the time interval. Uncertainties for Modern are based on 2019 estimates. Last Millennium rate of range shows lowest and highest values attained during this period; LDT shows highest rate of change. N/A indicates that values are not available. See chapter data table for bibliographic citation and auxiliary information for each dataset (Table 2.SM.1).

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

| '''CO''' <sub>2</sub> '''Concentration (ppm) and Dataset Details'''

| '''Rate of Change (ppm per century)'''

|-
| Modern

(1995–2014)

| 359.6 to 360.4 ''→'' 396.7 to 397.5 (AR6)

| 192.3 to 198.3 <sup>a</sup> (AR6)

|-
| Last 100 years (1919–2019)

| 302.8 to 306.0 ''→'' 409.5 to 410.3 (AR6)

| 103.9 to 107.1 (AR6)

|-
| Approximate pre-industrial baseline (1850–1900; see Cross-Chapter Box 1.2)

| 283.4 to 287.6 ''→'' 294.8 to 298.0 (AR6);

284.3 <sup>b</sup> ''→'' 295.7 <sup>b</sup> (CMIP6)

| 16.5 to 27.1 <sup>a</sup> (AR6)

22.8 <sup>b,a</sup> (CMIP6)

|-
| Last millennium (1000–1750)

| 278.0 to 285.0 (AR6; average of WAIS Divide, Law Dome and EDML core data)

| –6.9 ~ 4.7 <sup>b</sup> (Law Dome); –1.9 ~ 3.2 <sup>b</sup> (EDML); –5.2 ~ 4.2 <sup>b</sup> (WAIS Divide)

|-
| MH

| 260.1 to 268.1 (Dome C; CMIP6)

| N/A

|-
| LDT

| 193.2 <sup>b</sup> ''→'' 271.2 <sup>b</sup> (AR6); 195.2 <sup>b</sup> ''→'' 265.3 <sup>b</sup> (Dome C); 191.2 <sup>b</sup> → 277.0 <sup>b</sup> (WAIS Divide)

| 9.6 <sup>b</sup> (WAIS Divide);

7.1 <sup>b</sup> (Dome C)

|-
| LGM

| 188.4 to 194.2 (AR6); 190.5 to 200.1 (WAIS Divide); 186.8 to 202.0 (Byrd); 184.9 to 193.1 (Dome C); 180.5 to 192.7 (Siple Dome); 190 <sup>b</sup> (PMIP6); 174.2 to 205.8 ( δ <sup>11</sup> B proxy)

| N/A

|-
| LIG

| 265.9 to 281.5 (AR6); 259.4 to 283.8 (Vostok); 266.2 to 285.4 (Dome C); 275 <sup>b</sup> (PMIP4) 282.2 to 305.8 ( δ <sup>11</sup> B proxy)

| N/A

|-
| MPWP (KM5c)

| 360 to 420 (AR6)

| N/A

|-
| EECO

| 1150 to 2500 (AR6)

| N/A

|-
| PETM

| 800 to 1000 ''→'' 1400 to 3150 (AR6)

| 4 to 42 (AR6)

|}

<sup>a</sup> Centennial rate of change estimated by extrapolation of data from a shorter time period. The values (x to y) represent ''very likely'' ranges (90% CIs).

<sup>b</sup> Data uncertainty is not estimated.

Levels were close to 1750 values during at least one prolonged interval during the Carboniferous and Permian (350–252 Ma). During the Triassic (251.9–201.3 Ma), atmospheric CO <sub>2</sub> mixing ratios reached a maximum of between 2000–5000 ppm (200–220 Ma). During the PETM (56 Ma) CO <sub>2</sub> rapidly rose from about 900 ppm to about 2000 ppm (Table 2.1; [[#Schubert--2013|Schubert and Jahren, 2013]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Anagnostou--2020|Anagnostou et al., 2020]] ) in 3–20 kyr ( [[#Zeebe--2016|Zeebe et al., 2016]] ; [[#Gutjahr--2017|Gutjahr et al., 2017]] ; [[#Turner--2018|Turner, 2018]] ). Estimated multi-millennial rates of CO <sub>2</sub> accumulation during this event range from 0.3–1.5 PgC yr <sup>–1</sup> ( [[#Gingerich--2019|Gingerich, 2019]] ), at least 4–5 times lower than current centennial rates (Section 5.3.1.1). Based on boron and carbon isotope data, supported by other proxies ( [[#Hollis--2019|Hollis et al., 2019]] ), atmospheric CO <sub>2</sub> during the EECO (50 Ma) was between 1150 and 2500 ppm ( ''medium confidence'' ), and then gradually declined over the last 50 Myr at a long-term rate of about 16 ppm Myr <sup>–1</sup> (Figure 2.3). The last time the CO <sub>2</sub> mixing ratio was as high as 1000 ppm (the level reached by some high emissions scenarios by 2100; Annex III) was prior to the Eocene-Oligocene transition (33.5 Ma; Figure 2.3) that was associated with the first major advance of the AIS ( [[#Pearson--2009|Pearson et al., 2009]] ; [[#Pagani--2011|Pagani et al., 2011]] ; [[#Anagnostou--2016|Anagnostou et al., 2016]] ; [[#Witkowski--2018|Witkowski et al., 2018]] ; [[#Hollis--2019|Hollis et al., 2019]] ). The compilation of [[#Foster--2017|Foster et al. (2017)]] constrained CO <sub>2</sub> concentration to between 290 and 450 ppm during the MPWP, based primarily on the boron-isotope data reported by [[#Martínez-Botí--2015b|Martínez-Botí et al. (2015b)]] , consistent with the AR5 range of 300–450 ppm. A more recent high-resolution boron isotope-based study revealed that CO <sub>2</sub> cycled during the MPWP from about 330 to about 390 ppm on orbital timescales, with a mean of about 370 ppm ( [[#de%20la%20Vega--2020|de la Vega et al., 2020]] ). Although data from other proxy types (e.g., stomatal density or δ <sup>13</sup> C of alkenones) have too low resolution to resolve the orbital-related variability of CO <sub>2</sub> during this interval (e.g., [[#Kürschner--1996|Kürschner et al., 1996]] ; [[#Stoll--2019|Stoll et al., 2019]] ) there is general agreement among the different proxy types with the boron-derived mean (e.g., [[#Stoll--2019|Stoll et al., 2019]] ). High-resolution sampling (about 1 sample per 3 kyr) with the boron-isotope proxy indicates mean CO <sub>2</sub> mixing ratios for the Marine Isotope Stage KM5c interglacial were 360–420 ppm ( ''medium confidence'' ) ( [[#de%20la%20Vega--2020|de la Vega et al., 2020]] ).

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[[File:d037d36a9535e99702f37968c89d6d5c IPCC_AR6_WGI_Figure_2_3.png]]

'''Figure 2.3 | The evolution of atmospheric CO''' <sub>2</sub> '''through the last 450 million years (450 Myr).''' The periods covered are 0–450 Ma '''(a)''' , 0–58 Ma '''(b)''' , and 0–3500 ka '''(c)''' , reconstructed from continental rock, marine sediment and ice core records. Note different time scales and axes ranges in panels (a), (b) and (c). Dark and light green bands in (a) are uncertainty envelopes at 68% and 95% uncertainty, respectively. 100 ppm in each panel is shown by the marker in the lower right-hand corner to aid comparison between panels. In panel (b) and (c) the major paleoclimate reference periods (CCB2.1) have been labelled, and in addition: MPT (Mid Pleistocene Transition), MCO (Miocene Climatic Optimum). Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Following the MPWP, the atmospheric CO <sub>2</sub> mixing ratio generally decreased at a rate of about 30 ppm Myr <sup>–1</sup> . It is ''very likely'' that CO <sub>2</sub> levels as high as the present were not experienced in the last 2 Myr ( [[#Hönisch--2009|Hönisch et al., 2009]] ; [[#Bartoli--2011|Bartoli et al., 2011]] ; [[#Martínez-Botí--2015a|Martínez-Botí et al., 2015a]] ; [[#Chalk--2017|Chalk et al., 2017]] ; [[#Dyez--2018|Dyez et al., 2018]] ; [[#Da--2019|Da et al., 2019]] ; [[#Stoll--2019|Stoll et al., 2019]] ). Related to the shift of glacial-interglacial cycle frequency from 40 to 100 kyr at 0.8–1.2 Ma, there was a decrease of glacial-period CO <sub>2</sub> ( [[#Chalk--2017|Chalk et al., 2017]] ; [[#Dyez--2018|Dyez et al., 2018]] ). These boron isotope-based CO <sub>2</sub> results agree with available records based on ancient ice exposed near the surface of the AIS ( [[#Yan--2019|Yan et al., 2019]] ), however, direct comparison is limited due to a lack of ancient ice cores with sufficiently continuous stratigraphy ( [[#Higgins--2015|Higgins et al., 2015]] ; [[#Brook--2018|Brook and Buizert, 2018]] ).

To conclude, there is ''high confidence'' that average EECO and MPWP (KM5c) CO <sub>2</sub> concentrations were higher than those preceding industrialization at 1150–2500 ppm and 360–420 ppm, respectively. Although there is some uncertainty due to the non-continuous nature of marine sediment records, the last time atmospheric CO <sub>2</sub> mixing ratio was as high as present was ''very likely'' more than 2 Ma.

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==== 2.2.3.2 Glacial–Interglacial WMGHG Fluctuations from 800 Ka ====

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Since AR5, the number of ice cores for the last 800 kyr has increased and their temporal resolution has improved (Figure 2.4), especially for the last 60 kyr and when combined with analyses of firn air, leading to improved quantification of greenhouse gas concentrations prior to the mid-20th century.

