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

=== 2.4.2 Carbonaceous aerosols ===

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Carbonaceous aerosols are one of the most abundant components of aerosol particles in continental areas of the atmosphere and a key land–atmosphere component (Contini et al. 2018 <sup>[[#fn:r836|836]]</sup> ). They can make up to 60–80% of PM2.5 (particulate matter with size less than 2.5 μm) in urban and remote atmospheres (Tsigaridis et al. 2014 <sup>[[#fn:r837|837]]</sup> ; Kulmala et al. 2011 <sup>[[#fn:r838|838]]</sup> ). It comprises an organic fraction (OC) and a refractory light-absorbing component, generally referred to as elemental carbon (EC), from which BC is the optically active absorption component of EC (Gilardoni et al. 2011 <sup>[[#fn:r839|839]]</sup> ; Bond et al. 2013 <sup>[[#fn:r840|840]]</sup> ).

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==== 2.4.2.1 Carbonaceous aerosol precursors of short-lived climate forcers from land ====

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OC is a major component of aerosol mass concentration, and it originates from different anthropogenic (combustion processes) and natural (natural biogenic emissions) sources (Robinson et al. 2007 <sup>[[#fn:r841|841]]</sup> ). A large fraction of OC in the atmosphere has a secondary origin, as it can be formed in the atmosphere through condensation to the aerosol phase of low vapour pressure gaseous compounds emitted as primary pollutants or formed in the atmosphere. This component is SOA (Hodzic et al. 2016 <sup>[[#fn:r842|842]]</sup> ). A third component of the optically active aerosols is the so-called brown carbon (BrC), an organic material that shows enhanced solar radiation absorption at short wavelengths (Wang et al. 2016b <sup>[[#fn:r843|843]]</sup> ; Laskin et al. 2015 <sup>[[#fn:r844|844]]</sup> ; Liu et al. 2016a <sup>[[#fn:r845|845]]</sup> ; Bond et al. 2013 <sup>[[#fn:r846|846]]</sup> ; Saturno et al. 2018 <sup>[[#fn:r847|847]]</sup> ).

OC and EC have distinctly different optical properties, with OC being important for the scattering properties of aerosols and EC central for the absorption component (Rizzo et al. 2013 <sup>[[#fn:r848|848]]</sup> ; Tsigaridis et al. 2014 <sup>[[#fn:r849|849]]</sup> ; Fuzzi et al. 2015 <sup>[[#fn:r850|850]]</sup> ). While OC is reflective and scatters solar radiation, it has a cooling effect on climate. On the other side, BC and BrC absorb solar radiation and they have a warming effect in the climate system (Bond et al. 2013 <sup>[[#fn:r851|851]]</sup> ).

OC is also characterised by a high solubility with a high fraction of water-soluble organic compounds (WSOC) and it is one of the main drivers of the oxidative potential of atmospheric particles. This makes particles loaded with oxidised OC an efficient cloud condensation nuclei (CCN) in most of the conditions (Pöhlker et al. 2016 <sup>[[#fn:r852|852]]</sup> ; Thalman et al. 2017 <sup>[[#fn:r853|853]]</sup> ; Schmale et al. 2018 <sup>[[#fn:r854|854]]</sup> ).

Biomass burning is a major global source of carbonaceous aerosols (Bowman et al. 2011 <sup>[[#fn:r855|855]]</sup> ; Harrison et al. 2010 <sup>[[#fn:r856|856]]</sup> ; Reddington et al. 2016 <sup>[[#fn:r857|857]]</sup> ; Artaxo et al. 2013 <sup>[[#fn:r858|858]]</sup> ). As knowledge of past fire dynamics improved through new satellite observations, new fire proxies’ datasets (Marlon et al. 2013 <sup>[[#fn:r859|859]]</sup> ; van Marle et al. 2017a <sup>[[#fn:r860|860]]</sup> ), process-based models (Hantson et al. 2016 <sup>[[#fn:r861|861]]</sup> ) and a new historic biomass burning emissions dataset starting in 1750 have been developed (van Marle et al. 2017b <sup>[[#fn:r862|862]]</sup> ) (Cross-Chapter Box 3 in this chapter). Revised versions of OC biomass burning emissions (van Marle et al. 2017b <sup>[[#fn:r863|863]]</sup> ) show, in general, reduced trends compared to the emissions derived by Lamarque et al. (2010) <sup>[[#fn:r864|864]]</sup> for CMIP5. CMIP6 global emissions pathways (Gidden et al. 2018 <sup>[[#fn:r865|865]]</sup> ; Hoesly et al. 2018 <sup>[[#fn:r866|866]]</sup> ) estimate global BC emissions in 2015 at 9.8 MtBC yr <sup>–1</sup> , while global OC emissions are 35 MtOC yr <sup>–1</sup> .