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[[File:2f2d17b5b6508d0f035037b4a5fb2f75 IPCC_AR6_WGI_Figure_2_4.png]]

'''Figure 2.4 | Atmospheric well-mixed greenhouse gas (WMGHG) concentrations''' '''from ice cores. (a)''' Records during the last 800 kyr with the Last Glacial Maximum (LGM) to Holocene transition as inset. '''(b)''' Multiple high-resolution records over the CE. The horizontal black bars in panel (a) inset indicate LGM and Last Deglacial Termination (LDT) respectively. The red and blue lines in (b) are 100-year running averages for CO <sub>2</sub> and N <sub>2</sub> O concentrations, respectively. The numbers with vertical arrows in (b) are instrumentally measured concentrations in 2019. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

<div id="2.2.3.2.1" class="h4-container"></div>

<span id="carbon-dioxide-co-2"></span>
===== 2.2.3.2.1 Carbon dioxide (CO <sub>2</sub> ) =====

<div id="h4-1-siblings" class="h4-siblings"></div>

Records of CO <sub>2</sub> from the AIS formed during the last glacial period and the LDT show century-scale fluctuations of up to 9.6 ppm ( [[#Ahn--2012|Ahn et al., 2012]] ; [[#Ahn--2014|Ahn and Brook, 2014]] ; [[#Marcott--2014|Marcott et al., 2014]] ; [[#Bauska--2015|Bauska et al., 2015]] ; [[#Rubino--2019|Rubino et al., 2019]] ). Although these rates are an order of magnitude lower than those directly observed over 1919–2019 CE ( [[#2.2.3.3.1|Section 2.2.3.3.1]] ), they provide information on non-linear responses of climate-biogeochemical feedbacks (Section 5.1.2). Multiple records for 0–1850 CE show CO <sub>2</sub> mixing ratios of 274–285 ppm. Offsets among ice core records are about 1%, but the long-term trends agree well and show coherent multi-centennial variations of about 10 ppm ( [[#Ahn--2012|Ahn et al., 2012]] ; [[#Bauska--2015|Bauska et al., 2015]] ; [[#Rubino--2019|Rubino et al., 2019]] ). Multiple records show CO <sub>2</sub> concentrations of 278.3 ± 2.9 ppm in 1750 and 285.5 ± 2.1 ppm in 1850 ( [[#Siegenthaler--2005|Siegenthaler et al., 2005]] ; [[#MacFarling%20Meure--2006|MacFarling Meure et al., 2006]] ; [[#Ahn--2012|Ahn et al., 2012]] ; [[#Bauska--2015|Bauska et al., 2015]] ). CO <sub>2</sub> concentration increased by 5.0 ± 0.8 ppm during 970–1130 CE, followed by a decrease of 4.6 ± 1.7 ppm during 1580–1700 CE. The greatest rate of change over the CE prior to 1750 is observed at about 1600 CE, and ranges from –6.9 to +4.7 ppm per century in multiple high-resolution ice core records ( [[#Siegenthaler--2005|Siegenthaler et al., 2005]] ; [[#MacFarling%20Meure--2006|MacFarling Meure et al., 2006]] ; [[#Ahn--2012|Ahn et al., 2012]] ; [[#Bauska--2015|Bauska et al., 2015]] ; [[#Rubino--2019|Rubino et al., 2019]] ). Although ice core records present low-pass filtered time series due to gas diffusion and gradual bubble close-off in the snow layer over the ice sheet ( [[#Fourteau--2020|Fourteau et al., 2020]] ), the rate of increase since 1850 CE (about 125 ppm increase over about 170 years) is far greater than implied for any 170-year period by ice core records that cover the last 800 ka ( ''very high confidence'' ).

<div id="2.2.3.2.2" class="h4-container"></div>

<span id="methane-ch-4"></span>
===== 2.2.3.2.2 Methane (CH <sub>4</sub> ) =====

<div id="h4-2-siblings" class="h4-siblings"></div>

CH <sub>4</sub> concentrations over the past 110 kyr are higher in the Northern Hemisphere (NH) than in the Southern Hemisphere (SH), but closely correlated on centennial and millennial timescales ( [[#Buizert--2015|Buizert et al., 2015]] ). On glacial to interglacial cycles, approximately 450 ppb oscillations in CH <sub>4</sub> concentrations have occurred ( [[#Loulergue--2008|Loulergue et al., 2008]] ). On millennial timescales, most rapid climate changes observed in Greenland and other regions are coincident with rapid CH <sub>4</sub> changes ( [[#Buizert--2015|Buizert et al., 2015]] ; [[#Rhodes--2015|Rhodes et al., 2015]] , 2017). The variability of CH <sub>4</sub> on centennial timescales during the early Holocene does not significantly differ from that of the late Holocene prior to about 1850 ( [[#Rhodes--2013|Rhodes et al., 2013]] ; [[#Yang--2017|Yang et al., 2017]] ). The LGM concentration was 390.5 ± 6.0 ppb ( [[#Kageyama--2017|Kageyama et al., 2017]] ). The global mean concentrations during 0–1850 CE varied between 625 and 807 ppb. High-resolution ice core records from Antarctica and Greenland exhibit the same trends with an inter-polar difference of 36–47 ppb ( [[#Sapart--2012|Sapart et al., 2012]] ; L. [[#Mitchell--2013|]] [[#Mitchell--2013|Mitchell et al., 2013]] ). There is a long-term positive trend of about 0.5 ppb per decade during the CE until 1750 CE. The most rapid CH <sub>4</sub> changes prior to industrialization were as large as 30–50 ppb on multi-decadal timescales. Global mean CH <sub>4</sub> concentrations estimated from Antarctic and Greenland ice cores are 729.2 ± 9.4 ppb in 1750 and 807.6 ± 13.8 ppb in 1850 (L. [[#Mitchell--2013|]] [[#Mitchell--2013|Mitchell et al., 2013]] ).

<div id="2.2.3.2.3" class="h4-container"></div>

<span id="nitrous-oxide-n-2-o"></span>
===== 2.2.3.2.3 Nitrous oxide (N <sub>2</sub> O) =====

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New records show that N <sub>2</sub> O concentration changes are associated with glacial-interglacial transitions ( [[#Schilt--2014|Schilt et al., 2014]] ). The most rapid change during the last glacial termination is a 30 ppb increase in a 200-year period, which is an order of magnitude smaller than the modern rate ( [[#2.2.3.3|Section 2.2.3.3]] ). During the LGM, N <sub>2</sub> O was 208.5 ± 7.7 ppb ( [[#Kageyama--2017|Kageyama et al., 2017]] ). Over the Holocene the lowest value was 257 ± 6.6 ppb during 6–8 ka, but millennial variation is not clearly detectable due to analytical uncertainty and insufficient ice core quality ( [[#Flückiger--2002|Flückiger et al., 2002]] ; [[#Schilt--2010|Schilt et al., 2010]] ). Recently acquired high-resolution records from Greenland and Antarctica for the last 2 kyr consistently show multi-centennial variations of about 5–10 ppb (Figure 2.4), although the magnitudes vary over time ( [[#Ryu--2020|Ryu et al., 2020]] ). Three high temporal resolution records exhibit a short-term minimum at about 600 CE of 261 ± 4 ppb ( [[#MacFarling%20Meure--2006|MacFarling Meure et al., 2006]] ; [[#Ryu--2020|Ryu et al., 2020]] ). It is ''very likely'' that industrial N <sub>2</sub> O increase started before 1900 CE ( [[#Machida--1995|Machida et al., 1995]] ; [[#Sowers--2001|Sowers, 2001]] ; [[#MacFarling%20Meure--2006|MacFarling Meure et al., 2006]] ; [[#Ryu--2020|Ryu et al., 2020]] ). Multiple ice cores show N <sub>2</sub> O concentrations of 270.1 ± 6.0 ppb in 1750 and 272.1 ± 5.7 ppb in 1850 ( [[#Machida--1995|Machida et al., 1995]] ; [[#Flückiger--1999|Flückiger et al., 1999]] ; [[#Sowers--2001|Sowers, 2001]] ; [[#Rubino--2019|Rubino et al., 2019]] ; [[#Ryu--2020|Ryu et al., 2020]] ).

<div id="2.2.3.3" class="h3-container"></div>

<span id="modern-measurements-of-wmghgs"></span>

==== 2.2.3.3 Modern Measurements of WMGHGs ====

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In this section and for calculation of ERF, surface global averages are determined from measurements representative of the well-mixed lower troposphere. Global averages that include sites subject to significant anthropogenic activities or influenced by strong regional biospheric emissions are typically larger than those from remote sites, and require weighting accordingly (Table 2.2). This section focusses on global mean mixing ratios estimated from networks with global spatial coverage, and updated from the CMIP6 historical dataset ( [[#Meinshausen--2017|Meinshausen et al., 2017]] ) for periods prior to the existence of global networks.

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Table 2.2 '''|''' '''Atmospheric global annual mean mixing ratios (dry-air mole fraction) for well-mixed greenhouse gases.''' The table provides observed values for 2011 and 2019, and relative changes since 2011, for selected well-mixed, radiatively important gases (ERF &gt;0.001 W m <sup>–2</sup> ), estimated from various measurement networks or compilations. Units are parts per million (ppm) for CO <sub>2</sub> , parts per billion (ppb) for CH <sub>4</sub> and N <sub>2</sub> O, parts per trillion (ppt) for all other gases. Time series since 1750, data for additional gases, references, and network information can be found in [[IPCC:Wg1:Chapter:Annex-iii|Annex III]] and the corresponding electronic supplement. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

{| class="wikitable"
|-
| '''Species'''

| '''Lifetime,'''

'''AR6, ERF'''

| '''2011'''

| '''2019'''

| '''Change'''

| '''Network'''

| rowspan="35"|
| '''Species'''

| '''Lifetime,'''

'''AR6, ERF'''

| '''2011'''

| '''2019'''

| '''Change'''

| '''Network'''

|-
| rowspan="5"| '''CO''' <sub>2</sub>

| \#

| 390.5

| 409.9 (0.17)

| 5.0%

| NOAA* <sup>a</sup>

| rowspan="3"| '''HCFC-22'''

| 11.9

| 212.6

| 246.8 (0.5)

| 16.1%

| NOAA*

|-
| 409.9 (0.4)

| 389.7

| 409.5 (0.37)

| 5.1%

| SIO

| 246.8 (0.6)

| 213.7

| 246.7 (0.4)

| 15.5%

| AGAGE*

|-
| 2.156

| 390.2

| 409.6 (0.31)

| 5.0%

| CSIRO

| 0.053

| 209.0

| 244.1 (3.0)

| 22.0%

| UCI

|-
|
| 390.9

| 410.5 (0.30)

| 5.0%

| WMO

| rowspan="3"| '''HCFC-141b'''

| 9.4

| 21.3

| 24.4 (0.1)

| 14.4%

| NOAA*

|-
|
| 390.9

|
| CMIP6

| 24.4 (0.3)

| 21.4

| 24.3 (0.1)

| 13.7%

| AGAGE*

|-
| rowspan="6"| '''CH''' <sub>4</sub>

| 9.1–11.8

| 1803.1

| 1866.6 (1.0)

| 3.5%

| NOAA*

| 0.004

| 20.8

| 26.0 (0.3)

| 25.0%

| UCI

|-
| 1866.3 (3.3)

| 1803.6

| 1866.1 (2.0)

| 3.5%

| AGAGE*

| rowspan="3"| '''HCFC-142b'''

| 18

| 20.9

| 22.0 (0.1)

| 5.3%

| NOAA*

|-
| 0.544

| 1791.8

| 1860.8 (3.5)

| 3.9%

| UCI

| 22.3 (0.4)

| 21.5

| 22.5 (0.1)

| 5.0%

| AGAGE*

|-
|
| 1802.3

| 1862.5 (2.4)

| 3.3%

| CSIRO

| 0.004

| 21.0

| 22.8 (0.2)

| 8.6%

| UCI

|-
|
| 1813

| 1877 (3)

| 3.5%

| WMO

| rowspan="3"| '''HFC-134a'''

| 14

| 62.7

| 107.8 (0.4)

| 72%

| NOAA*

|-
|
| 1813.1

|
| CMIP6

| 107.6 (1.0)

| 62.8

| 107.4 (0.2)

| 71%

| AGAGE*

|-
| rowspan="5"| '''N''' <sub>2</sub> '''O'''

| 116–109

| 324.2

| 331.9 (0.2)

| 2.4%

| NOAA*

| 0.018

| 63.4

| 107.6 (1.7)

| 70%

| UCI

|-
| 332.1 (0.4)

| 324.7

| 332.3 (0.1)

| 2.4%

| AGAGE*

| rowspan="3"| '''HFC-125'''

| 30

| 10.1

| 29.1 (0.3)

| 187%

| NOAA*

|-
| 0.208

| 324.0

| 331.6 (0.3)

| 2.3%

| CSIRO

| 29.4 (0.6)

| 10.4

| 29.7 (0.1)

| 186%

| AGAGE*

|-
|
| 324.3

| 332.0 (0.2)

| 2.4%

| WMO

| 0.007

|
|-
|
| 324.2

|
| CMIP6

| rowspan="3"| '''HFC-23'''

| 228

| 24.1

| 32.4 (0.1)

| 35%

| AGAGE*

|-
| rowspan="3"| '''CFC-12'''

| 102

| 526.9

| 501.5 (0.3)

| –4.8%

| NOAA*

| 32.4 (0.1)

|
|-
| 503.1 (3.2)

| 529.6

| 504.6 (0.2)

| –4.7%

| AGAGE*

| 0.006

|
|-
| 0.180

| 525.3

| 508.4 (2.5)

| –3.2%

| UCI

| rowspan="3"| '''HFC-143a'''

| 51

| 11.9

| 23.8 (0.1)

| 100%

| NOAA*

|-
| rowspan="3"| '''CFC-11'''

| 52

| 237.2

| 226.5 (0.2)

| –4.5%

| NOAA*

| 24.0 (0.4)

| 12.1

| 24.2 (0.1)

| 100%

| AGAGE*

|-
| 226.2 (1.1)

| 237.4

| 225.9 (0.1)

| –4.8%

| AGAGE*

| 0.004

|
|-
| 0.066

| 237.9

| 224.9 (1.3)

| –5.5%

| UCI

| rowspan="3"| '''HFC-32'''

| 5.4

| 4.27

| 19.2 (0.3)

| 350%

| NOAA*

|-
| rowspan="3"| '''CFC-113'''

| 93

| 74.5

| 69.7 (0.1)

| –6.4%

| NOAA*

| 20.0 (1.4)

| 5.15

| 20.8 (0.2)

| 304%

| AGAGE*

|-
| 69.8 (0.3)

| 74.6

| 69.9 (0.1)

| –6.3%

| AGAGE*

| 0.002

|
|-
| 0.021

| 74.9

| 70.0 (0.5)

| –6.5%

| UCI

| rowspan="3"| '''CF''' <sub>4</sub>

| 50,000

| 79.0

| 85.5 (0.1)

| 8.2%

| AGAGE*

|-
| rowspan="3"| '''CFC-114'''

| 189

| 16.36

| 16.28 (0.03)

| –0.5%

| AGAGE*

| 85.5 (0.2)

|
|-
| 16.0 (0.05)

|
| 0.005

|
|-
| 0.005

|
| rowspan="3"| '''C''' <sub>2</sub> '''F''' <sub>6</sub>

| 10,000

| 4.17

| 4.85 (0.01)

| 16.3%

| AGAGE*

|-
| rowspan="3"| '''CFC-115'''

| 540

| 8.39

| 8.67 (0.02)

| 3.3%

| AGAGE*

| 4.85 (0.1)

|
|-
| 8.67 (0.02)

|
| 0.001

|
|-
| 0.002

|
| rowspan="3"| '''SF''' <sub>6</sub> <sub></sub>

| About 1000

| 7.32

| 9.96 (0.02)

| 36.1%

| NOAA*

|-
| rowspan="3"| '''CCl''' <sub>4</sub>

| 32

| 86.9

| 78.4 (0.1)

| –9.8%

| NOAA*

| 9.95 (0.01)

| 7.28

| 9.94 (0.02)

| 36.5%

| AGAGE*

|-
| 77.9 (0.7)

| 85.3

| 77.3 (0.1)

| –9.4%

| AGAGE*

| 0.006

|
|-
| 0.013

| 87.8

| 77.7 (0.7)

| –11.5%

| UCI

|
|}

AGAGE: Advanced Global Atmospheric Gases Experiment; SIO: Scripps Institution of Oceanography; NOAA: National Oceanic and Atmospheric Administration, Global Monitoring Laboratory; UCI: University of California, Irvine; CSIRO: Commonwealth Scientific and Industrial Research Organization, Aspendale, Australia; WMO: World Meteorological Organization, Global Atmosphere Watch, CMIP6 (Climate Model Intercomparison Project Phase 6). Mixing ratios denoted by AR6 are representative of the remote, unpolluted troposphere, derived from one or more measurement networks (denoted by *). Minor differences between 2011 values reported here and in the previous Assessment Report (AR5) are due to updates in calibration and data processing. ERF in 2019 is taken from Table 7.5, and the difference with the AR5 assessment reflects updates in the estimates of AR6 global mixing ratios and updated radiative calculations. Uncertainties, in parenthesis, are estimated at 90% confidence interval. Networks use different methods to estimate uncertainties. Some uncertainties have been rounded up to be consistent with the number of decimal places shown. Lifetime is reported in years: # indicates multiple lifetimes for CO <sub>2</sub> . For CH <sub>4</sub> and N <sub>2</sub> O the two values represent total atmospheric lifetime and perturbation lifetime.

<div id="2.2.3.3.1" class="h4-container"></div>

<span id="carbon-dioxide-co-2-1"></span>
===== 2.2.3.3.1 Carbon dioxide (CO <sub>2</sub> ) =====

<div id="h4-4-siblings" class="h4-siblings"></div>

There has been a positive trend in globally averaged surface CO <sub>2</sub> mixing ratios since 1958 (Figure 2.5a), that reflects the imbalance of sources and sinks (Section 5.2). The growth rate has increased overall since the 1960s (Figure 2.5a inset), while annual growth rates have varied substantially, for example, reaching a peak during the strong El Niño events of 1997–1998 and 2015–2016 ( [[#Bastos--2013|Bastos et al., 2013]] ; [[#Betts--2016|Betts et al., 2016]] ). The average annual CO <sub>2</sub> increase from 2000 through 2011 was 2.0 ppm yr <sup>–1</sup> (standard deviation 0.3 ppm yr <sup>–1</sup> ), similar to what was reported in AR5. From 2011 through 2019 it was 2.4 ppm yr <sup>–</sup> <sup>1</sup> (standard deviation 0.5 ppm yr <sup>–1</sup> ), which is higher than that of any comparable time period since global measurements began. Global networks consistently show that the globally averaged annual mean CO <sub>2</sub> has increased by 5.0% since 2011, reaching 409.9 ± 0.4 ppm in 2019 (NOAA measurements). Further assessment of changing seasonality is undertaken in [[#2.3.4.1|Section 2.3.4.1]] .

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[[File:3e2fe664c93efaa7c3c0198d50490cdd IPCC_AR6_WGI_Figure_2_5.png]]

'''Figure 2.5 |''' '''Globally averaged dry-air mole fractions of greenhouse gases. (a)''' CO <sub>2</sub> from SIO, CSIRO, and NOAA/GML '''(b)''' CH <sub>4</sub> from NOAA, AGAGE, CSIRO, and UCI; and '''(c)''' N <sub>2</sub> O from NOAA, AGAGE, and CSIRO (Table 2.2). Growth rates, calculated as the time derivative of the global means after removing seasonal cycle are shown as inset figures. Note that the CO <sub>2</sub> series is 1958–2019 whereas CH <sub>4</sub> , and N <sub>2</sub> O are 1979–2019. Units are parts per million (ppm) or parts per billion (ppb). Further details on data are in Annex III, and on data sources and processing are available in the chapter data table (Table 2.SM.1).