Land use change is critically important for carbonaceous aerosols, since biomass-burning emissions consist mostly of organic aerosol, and the undisturbed forest is also a large source of organic aerosols (Artaxo et al. 2013 <sup>[[#fn:r867|867]]</sup> ). Additionally, urban aerosols are also mostly carbonaceous because of the source composition (traffic, combustion, industry, etc.) (Fuzzi et al. 2015 <sup>[[#fn:r868|868]]</sup> ). Burning of fossil fuels, biomass- burning emissions and SOA from natural BVOC emissions are the main global sources of carbonaceous aerosols. Any change in each of these components directly influence the radiative forcing (Contini et al. 2018 <sup>[[#fn:r869|869]]</sup> ; Boucher et al. 2013 <sup>[[#fn:r870|870]]</sup> ; Bond et al. 2013 <sup>[[#fn:r871|871]]</sup> ).

One important component of carbonaceous aerosols is the primary biological aerosol particles (PBAP), also called bioaerosols, that correspond to a significant fraction of aerosols in forested areas (Fröhlich-Nowoisky et al. 2016 <sup>[[#fn:r872|872]]</sup> ; Pöschl and Shiraiwa 2015 <sup>[[#fn:r873|873]]</sup> ). They are emitted directly by the vegetation as part of the biological processes (Huffman et al. 2012 <sup>[[#fn:r874|874]]</sup> ). Airborne bacteria, fungal spores, pollen, archaea, algae and other bioparticles are essential for the reproduction and spread of organisms across various terrestrial ecosystems. They can serve as nuclei for cloud droplets, ice crystals and precipitation, thus influencing the hydrological cycle and climate (Whitehead et al. 2016 <sup>[[#fn:r875|875]]</sup> ; Scott et al. 2015 <sup>[[#fn:r876|876]]</sup> ; Pöschl et al. 2010 <sup>[[#fn:r877|877]]</sup> ).

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==== 2.4.2.2 Effects of past climate change on carbonaceous aerosols emissions and feedbacks ====

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Annual global emission estimates of BC range from 7.2–7.5 Tg yr <sup>–1</sup> (using bottom-up inventories) (Bond et al. 2013 <sup>[[#fn:r878|878]]</sup> ; Klimont et al. 2017 <sup>[[#fn:r879|879]]</sup> ) up to 17.8 ± 5.6 Tg yr <sup>–1</sup> (using a fully coupled climate-aerosol-urban model constrained by aerosol measurements) (Cohen and Wang 2014 <sup>[[#fn:r880|880]]</sup> ), with considerably higher BC emissions for Eastern Europe, southern East Asia, and Southeast Asia, mostly due to higher anthropogenic BC emissions estimates. A significant source of BC, the net trend in global burned area from 2000–2012 was a modest decrease of 4.3 Mha yr <sup>–1</sup> (–1.2% yr <sup>–1</sup> ).

Carbonaceous aerosols are important in urban areas as well as pristine continental regions, since they can be responsible for 50–85% of PM2.5 (Contini et al. 2018 <sup>[[#fn:r881|881]]</sup> ; Klimont et al. 2017 <sup>[[#fn:r882|882]]</sup> ). In boreal and tropical forests, carbonaceous aerosols originate from BVOC oxidation (Section 2.4.3). The largest global source of BC aerosols is open burning of forests, savannah and agricultural lands with emissions of about 2700 Gg yr <sup>–1</sup> in the year 2000 (Bond et al. 2013 <sup>[[#fn:r883|883]]</sup> ).

ESMs most likely underestimate globally averaged EC emissions (Bond et al. 2013 <sup>[[#fn:r884|884]]</sup> ; Cohen and Wang 2014 <sup>[[#fn:r885|885]]</sup> ), although recent emission inventories have included an upwards adjustment in these numbers (Hoesly et al. 2018 <sup>[[#fn:r886|886]]</sup> ). Vertical EC profiles have also been shown to be poorly constrained (Samset et al. 2014 <sup>[[#fn:r887|887]]</sup> ), with a general tendency of too much EC at high altitudes. Models differ strongly in the magnitude and importance of the coating-enhancement of ambient EC absorption (Boucher et al. 2016 <sup>[[#fn:r888|888]]</sup> ; Gustafsson and Ramanathan 2016 <sup>[[#fn:r889|889]]</sup> ) in their estimated lifetime of these particles, as well as in dry and wet removal efficiency ( ''limited evidence, medium agreement'' ) (Mahmood et al. 2016 <sup>[[#fn:r890|890]]</sup> ).