<div id="2.2.3.3.2" class="h4-container"></div>

<span id="methane-ch-4-1"></span>
===== 2.2.3.3.2 Methane (CH <sub>4</sub> ) =====

<div id="h4-5-siblings" class="h4-siblings"></div>

The globally averaged surface mixing ratio of CH <sub>4</sub> in 2019 was 1866.3 ± 3.3 ppb, which is 3.5% higher than 2011, while observed increases from various networks range from 3.3–3.9% (Table 2.2 and Figure 2.5b). There are marked growth rate changes over the period of direct observations, with a decreasing rate from the late-1970s through the late-1990s, very little change in concentrations from 1999–2006, and resumed increases since 2006. Atmospheric CH <sub>4</sub> fluctuations result from complex variations of sources and sinks. A detailed discussion of recent methane trends and our understanding of their causes is presented in Cross-Chapter Box 5.2.

<div id="2.2.3.3.3" class="h4-container"></div>

<span id="nitrous-oxide-n-2-o-1"></span>
===== 2.2.3.3.3 Nitrous oxide (N <sub>2</sub> O) =====

<div id="h4-6-siblings" class="h4-siblings"></div>

The AR5 reported 324.2 ± 0.1 ppb for global surface annual mean N <sub>2</sub> O in 2011; since then, it has increased by 2.4% to 332.1 ± 0.4 ppb in 2019. Independent measurement networks agree well for both the global mean mixing ratio and relative change since 2011 (Table 2.2). Over 1995–2011, N <sub>2</sub> O increased at an average rate of 0.79 ± 0.05 ppb yr <sup>–1</sup> . The growth rate has been higher in recent years, amounting to 0.96 ± 0.05 ppb yr <sup>–</sup> <sup>1</sup> from 2012 to 2019 (Figure 2.5c and Section 5.2.3.5).

<div id="2.2.3.4" class="h3-container"></div>

<span id="summary-of-changes-in-wmghgs"></span>

==== 2.2.3.4 Summary of Changes in WMGHGs ====

<div id="h3-4-siblings" class="h3-siblings"></div>

In summary, CO <sub>2</sub> has fluctuated by at least 2000 ppm over the last 450 Myr ( ''medium confidence'' ). The last time CO <sub>2</sub> concentrations were similar to the present-day was over 2 Ma ( ''high confidence'' ). Further, it is certain that WMGHG mixing ratios prior to industrialization were lower than present-day levels and the growth rates of the WMGHGs from 1850 are unprecedented on centennial timescales in at least the last 800 kyr. During the glacial-interglacial climate cycles over the last 800 kyr, the concentration variations of the WMGHG were 50–100 ppm for CO <sub>2</sub> , 210–430 ppb for CH <sub>4</sub> and 60–90 ppb for N <sub>2</sub> O. Between 1750–2019 mixing ratios increased by 131.6 ± 2.9 ppm (47%), 1137 ± 10 ppb (156%), and 62 ± 6 ppb (23%), for CO <sub>2</sub> , CH <sub>4</sub> , and N <sub>2</sub> O, respectively ( ''very high confidence'' ). Since 2011 (AR5) mixing ratios of CO <sub>2</sub> , CH <sub>4</sub> , and N <sub>2</sub> O have further increased by 19 ppm, 63 ppb, and 7.7 ppb, reaching in 2019 levels of 409.9 (± 0.4) ppm, 1866.3 (± 3.3) ppb, and 332.1 (± 0.4) ppb, respectively. By 2019, the combined ERF (relative to 1750) of CO <sub>2</sub> , CH <sub>4</sub> and N <sub>2</sub> O was 2.9 ± 0.5 W m <sup>–2</sup> (Table 2.2; Section 7.3.2).

<div id="2.2.4" class="h2-container"></div>

<span id="halogenated-greenhouse-gases-cfcs-hcfcs-hfcs-pfcs-sf6-and-others"></span>
=== 2.2.4 Halogenated Greenhouse Gases (CFCs, HCFCs, HFCs, PFCs, SF6 and others) ===

<div id="h2-8-siblings" class="h2-siblings"></div>

This category includes ozone depleting substances (ODS), their replacements, and gases used industrially or produced as by-products. Some have natural sources (Section 6.2.2.4). The AR5 reported that atmospheric abundances of chlorofluorocarbons (CFCs) were decreasing in response to controls on production and consumption mandated by the Montreal Protocol on Substances that Deplete the Ozone Layer and its amendments. In contrast, abundances of both hydrochlorofluorocarbons (HCFCs, replacements for CFCs) and hydrofluorocarbons (HFCs, replacements for HCFCs) were increasing. Atmospheric abundances of perfluorocarbons (PFCs), SF <sub>6</sub> , and NF <sub>3</sub> were also increasing.

Further details on ODS and other minor greenhouse gases can be found in the Scientific Assessment of Ozone Depletion: 2018 ( [[#Engel--2018|Engel et al., 2018]] ; [[#Montzka--2018b|Montzka et al., 2018b]] ). Updated mixing ratios of the most radiatively important gases (ERF &gt;0.001 W m <sup>–2</sup> ) are reported in Table 2.2, and additional gases (ERF &lt;0.001 W m <sup>–2</sup> ) are shown in Annex III.

<div id="2.2.4.1" class="h3-container"></div>

<span id="chlorofluorocarbons-cfcs"></span>
==== 2.2.4.1 Chlorofluorocarbons (CFCs) ====

<div id="h3-5-siblings" class="h3-siblings"></div>

Atmospheric abundances of most CFCs have continued to decline since 2011 (AR5). The globally-averaged abundance of CFC-12 decreased by 25 ppt (4.8%) from 2011 to 2019, while CFC-11 decreased by about 11 ppt (4.7%) over the same period (Table 2.2 and Figure 2.6). Atmospheric abundances of some minor CFCs (CFC-13, CFC-115, CFC-113a) have increased since 2011 (Annex III), possibly related to use of HFCs ( [[#Laube--2014|Laube et al., 2014]] ). Overall, as of 2019 the ERF from CFCs has declined by 9 ± 0.5% from its maximum in 2000, and 4.7 ± 0.6% since 2011 (Table 7.5).

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[[File:d1f994b46cfc8ab9668df2eb9ef8c20c IPCC_AR6_WGI_Figure_2_6.png]]

'''Figure 2.6''' '''|''' '''Global mean atmospheric mixing ratios of select ozone-depleting''' '''substances and other greenhouse gases.''' Data shown are based on the CMIP6 historical dataset and data from NOAA and AGAGE global networks. PFCs include CF <sub>4</sub> , C <sub>2</sub> f <sub>6</sub> , and C <sub>3</sub> F <sub>8</sub> , and ''c'' -C <sub>4</sub> F <sub>8</sub> ; Halons include halon-1211, halon-1301, and halon-2402; other HFCs include HFC-23, HFC-32, HFC-125, HFC-143a, HFC-152a, HFC-227ea, HFC-236fa, HFC-245fa, and HFC-365mfc, and HFC-43-10mee. Note that the y-axis range is different for '''(a)''' , '''(b)''' and '''(c)''' and a 25 parts per trillion (ppt) yardstick is given next to each panel to aid interpretation. Further data are in [[IPCC:Wg1:Chapter:Annex-iii|Annex III]] and details on data sources and processing are available in the chapter data table (Table 2.SM.1).

While global reporting indicated that CFC-11 production had essentially ceased by 2010, and the atmospheric abundance of CFC-11 is still decreasing, emissions inferred from atmospheric observations began increasing in 2013–2014 and remained elevated for 5–6 years, suggesting renewed and unreported production ( [[#Montzka--2018a|Montzka et al., 2018a]] , 2021; [[#Rigby--2019|Rigby et al., 2019]] ; [[#Park--2021|Park et al., 2021]] ). The global lifetimes of several ozone-depleting substances have been updated (SPARC, 2013), in particular for CFC-11 from 45 to 52 years.

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<span id="hydrochlorofluorocarbons-hcfcs"></span>
==== 2.2.4.2 Hydrochlorofluorocarbons (HCFCs) ====

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The atmospheric abundances of the major HCFCs (HCFC-22, HCFC-141b, HCFC-142b), primarily used in refrigeration and foam blowing, are increasing, but rates of increase have slowed in recent years (Figure 2.6). Global mean mixing ratios (Table 2.2) showed good concordance at the time of AR5 for the period 2005–2011. For the period 2011–2019, the UCI network detected larger increases in HCFC-22, HCFC-141b, and HCFC-142b compared to the NOAA and AGAGE networks. Reasons for the discrepancy are presently unverified, but could be related to differences in sampling locations in the networks ( [[#Simpson--2012|Simpson et al., 2012]] ). Emissions of HCFC-22, derived from atmospheric data, have remained relatively stable since 2012, while those of HCFC-141b and HCFC-142b have declined ( [[#Engel--2018|Engel et al., 2018]] ). Minor HCFCs, HCFC-133a and HCFC-31, have been detected in the atmosphere (currently less than 1 ppt) and may be unintentional by-products of HFC production ( [[#Engel--2018|Engel et al., 2018]] ).

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<span id="hydrofluorocarbons-hfcs-perfluorocarbons-pfcs-sulphur-hexafluoride-sf-6-and-other-radiatively-important-halogenated-gases"></span>
==== 2.2.4.3 Hydrofluorocarbons (HFCs), Perfluorocarbons (PFCs), Sulphur Hexafluoride (SF <sub>6</sub> ) and Other Radiatively Important Halogenated Gases ====

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Hydrofluorocarbons (HFCs) are replacements for CFCs and HCFCs. The atmospheric abundances of many HFCs increased between 2011 and 2019. HFC-134a (mobile air conditioning, foam blowing, and domestic refrigerators) increased by 71% from 63 ppt in 2011 to 107.6 ppt in 2019 (Table 2.2). The UCI network detected a slightly smaller relative increase (53%). HFC-23, which is emitted as a by-product of HCFC-22 production, increased by 8.4 ppt (35%) over 2011–2019. HFC-32 used as a substitute for HCFC-22, increased at least by 300%, and HFC-143a and HFC-125 showed increases of 100% and 187%, respectively. While the ERF of HFC-245fa is currently &lt;0.001 W m <sup>–2</sup> , its atmospheric abundance doubled since 2011 to 3.1 ppt in 2019 (Annex III). In contrast, HFC-152a is showing signs of stable (steady-state) abundance.

Other radiatively important gases with predominantly anthropogenic sources also continue to increase in abundance. SF <sub>6</sub> , used in electrical distribution systems, magnesium production, and semi-conductor manufacturing, increased from 7.3 ppt in 2011 to 10.0 ppt in 2019 (+36%). Alternatives to SF <sub>6</sub> or SF <sub>6</sub> -free equipment for electrical systems have become available in recent years, but SF <sub>6</sub> is still widely in use in electrical switch gear ( [[#Simmonds--2020|Simmonds et al., 2020]] ). The global lifetime of SF <sub>6</sub> has been revised from 3200 years to about 1000 years ( [[#Kovács--2017|Kovács et al., 2017]] ; [[#Ray--2017|Ray et al., 2017]] ) with implications for climate emissions metrics (Section 7.6.2). NF <sub>3</sub> , which is used in the semi-conductor industry, increased 147% over the same period to 2.05 ppt in 2019. Its contribution to ERF remains small, however, at 0.0004 W m <sup>–2</sup> . The atmospheric abundance of SO <sub>2</sub> f <sub>2</sub> , which is used as a fumigant in place of ozone-depleting methyl bromide, reached 2.5 ppt in 2019, a 46% increase from 2011. Its ERF also remains small at 0.0005 W m <sup>–2</sup> .

The global abundance of CCl <sub>4</sub> continues to decline, down about 9.6% since 2011. Following a revision of the global lifetime from 26 to 32 years, and discovery of previously unknown sources (e.g., biproducts of industrial emissions), knowledge of the CCl <sub>4</sub> budget has improved. There is now better agreement between top-down emissions estimates (based on atmospheric measurements) and industry-based estimates ( [[#Engel--2018|Engel et al., 2018]] ). Halon-1211, mainly used for fire suppression, is also declining, and its ERF dropped below 0.001 W m <sup>–2</sup> in 2019. While CH <sub>2</sub> cl <sub>2</sub> has a short atmospheric lifetime (6 months), and is not well-mixed, its abundance is increasing and its ERF is approaching 0.001 W m <sup>–2</sup> .

Perfluorocarbons CF <sub>4</sub> and C <sub>2</sub> f <sub>6</sub> , which have exceedingly long global lifetimes, showed modest increases from 2011 to 2019. CF <sub>4</sub> , which has both natural and anthropogenic sources, increased 8.2% to 85.5 ppt, and C <sub>2</sub> f <sub>6</sub> increased 16.3% to 4.85 ppt. ''c'' <sup>–</sup> C <sub>4</sub> F <sub>8</sub> , which is used in the electronics industry and may also be generated during the production of polytetrafluoroethylene (PTFE, also known as ‘Teflon’) and other fluoropolymers ( [[#Mühle--2019|Mühle et al., 2019]] ), has increased 34% since 2011 to 1.75 ppt, although its ERF remains below 0.001 W m <sup>–2</sup> . Other PFCs, present at mixing ratios &lt;1 ppt, have also been quantified ( [[#Droste--2020|Droste et al., 2020]] ; see Annex III).

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[[File:c7b7cf9bf1bbed375a5977b87d763289 IPCC_AR6_WGI_Figure_2_37.png]]

'''Figure 2.37 |''' '''Indices of interannual climate variability from 1950–2019 based upon several sea surface temperature data products.''' Shown are the following indices from top to bottom: IOB mode, IOD, Niño3.4, AMM and AZM. All indices are based on area-averaged annual data (see Annex IV). Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

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<span id="summary-of-changes-in-halogenated-gases"></span>
==== 2.2.4.4 Summary of Changes in Halogenated Gases ====

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In summary, by 2019 the ERF of halogenated GHGs had increased by 3.5% since 2011, reflecting predominantly a decrease in the atmospheric mixing ratios of CFCs and an increase in their replacements. However, average annual ERF growth rates associated with halogenated gases since 2011 are a factor of seven lower than in the 1970s and 1980s. Direct radiative forcings from CFCs, HCFCs, HFCs, and other halogenated greenhouse gases were 0.28, 0.06, 0.04, and 0.03 W m <sup>–2</sup> respectively, totalling 0.41 ± 0.07 W m <sup>–2</sup> in 2019 (see Table 7.5).

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<span id="other-short-lived-gases"></span>
=== 2.2.5 Other Short-lived Gases ===

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<span id="stratospheric-water-vapour"></span>
==== 2.2.5.1 Stratospheric Water Vapour ====

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The AR5 assessed ''low confidence'' in stratospheric water vapour (SWV) trends based on substantial seasonal and interannual variability in satellite data from 1992 to 2011. The 1980–2010 record of balloon-borne frost point hygrometer measurements over Boulder, Colorado (40°N), showed an average net increase of 1.0 ± 0.2 ppm (27 ± 6%) in the 16–26 km layer.

Since AR5, bias-adjusted spatially comprehensive SWV measurements by different satellite sensors were merged to form continuous records ( [[#Hegglin--2014|Hegglin et al., 2014]] ; [[#Froidevaux--2015|Froidevaux et al., 2015]] ; [[#Davis--2016|Davis et al., 2016]] ). These indicate no net global increase of SWV in the lower stratosphere since the late 1980s. [[#Hegglin--2014|Hegglin et al. (2014)]] reported a latitudinal dependence of SWV trends and suggested that the upward trend over Boulder should not be considered representative of the global stratosphere, while [[#Lossow--2018|Lossow et al. (2018)]] showed insignificant differences between SWV trends at Boulder and those for the 35–45°N zonal mean from 1980 to 2010 using model simulations and satellite observations.

Recent studies of dynamical influences on SWV ( [[#Eguchi--2015|Eguchi et al., 2015]] ; [[#Evan--2015|Evan et al., 2015]] ; [[#Tao--2015|Tao et al., 2015]] ; [[#Konopka--2016|Konopka et al., 2016]] ; [[#Diallo--2018|Diallo et al., 2018]] ; [[#Garfinkel--2018|Garfinkel et al., 2018]] ) have demonstrated that the quasi-biennial oscillation (QBO), El Niño–Southern Oscillation (ENSO), Sudden Stratospheric Warming (SSW) events and possibly also Pacific Decadal Variability (PDV; W. [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|]] [[#Wang--2016|Wang et al., 2016]] ), can significantly influence SWV abundance and the tropical cold point tropopause temperatures that largely control water vapour entering the stratosphere. It has also been shown that the convective lofting of ice can moisten the lower stratosphere over large regions ( [[#Dessler--2016|Dessler et al., 2016]] ; [[#Anderson--2017|Anderson et al., 2017]] ; [[#Avery--2017|Avery et al., 2017]] ). Near-global observations of SWV have revealed unusually strong and abrupt interannual changes, especially in the tropical lower stratosphere. Between December 2015 and November 2016, the tropical mean SWV anomaly at 82 hPa dropped from 0.9 ± 0.1 ppm to –1.0 ± 0.1 ppm, accompanied by highly anomalous QBO-related dynamics in the tropical stratosphere ( [[#Newman--2016|]] [[#Newman--2016|P.A. Newman et al., 2016]] ; [[#Tweedy--2017|Tweedy et al., 2017]] ) and the transition of ENSO from strong El Niño to La Niña conditions ( [[#Davis--2017|Davis et al., 2017]] ). The tropical mean SWV anomaly then rose sharply to 0.7 ± 0.1 ppm in June 2017 as warm westerlies returned to the tropical lower stratosphere and ENSO neutral conditions prevailed ( [[#Davis--2017|Davis et al., 2017]] ).

In summary, in situ measurements at a single mid-latitude location indicate about a 25% net increase in stratospheric water vapour since 1980, while merged satellite data records since the late 1980s suggest little net change. Recent studies of dynamical influences on SWV have highlighted their substantial roles in driving large interannual variability that complicates trend detection. There thus continues to be ''low confidence'' in trends of SWV over the instrumental period. Disregarding dynamic influences on SWV, an ERF of 0.05 ± 0.05 W m <sup>–2</sup> is estimated for SWV produced by CH <sub>4</sub> oxidation (Section 7.3.2.6), unchanged from AR5.

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<span id="stratospheric-ozone"></span>
==== 2.2.5.2 Stratospheric Ozone ====

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The AR5 assessed that it was certain that global stratospheric ozone from the mid-1990s to 2011 was nearly constant and about 3.5% lower than in the reference period 1964–1980. Most of the declines occurred prior to the mid-1990s.

Global annual mean total ozone (Figure 2.7) significantly declined by about 3.5% during the 1980s and the early 1990s and by 2.5% over 60°S–60°N (near-global). Then, during 2000–2017, both global and near-global concentrations increased slightly, but not significantly, all in line with production and consumption limits of ODS regulated under the Montreal Protocol and its amendments. Near-global 2014–2017 mean total ozone is about 2.2% below the pre-ozone depletion 1964–1980 average ( [[#Braesicke--2018|Braesicke et al., 2018]] ). At southern and northern mid-latitudes, declines are 5.5% and 3.0% compared to the 1964–1980 average respectively. Total ozone remained practically unchanged in the tropics ( [[#Braesicke--2018|Braesicke et al., 2018]] ). Emission of ODS started before 1980 and some estimates suggest that as much as 40% of the long-term ozone loss occurred between 1960 and 1980 ( [[#Shepherd--2014|Shepherd et al., 2014]] ), lowering the 1964–1980 baseline values by about 1% (outside the polar regions), a value close to observational uncertainties. The world’s longest record of total ozone measurements from Arosa, Switzerland, initiated in 1926, does not show any substantial long-term changes before about 1980 ( [[#Staehelin--2018|Staehelin et al., 2018]] ).