The equilibrium in emissions and concentrations between the scattering properties of organic aerosol versus the absorption component of BC is a key ingredient in the future climatic projections of aerosol effects ( ''limited evidence, high agreement'' ). The uncertainties in net climate forcing from BC-rich sources are substantial, largely due to lack of knowledge about cloud interactions with both BC and co-emitted OC. A strong positive forcing of about 1.1 wm <sup>–2</sup> was calculated by Bond et al. (2013), but this forcing is balanced by a negative forcing of –1.45 wm <sup>–2</sup> , and shows clearly a need to work on the co-emission issue for carbonaceous aerosols. The forcing will also depend on the aerosol-cloud interactions, where carbonaceous aerosol can be coated and change their CCN capability. It is difficult to estimate the changes in any of these components in a future climate, but this will strongly influence the radiative forcing ( ''high confidence'' ) (Contini et al. 2018 <sup>[[#fn:r891|891]]</sup> ; Boucher et al. 2013 <sup>[[#fn:r892|892]]</sup> ; Bond et al. 2013 <sup>[[#fn:r893|893]]</sup> ).

De Coninck et al. (2018) <sup>[[#fn:r894|894]]</sup> reported studies estimating a lower global temperature effect from BC mitigation (e.g., Samset et al. 2014 <sup>[[#fn:r895|895]]</sup> ; Boucher et al. 2016 <sup>[[#fn:r896|896]]</sup> ), although commonly used models do not capture properly observed effects of BC and co-emissions on climate (e.g., Bond et al. 2013 <sup>[[#fn:r897|897]]</sup> ). Regionally, the warming effects can be substantially larger, for example, in the Arctic (Sand et al. 2015 <sup>[[#fn:r898|898]]</sup> ) and high mountain regions near industrialised areas or areas with heavy biomass-burning impacts ( ''high confidence'' ) (Ming et al. 2013 <sup>[[#fn:r899|899]]</sup> ).

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==== 2.4.2.3 Future changes of carbonaceous aerosol emissions ====

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Due to the short atmospheric lifetime of carbonaceous aerosols in the atmosphere, of the order of a few days, most studies dealing with the future concentration levels have a regional character (Cholakian et al. 2018 <sup>[[#fn:r900|900]]</sup> ; Fiore et al. 2012 <sup>[[#fn:r901|901]]</sup> ). The studies agree that the uncertainties in changes in emissions of aerosols and their precursors are generally higher than those connected to climate change itself. Confidence in future changes in carbonaceous aerosol concentration projections is limited by the reliability of natural and anthropogenic emissions (including wildfires, largely caused by human activity) of primary aerosol as well as that of the precursors. The Aerosol Chemistry Model Intercomparison Project (AerChemMIP) is endorsed by the Coupled-Model Intercomparison Project 6 (CMIP6) and is designed to quantify the climate impacts of aerosols and chemically reactive gases (Lamarque et al. 2013 <sup>[[#fn:r902|902]]</sup> ). These simulations calculated future responses to SLCF emissions for the RCP scenarios in terms of concentration changes and radiative forcing. Carbonaceous aerosol emissions are expected to increase in the near future due to possible increases in open biomass-burning emissions (from forest, savannah and agricultural fires), and increase in SOA from oxidation of BVOCs ( ''medium confidence'' ) (Tsigaridis et al. 2014 <sup>[[#fn:r903|903]]</sup> ; van Marle et al. 2017b <sup>[[#fn:r904|904]]</sup> ; Giglio et al. 2013 <sup>[[#fn:r905|905]]</sup> ).

More robust knowledge has been produced since the conclusions reported in AR5 (Boucher et al. 2013 <sup>[[#fn:r906|906]]</sup> ) and all lines of evidence now agree on a small effect on carbonaceous aerosol global burden due to climate change ( ''medium confidence'' ). The regional effects, however, are predicted to be much higher (Westervelt et al. 2015 <sup>[[#fn:r907|907]]</sup> ). With respect to possible changes in the chemical composition of PM as a result of future climate change, only a few sparse data are available in the literature and the results are, as yet, inconclusive. The co-benefits of reducing aerosol emissions due to air quality issues will play an important role in future carbonaceous aerosol emissions ( ''high confidence'' ) (Gonçalves et al. 2018 <sup>[[#fn:r908|908]]</sup> ; Shindell et al. 2017 <sup>[[#fn:r909|909]]</sup> ).

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