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[[File:51f32fe55a5202814a50fe02a6e8fc7a IPCC_AR6_WGI_Figure_2_7.png]]

'''Figure''' '''2.7 |''' '''Time series of annual mean total column ozone from 1964–2019.''' Values are in Dobson Units (DU), a good proxy for vertically integrated stratospheric ozone. Time series are shown for '''(a)''' near-global domain; '''(b–d)''' three zonal bands; and '''(e)''' polar (60°–90°) total ozone in March (Northern Hemisphere) and October (Southern Hemisphere): the months when polar ozone losses usually are largest. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

ERF depends strongly on the altitude of ozone changes. Two stratospheric regions are mainly responsible for long-term changes outside the polar regions. In the upper stratosphere (35–45 km), there was a strong decline (about 10%) from the start of observations in 1979 up to the mid-1990s and a subsequent increase by about 4% to present (SPARC/IO3C/GAW, 2019). In the lower stratosphere (20–25 km), there also was a statistically significant decline (7–8%) up to the mid-1990s, followed by stabilization or a small further decline ( [[#Ball--2018|Ball et al., 2018]] , 2019), although the natural variability is too strong to make a conclusive statement ( [[#Chipperfield--2018|Chipperfield et al., 2018]] ).

The strongest ozone loss in the stratosphere continues to occur in austral spring over Antarctica (ozone hole) with emergent signs of recovery after 2000 ( [[#Langematz--2018|Langematz et al., 2018]] ). Interannual variability in polar stratospheric ozone is driven by large scale winds and temperatures, and, to a lesser extent, by the stratospheric aerosol loading and the solar cycle. This variability is particularly large in the Arctic, where the largest depletion events, comparable to a typical event in the Antarctic, occurred in 2011 ( [[#Manney--2011|Manney et al., 2011]] ; [[#Langematz--2018|Langematz et al., 2018]] ) and again in 2020 ( [[#Manney--2020|Manney et al., 2020]] ; [[#Grooß--2021|Grooß and Müller, 2021]] ). Further details on trends and ERF can be found in Sections 6.3.2 and 7.3.2.5.

In summary, compared to the 1964–1980 average, stratospheric ozone columns outside polar regions (60°S–60°N) declined by about 2.5% over 1980–1995, and stabilized after 2000, with 2.2% lower values in 2014–2017. Large ozone depletions continue to appear in spring in the Antarctic and, in particularly cold years, also in the Arctic. Model-based estimates disagree on the sign of the ERF due to stratospheric ozone changes, but agree that it is much smaller in magnitude than that due to tropospheric ozone changes (Section 7.3.2.5).

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<span id="tropospheric-ozone"></span>
==== 2.2.5.3 Tropospheric Ozone ====

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The AR5 assessed ''medium confidence'' in large-scale increases of tropospheric ozone at rural surface sites across the NH (1970–2010), and in a doubling of European surface ozone during the 20th century, with the increases of surface ozone in the SH being of ''low confidence'' . Surface ozone ''likely'' increased in East Asia, but levelled off or decreased in the eastern USA and western Europe. Free tropospheric trends (1971–2010) from ozonesondes and aircraft showed positive trends in most, but not all, assessed regions, and for most seasons and altitudes. This section focuses on large scale ozone changes; chemical and physical processes and regional changes in tropospheric ozone are assessed in Section 6.3.2.1 and Section 7.3.2.5 assesses radiative forcing.

Prior to 1850 ozone observations do not exist, but a recent analysis using clumped-isotope composition of molecular oxygen ( <sup>18</sup> O <sup>18</sup> O in O <sub>2</sub> ) trapped in polar firn and ice, combined with atmospheric chemistry model simulations, constrains the global tropospheric ozone increase to less than 40% between 1850 and 2005, with most of this increase occurring between 1950 and 1980 ( [[#Yeung--2019|Yeung et al., 2019]] ). Recently, the Tropospheric Ozone Assessment Report identified and evaluated 60 records of surface ozone observations collected at rural locations worldwide between 1896 and 1975, which were based on a range of measurement techniques with potentially large uncertainties ( [[#Tarasick--2019|Tarasick et al., 2019]] ). They found that from the mid-20th century (1930s to the early 1970s) to 1990–2014, rural surface ozone increased by 30–70% across the northern extra-tropics. This is smaller than the 100% 20th-century increase reported in AR5, which relied on far fewer measurement sites, all in Europe. In the northern tropics limited low-elevation historical data (1954–1975) provide no clear indication of surface ozone increases ( [[#Tarasick--2019|Tarasick et al., 2019]] ). However, similar to the northern mid-latitude increases, lower-free tropospheric ozone at Mauna Loa, Hawaii increased by approximately 50% from the late 1950s to present ( [[#Cooper--2020|Cooper et al., 2020]] ). Historical observations are too limited to draw conclusions on surface ozone trends in the SH tropics and mid-latitudes since the mid-20th century, with tropospheric ozone exhibiting little change across Antarctica ( [[#Tarasick--2019|Tarasick et al., 2019]] ; [[#Cooper--2020|Cooper et al., 2020]] ). Based on reliable UV absorption measurements at remote locations (surface and lower troposphere), ozone trends since the mid-1990s varied spatially at northern mid-latitudes, but increased in the northern tropics (2–17%; 1–6 ppbv per decade; ( [[#Cooper--2020|Cooper et al., 2020]] ; [[#Gaudel--2020|Gaudel et al., 2020]] ). Across the SH these more recent observations are too limited to determine zonal trends (e.g., tropics, mid-latitudes, high latitudes).

The earliest observations of free tropospheric ozone (1934–1955) are available from northern mid-latitudes where limited data indicate a tropospheric column ozone increase of 48 ± 30% up to 1990–2012 ( [[#Tarasick--2019|Tarasick et al., 2019]] ). Starting in the 1960s, records from ozonesondes show no significant changes in the free troposphere over the Arctic and mid-latitude regions of Canada, but trends are mainly positive elsewhere in the northern mid-latitudes ( [[#Oltmans--2013|Oltmans et al., 2013]] ; [[#Cooper--2020|Cooper et al., 2020]] ). Tropospheric column and free tropospheric trends since the mid-1990s based on commercial aircraft, ozonesonde observations and satellite retrievals (Figure 2.8b,c), are overwhelmingly positive across the northern mid-latitudes (2–7%; 1–4 ppbv per decade) and tropics (2–14%; 1–5 ppbv per decade), with the largest increases (8–14%; 3–6 ppbv per decade) in the northern tropics in the vicinity of southern Asia and Indonesia. Observations in the SH are limited, but indicate average tropospheric column ozone increases of 2–12% (1–5 ppbv) per decade in the tropics (Figure 2.8c), and weak tropospheric column ozone increases (&lt;5%, &lt;1 ppbv per decade) at mid-latitudes ( [[#Cooper--2020|Cooper et al., 2020]] ). Above Antarctica, mid-tropospheric ozone has increased since the late 20th century ( [[#Oltmans--2013|Oltmans et al., 2013]] ). The total ozone ERF from 1750 to 2019 best estimate is assessed as 0.47 W m <sup>–2</sup> (Section 7.3.2.5) and this is dominated by increases in the troposphere. The underlying modelled global tropospheric ozone column increase ( [[#Skeie--2020|Skeie et al., 2020]] ) from 1850 to 2010 of 40–60%, is somewhat higher than the isotope based upper-limit of [[#Yeung--2019|Yeung et al. (2019)]] . At mid-latitudes (30°–60°N) model increases of 30–40% since the mid-20th century are broadly consistent with observations.

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[[File:29dfc27e28be57444bcf16f98ff974d1 IPCC_AR6_WGI_Figure_2_8.png]]

'''Figure 2.8''' '''|''' '''Surface and tropospheric ozone trends. (a)''' Decadal ozone trends by latitude at 28 remote surface sites and in the lower free troposphere (650 hPa, about 3.5 km) as measured by IAGOS aircraft above 11 regions. All trends are estimated for the time series up to the most recently available year, but begin in 1995 or 1994. Colours indicate significance (p-value) as denoted in the in-line key. See Figure 6.5 for a depiction of these trends globally. '''(b)''' Trends of ozone since 1994 as measured by IAGOS aircraft in 11 regions in the mid-troposphere (700–300 hPa; about 3–9 km) and upper troposphere (about 10–12 km), as measured by IAGOS aircraft and ozonesondes. '''(c)''' Trends of average tropospheric column ozone mixing ratios from the TOST composite ozonesonde product and three composite satellite products based on TOMS, OMI/MLS (Sat1), GOME, SCIAMACHY, OMI, GOME-2A, GOME-2B (Sat2), and GOME, SCIAMACHY, GOME-II (Sat3). Vertical bars indicate the latitude range of each product, while horizontal lines indicate the ''very likely'' uncertainty range. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

In summary, ''limited'' available isotopic ''evidence'' constrains the global tropospheric ozone increase to less than 40% between 1850 and 2005 ( ''low confidence'' ). Based on sparse historical surface/low altitude data tropospheric ozone has increased since the mid-20th century by 30–70% across the NH ( ''medium confidence'' ). Surface/low altitude ozone trends since the mid-1990s are variable at northern mid-latitudes, but positive in the tropics [2 to 17% per decade] ( ''high confidence'' ). Since the mid-1990s, free tropospheric ozone has increased by 2–7% per decade in most regions of the northern mid-latitudes, and 2–12% per decade in the sampled regions of the northern and southern tropics ( ''high confidence'' ). Limited coverage by surface observations precludes identification of zonal trends in the SH, while observations of tropospheric column ozone indicate increases of less than 5% per decade at southern mid-latitudes ( ''medium confidence'' ).

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<span id="aerosols"></span>
=== 2.2.6 Aerosols ===

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The AR5 assessed large-scale aerosol optical depth (AOD) trends over 2000–2009, concluding that there was ''low confidence'' in a global trend, but that AOD ''very likely'' decreased from 1990 onwards over Europe and the eastern USA, and increased since 2000 over eastern and southern Asia. The ERF associated with aerosol–radiation interactions for 2011 (relative to 1750) was estimated to be –0.45 ± 0.5 W m <sup>–2</sup> and of aerosol–cloud interaction estimated as –0.45 [–1.2 to 0.0] W m <sup>–2</sup> . Aerosol ERF uncertainty was assessed as the largest contributor to the overall ERF uncertainty since 1750.

This section assesses the observed large-scale temporal evolution of tropospheric aerosols. Aerosol-related processes, chemical and physical properties, and links to air quality, are assessed in Chapter 6. An in-depth assessment of aerosol interactions with radiation and clouds is provided in Section 7.3.3.

Aerosol proxy records of improved temporal resolution and quality are now available ( [[#Kylander--2016|Kylander et al., 2016]] ; [[#Stevens--2016|Stevens et al., 2016]] , 2018; [[#Jacobel--2017|Jacobel et al., 2017]] ; [[#Dornelas--2018|Dornelas et al., 2018]] ; [[#Middleton--2018|Middleton et al., 2018]] ), which further advance synthesis of new global compilations of aerosol loadings ( [[#Lambert--2015|Lambert et al., 2015]] ; [[#Albani--2016|Albani et al., 2016]] ). Estimates of the glacial/interglacial ratio in global dust deposition are within the range of 2–4 ( [[#Albani--2015|Albani et al., 2015]] ; [[#Lambert--2015|Lambert et al., 2015]] ). New reconstructions indicate a ratio of 3–5 for the glacial/interglacial loadings for mid- and high-latitude ocean of both hemispheres ( [[#Lamy--2014|Lamy et al., 2014]] ; [[#Martinez-Garcia--2014|Martinez-Garcia et al., 2014]] ; [[#Serno--2015|Serno et al., 2015]] ). Improved quantification of changes in dust deposition from North Africa and North Atlantic sediment records confirms dust deposition rates lower by a factor 2–5 during the African Humid Period (10–5 ka) compared to the late Holocene ( [[#McGee--2013|McGee et al., 2013]] ; [[#Albani--2015|Albani et al., 2015]] ; [[#Middleton--2018|Middleton et al., 2018]] ; [[#Palchan--2019|Palchan and Torfstein, 2019]] ). During the Holocene, biogenic emissions and volcanic activity drove significant variability (up to one order of magnitude) in sulphate concentrations ( [[#Schüpbach--2018|Schüpbach et al., 2018]] ).

Ice cores allow for estimation of multi-centennial trends in mid- and high-latitude aerosol deposition, including those for sulphate and black carbon (Figure 2.9a,b). Sulphate in ice cores increased by a factor of 8 from the end of the 19th century to the 1970s in continental Europe, by a factor of 4 from the 1940s to the 1970s in Russia, and by a factor of 3 from the end of the 19th century to 1950 in the Arctic (Svalbard). In all regions studied, concentrations have declined by about a factor of 2 following their peak (around 1970 in Europe and Russia, and 1950 in the Arctic). Strong increases of black carbon (BC) were observed in the 20th century over Europe, Russia, Greenland (primarily originating from emissions from North America), and in the Arctic (Svalbard). South America exhibits a small positive trend (Figure 2.9). BC concentrations in various Antarctic ice cores were below 1 ng g <sup>–1</sup> without a clear trend.

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[[File:d620688ec652c435c624d8090aff573c IPCC_AR6_WGI_Figure_2_9.png]]

'''Figure 2.9''' '''|''' '''Aerosol evolution from ice-core measurements.''' Changes are shown as 10-year averaged time series '''(a, b)''' and trends in remote-sensing aerosol optical depth (AOD) and AODf '''(c, d). (a)''' Concentrations of non-sea salt (nss) sulphate (ng g <sup>–1</sup> ). '''(b)''' Black carbon (BC) in glacier ice from the Arctic (Lomonosovfonna), Russia (Belukha), Europe (Colle Gnifetti), South America (Illimani), Antarctica (stacked sulphate record, and BC from the B40 core), and BC from Greenland (stacked rBC record from Greenland and eastern Europe (Elbrus)). '''(c)''' Linear trend in annual mean AOD retrieved from satellite data for the 2000–2019 period (% yr <sup>–1</sup> ). The average trend from MODerate Resolution Imaging Spectroradiometer (MODIS) and Multi-Angle Imaging Spectroradiometer (MISR) is shown. Trends are calculated using OLS regression with significance assessed following AR(1) adjustment after [[#Santer--2008|Santer et al. (2008)]] . Superimposed are the trends in annual-mean AOD from the AERONET surface sunphotometer network for 2000–2019. '''(d)''' Linear trend in 2000–2019 as in (c), but for fine-mode AOD, AODf, and using only MISR over land. ‘×’ marks denote non-significant trends. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Spatially resolved trends of AOD derived from Aqua/Terra MISR and MODIS instruments over 2000–2019 range between –2% and +2% per year (Figure 2.9c). Ground-based solar attenuation networks help to constrain and improve the satellite-derived retrievals of AOD, and trends derived from the AERONET network (Figure 2.9c,d) corroborate satellite results ( [[#Georgoulias--2016|Georgoulias et al., 2016]] ; [[#Wei--2019|Wei et al., 2019]] ; [[#Bauer--2020|Bauer et al., 2020]] ; H. [[#Yu--2020|]] [[#Yu--2020|Yu et al., 2020]] ) in particular for declines over Europe ( [[#Stjern--2011|Stjern et al., 2011]] ; [[#Cherian--2014|Cherian et al., 2014]] ; [[#Li--2014|Li et al., 2014]] ) and the USA ( [[#Li--2014|Li et al., 2014]] ; [[#Jongeward--2016|Jongeward et al., 2016]] ). The tendency in AOD over East Asia reversed from positive (2000–2010) to negative (since 2010) ( [[#Sogacheva--2018|Sogacheva et al., 2018]] ; [[#Filonchyk--2019|Filonchyk et al., 2019]] ; [[#Ma--2019|Ma et al., 2019]] ; [[#Samset--2019|Samset et al., 2019]] ). Over southern Asia, however, AOD from satellite (MODIS/MISR) and AERONET retrievals show continuing increases ( [[#Li--2014|Li et al., 2014]] ; [[#Zhao--2017|Zhao et al., 2017]] ), with similar trends from UV-based aerosol retrievals from the Ozone Monitoring Instrument (OMI) on the Aura satellite ( [[#Dahutia--2018|Dahutia et al., 2018]] ; [[#Hammer--2018|Hammer et al., 2018]] ). A comparison of MODIS and MISR radiometric observations with the broadband CERES satellite instrument ( [[#Corbett--2015|Corbett and Loeb, 2015]] ) showed that drifts in calibration are ''unlikely'' to affect the satellite derived trends. CERES shows patterns for clear-sky broadband radiation consistent with the aerosol spatio-temporal changes ( [[#Loeb--2018|Loeb et al., 2018]] ; [[#Paulot--2018|Paulot et al., 2018]] ).

Satellite-derived trends are further supported by in situ regional surface concentration measurements, operational since the 1980s (sulphate) and 1990s (PM2.5) from a global compilation ( [[#Collaud%20Coen--2020|Collaud Coen et al., 2020]] ) of networks over Europe ( [[#Stjern--2011|Stjern et al., 2011]] ), North America ( [[#Jongeward--2016|Jongeward et al., 2016]] ), and China ( [[#Zheng--2018|Zheng et al., 2018]] ). [[#Collaud%20Coen--2020|Collaud Coen et al. (2020)]] report from surface observations across the NH mid-latitudes that aerosol absorption coefficients decreased since the first decade of the 21st century.

Anthropogenic aerosol is predominantly found in the fraction of particles with radii &lt;1 µm that comprise the fine-mode AOD (AODf; Figure 2.9d; [[#Kinne--2019|Kinne, 2019]] ). A significant decline in AODf of more than 1.5% per year from 2000 to 2019 has occurred over Europe and North America, while there have been positive trends of up to 1.5% per year over Southern Asia and East Africa. The global-scale trend in AODf of –0.03% per year (Figure 2.9) is significant. The results are consistent with trend estimates from an aerosol reanalysis ( [[#Bellouin--2020|Bellouin et al., 2020]] ), and the trends in satellite-derived cloud droplet number concentrations are consistent with the aerosol trends ( [[#Cherian--2020|Cherian and Quaas, 2020]] ). Cloudiness and cloud radiative properties trends are, however, less conclusive possibly due to their large variability ( [[#Norris--2016|Norris et al., 2016]] ; [[#Cherian--2020|Cherian and Quaas, 2020]] ). Further details on aerosol-cloud interactions are assessed in Section 7.3.3.2.

To conclude, atmospheric aerosols sampled by ice cores, influenced by northern mid-latitude emissions, show positive trends from 1700 until the last quarter of the 20th century and decreases thereafter ( ''high confidence'' ), but there is ''low confidence'' in observations of systematic changes in other parts of the world in these periods. Satellite data and ground-based records indicate that AOD exhibits predominantly negative trends since 2000 over NH mid-latitudes and SH continents, but increased over South Asia and East Africa ( ''high confidence'' ). A globally deceasing aerosol abundance is thus assessed with ''medium confidence'' . This implies increasing net positive ERF, since the overall negative aerosol ERF has become smaller.

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=== 2.2.7 Land Use and Land Cover ===

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The AR5 assessed that land use change ''very likely'' increased the Earth’s albedo with a radiative forcing of –0.15 (± 0.10) W m <sup>–2</sup> . AR5 also assessed that a net cooling of the surface, accounting for processes that are not limited to the albedo, was ''about as likely as not'' . The SRCCL concluded with ''medium confidence'' that the biophysical effects of land cover change (mainly increased albedo) had a cooling effect on surface temperatures. The SRCCL also concluded with ''very high confidence'' that the biogeochemical effects of land cover change (i.e., GHG emissions) resulted in a mean annual surface warming.

Much of the global land surface has been modified or managed to some extent by human activities during the Holocene. Reconstructions based on pollen data indicate that natural vegetation probably covered most of the Earth’s ice-free terrestrial surface until roughly the mid-Holocene ( [[#Marquer--2017|Marquer et al., 2017]] ; [[#Harrison--2020|Harrison et al., 2020]] ; F. [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ). Reconstructions based on pollen, archaeological, and historical data indicate deforestation at the regional scale since at least 6 ka ( [[#Marquer--2017|Marquer et al., 2017]] ; [[#Stephens--2019|Stephens et al., 2019]] ; [[#Harrison--2020|Harrison et al., 2020]] ; F. [[#Li--2020|]] [[#Li--2020|Li et al., 2020]] ). From a global perspective, land-use forcing datasets ( [[#Lawrence--2016|Lawrence et al., 2016]] ) estimate that changes in land use (and related deforestation) were small on the global scale until the mid-19th century and accelerated markedly thereafter, with larger uncertainties prior to industrialization ( [[#Kaplan--2017|Kaplan et al., 2017]] ). Since the early 1980s, about 60% of all land cover changes have been associated with direct human activities, with spatial patterns emphasizing the regional character of land use and land management, including tropical deforestation, temperate afforestation, cropland intensification, and increased urbanization ( [[#Song--2018|Song et al., 2018]] ; [[#Zeng--2018|Zeng et al., 2018]] ). At present, nearly three-quarters of the ice-free terrestrial surface is under some form of human use ( [[#Venter--2016|Venter et al., 2016]] ; [[#Erb--2017|Erb et al., 2017]] ), particularly in agriculture and forest management.

The impact of historical land-cover change on global climate is assessed with model simulations that consider multiple climate and biophysical processes (e.g., changes in albedo, evapotranspiration, and roughness) and/or biogeochemical processes (e.g., changes in atmospheric composition such as carbon release from deforestation). The dominant biophysical response to land cover changes is albedo, which is estimated (using a MODIS albedo product and a historical land-use harmonization product) to have increased gradually prior to the mid-19th century and then strongly through the mid-20th century, with a slightly slower rise thereafter ( [[#Ghimire--2014|Ghimire et al., 2014]] ). Recent radiative forcing estimates arising from biophysical processes generally fall at the lower end of the AR5 assessed range. For instance, based on historical simulations from 13 CMIP6 models, C.J. [[#Smith--2020|]] [[#Smith--2020|Smith et al. (2020)]] estimated that the ERF from surface albedo changes (including snow cover and leaf area) was –0.08 [–0.22 to +0.06] W m <sup>–2</sup> since 1850. Similarly, based on simulations from 13 CMIP5 models, [[#Lejeune--2020|Lejeune et al. (2020)]] estimated the radiative forcing from transitions between trees, crops, and grasslands was –0.11 [–0.16 to +0.04] W m <sup>–2</sup> since 1860. [[#Andrews--2017|Andrews et al. (2017)]] identified an ERF of –0.40 W m <sup>–2</sup> since 1860, ascribing much of the effect to increases in albedo (including the unmasking of underlying snow cover); notably, however, the analysis was based on a single model with a known tendency to overestimate the ERF ( [[#Collins--2011|Collins et al., 2011]] ). [[#Ward--2014|Ward et al. (2014)]] examined the combined effects of biophysical and biogeochemical processes, obtaining an RF of 0.9 ± 0.5 W m <sup>–2</sup> since 1850 that was driven primarily by increases in land-use related GHG emissions from deforestation and agriculture ( [[#Ward--2015|Ward and Mahowald, 2015]] ). According to a large suite of historical simulations, the biophysical effects of changes in land cover (i.e., increased surface albedo and decreased turbulent heat fluxes) led to a net global cooling of 0.10°C ± 0.14°C at the surface (SRCCL). Available model simulations suggest that biophysical and biogeochemical effects jointly may have contributed to a small global warming of 0.078°C ± 0.093°C at the surface over about the past two centuries (SRCCL), with a potentially even larger warming contribution over the Holocene as a whole ( [[#He--2014|He et al., 2014]] ).

In summary, biophysical effects from historical changes in land use have an overall negative ERF ( ''medium confidence'' ). The best-estimate ERF from the increase in global albedo is –0.15 W m <sup>–2</sup> since 1700 and –0.12 W m <sup>–2</sup> since 1850 ( ''medium confidence'' ) (Section 7.3.4.1). Biophysical effects of land-use change ''likely'' resulted in a net global cooling of about 0.1°C since 1750 ( ''medium confidence'' ) (Section 7.3.5.3).

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<span id="effective-radiative-forcing-erf-exerted-by-the-assessed-climate-drivers"></span>
=== 2.2.8 Effective Radiative Forcing (ERF) Exerted by the Assessed Climate Drivers ===

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The AR5 concluded that changes in climate drivers over the industrial period corresponded to a positive ERF which increased more rapidly after 1970 than before. There was ''very high confidence'' in the positive ERF due to WMGHG, with CO <sub>2</sub> the single largest contributor. The AR5 concluded that there was ''high confidence'' that aerosols have offset a substantial portion of the WMGHG forcing.

This section reports the evolution in ERF with respect to 1750 as assessed in Section 7.3 and relies on the observed changes in climate drivers as assessed in [[#2.2|Section 2.2]] wherever possible, and models otherwise. The ERF is assessed using the methods and details described in Section 7.3.1 and includes, in addition to the radiative forcing, the rapid adjustments, especially implied by clouds. The time series are shown in Figure 2.10.

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[[File:b59110924317078cfe585f9691ea0777 IPCC_AR6_WGI_Figure_2_10.png]]

'''Figure 2.1''' '''0 |''' '''Temporal evolution of effective radiative forcing (ERF) related to the drivers assessed in [[#2.2|Section 2.2]] .''' ERFs are based upon the calculations described in Chapter 7, of which the global annual mean, central assessment values are shown as lines and the 5 to 95% uncertainty range as shading (Section 7.3, see Figures 7.6 to 7.8 for more detail on uncertainties). The inset plot shows the rate of change (linear trend) in total anthropogenic ERF (total without TSI and volcanic ERF) for 30-year periods centred at each dot. Further details on data sources and processing are available in the chapter data table (Table 2.SM.1).

Increasing TSI ( [[#2.2.1|Section 2.2.1]] ) implies a small ERF of less than 0.1 W m <sup>–2</sup> between 1900 and 1980. TSI varies over the 11-year solar cycle with ERF of order ± 0.1 W m <sup>–2</sup> in the assessed period. Strong volcanic eruptions ( [[#2.2.2|Section 2.2.2]] ) with periods of strong negative ERF lasting 2–5 years in duration occurred in the late 19th and early 20th centuries. There followed a relatively quiescent period between about 1920 and 1960, and then three strong eruptions in 1963, 1982 and 1991, and only small-to-moderate eruptions thereafter ( [[#Schmidt--2018|Schmidt et al., 2018]] ).

The atmospheric concentrations of WMGHGs ( [[#2.2.3|Section 2.2.3]] ) have continuously increased since the early 19th century, with CO <sub>2</sub> contributing the largest share of the positive ERF. Compared to the last two decades of the 20th century, the growth rate of CO <sub>2</sub> in the atmosphere increased in the 21st century, showed strong fluctuations for CH <sub>4</sub> , and was about constant for N <sub>2</sub> O. Mixing ratios of the most abundant CFCs declined ( [[#2.2.4|Section 2.2.4]] ). Mixing ratios of HCFCs increased, but growth rates are starting to decelerate. Mixing ratios of HFCs and some other human-made components are increasing ( [[#2.2.4|Section 2.2.4]] ). The ERF for CO <sub>2</sub> alone is stronger than for all the other anthropogenic WMGHGs taken together throughout the industrial period, and its relative importance has increased in recent years (Figures 2.10 and 7.6).

Among the gaseous short-lived climate forcers ( [[IPCC:Wg1:Chapter:Chapter-6|Chapter 6]] and Sections 2.2.5 and 7.3; excluding CH <sub>4</sub> here), ozone (O <sub>3</sub> ) is the component with the largest (positive) ERF. Concentrations from direct observations have increased since the mid-20th century and, mostly based on models, this extends to since 1750. Other gaseous short-lived climate forcers have small contributions to total ERF.

The net effect of aerosols (Sections 2.2.6 and 6.4) on the radiation budget, including their effect on clouds, and cloud adjustments, as well as the deposition of black carbon on snow (Section 7.3.4.3), was negative throughout the industrial period ( ''high confidence'' ). The net effect strengthened (becoming more negative) over most of the 20th century, but ''more likely than not'' weakened (becoming less negative) since the late 20th century. These trends are reflected in measurements of surface solar radiation (Section 7.2.2.3) and the Earth’s energy imbalance (Section 7.2.2.1). The relative importance of aerosol forcing compared to other forcing agents has decreased globally in the most recent 30 years ( ''medium confidence'' ) and the reduction of the negative forcing in the 21st century enhances the overall positive ERF.

Land use and land cover changes ( [[#2.2.7|Section 2.2.7]] ) over the industrial period introduce a negative radiative forcing by increasing the surface albedo. This effect increased since 1750, reaching current values of about –0.20 W m <sup>–2</sup> ( ''medium confidence'' ). This ERF value is taken from Section 7.3.4.1 and is different from the assessment in [[#2.2.7|Section 2.2.7]] in that it also includes the effect of irrigation. It also includes uncertain rapid adjustments and thus there is ''low confidence'' in its magnitude. Biogeochemical feedbacks can be substantial (Section 5.4) and are not included in ERF.

In conclusion, the net ERF due to all observed changes in climate drivers is positive, except for short periods (up to a few years in duration) following moderate to large volcanic eruptions, and has grown in magnitude since the late 19th century. The rate of change ''likely'' has increased in the last 30 years, since CO <sub>2</sub> concentrations increased at an increasing rate due to growing CO <sub>2</sub> emissions ( ''very likely'' ) , and since the aerosol forcing became less negative ( ''more likely than not'' ).

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