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

== 2.4 Emissions and impacts of short-lived climate forcers (SLCF) from land ==

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While the rising atmospheric concentration of GHGs is the largest driver of anthropogenic changes in climate, the levels of short-lived climate forcers (SLCF) can significantly modulate regional climate by altering radiation exchanges and hydrological cycle and impact ecosystems ( ''high confidence'' ) (Boucher et al. 2013 <sup>[[#fn:r793|793]]</sup> ; Rogelj et al. 2014 <sup>[[#fn:r794|794]]</sup> ; Kok et al. 2018 <sup>[[#fn:r795|795]]</sup> ). This section assesses the current state of knowledge with respect to past and future emissions of the three major SLCFs and their precursors: mineral dust, carbonaceous aerosols (black carbon (BC) and organic carbon (OC)) and BVOCs. This section also reports on implications of changes in their emissions for climate. Aerosols particles with diameters between about 0.010 μm to about 20 μm are recognised as SLCFs, a term that refers to their short atmospheric lifetime (a few days). BVOCs are important precursors of ozone and OC, both important climate forcing agents with short atmospheric lifetimes.

While the AR5 did not assess land aerosols emissions in depth, their findings stated that although progress in quantifying regional emissions of anthropogenic and natural land aerosols has been made, considerable uncertainty still remains about their historical trends, their inter-annual and decadal variability and about any changes in the future (Calvo et al. 2013 <sup>[[#fn:r796|796]]</sup> ; Klimont et al. 2017 <sup>[[#fn:r797|797]]</sup> ). Some new and improved understanding of processes controlling emissions and atmospheric processing has been developed since AR5, for example, a better understanding of the climatic role of BC as well as the understanding of the role of BVOCs in formation of secondary organic aerosols (SOA).

Depending on the chemical composition and size, aerosols can absorb or scatter sunlight and thus directly affect the amount of absorbed and scattered radiation (Fuzzi et al. 2015 <sup>[[#fn:r798|798]]</sup> ; Nousiainen 2011 <sup>[[#fn:r799|799]]</sup> ; de Sá et al. 2019 <sup>[[#fn:r800|800]]</sup> ) Aerosols affect cloud formation and development, and thus can also influence precipitation patterns and amounts (Suni et al. 2015 <sup>[[#fn:r801|801]]</sup> ). In addition, deposition of aerosols – especially BC – on snow and ice surfaces can reduce albedo and increase warming as a self-reinforcing feedback. Aerosols deposition also changes biogeochemical cycling in critical terrestrial ecosystems, with deposition of nutrients such as nitrogen and phosphorus (Andreae et al. 2002 <sup>[[#fn:r802|802]]</sup> ). Primary land aerosols are emitted directly into the atmosphere due to natural or anthropogenic processes and include mineral aerosols (or dust), volcanic dust, soot from combustion, organic aerosols from industry, vehicles or biomass burning, bioaerosols from forested regions and others. SOAs are particulates that are formed in the atmosphere by the gas-to-particles conversion processes from gaseous precursors, such as BVOCs, and account for a large fraction of fine mode (particles less than 2.5μm) aerosol mass (Hodzic et al. 2016 <sup>[[#fn:r803|803]]</sup> ; Manish et al. 2017 <sup>[[#fn:r804|804]]</sup> ). Land use change can affect the climate through changed emissions of SLCFs such as aerosols, ozone precursors and methane.

Aerosols from air pollution will decline in the coming years as a means for improving urban and regional air, but their removal will lead to additional warming (Boucher et al. 2013 <sup>[[#fn:r805|805]]</sup> ), with important regional variability, and partially offsetting projected mitigation effects for two to three decades in 1.5°C consistent pathways ( ''high confidence'' ) (IPCC 2018 <sup>[[#fn:r806|806]]</sup> ). It is important to emphasise that changes in emissions can either be due to external forcing or through a feedback in the climate system (Box 2.1). For instance, enhanced dust emissions due to reduced vegetation could be a forcing if overgrazing is the cause of larger dust emission, or a feedback if dryer climate is the cause. This distinction is important in terms of mitigation measures to be implemented.

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=== 2.4.1 Mineral dust ===

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One of the most abundant atmospheric aerosols emitted into the atmosphere is mineral dust, a ‘natural’ aerosol that is produced by wind strong enough to initiate the emissions process of sandblasting. Mineral dust is preferentially emitted from dry and unvegetated soils in topographic depressions where deep layers of alluvium have been accumulated (Prospero et al. 2002 <sup>[[#fn:r807|807]]</sup> ). Dust is also emitted from disturbed soils by human activities, with a 25% contribution to global emissions based on a satellite-based estimate (Ginoux et al. 2012 <sup>[[#fn:r808|808]]</sup> ).

Dust is then transported over long distances across continents and oceans. The dust cycle, which consists of mineral dust emission, transport, deposition and stabilisation, has multiple interactions with many climate processes and biogeochemical cycles.

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<span id="mineral-dust-as-a-short-lived-climate-forcer-from-land"></span>
==== 2.4.1.1 Mineral dust as a short-lived climate forcer from land ====

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Depending on the dust mineralogy, mixing state and size, dust particles can absorb or scatter shortwave and longwave radiation. Dust particles serve as cloud condensation nuclei and ice nuclei. They can influence the microphysical properties of clouds, their lifetime and precipitation rate (Kok et al. 2018 <sup>[[#fn:r809|809]]</sup> ). New and improved understanding of processes controlling emissions and transport of dust, its regional patterns and variability, as well as its chemical composition, has been developed since AR5.

While satellites remain the primary source of information to locate dust sources and atmospheric burden, in-situ data remains critical to constrain optical and mineralogical properties of the dust (DiBiagio et al. 2017 <sup>[[#fn:r810|810]]</sup> ; Rocha-Lima et al. 2018 <sup>[[#fn:r811|811]]</sup> ). Dust particles are composed of minerals, including iron oxides which strongly absorb shortwave radiation and provide nutrients for marine ecosystems. Another mineral such as feldspar is an efficient ice nuclei (Harrison et al. 2016 <sup>[[#fn:r812|812]]</sup> ). Dust mineralogy varies depending on the native soils, so global databases were developed to characterise the mineralogical composition of soils for use in weather and climate models (Journet et al. 2014 <sup>[[#fn:r813|813]]</sup> ; Perlwitz et al. 2015 <sup>[[#fn:r814|814]]</sup> ). New field campaigns, as well as new analyses of observations from prior campaigns, have produced insights into the role of dust in western Africa in climate system, such as long-ranged transport of dust across the Atlantic (Groß et al. 2015 <sup>[[#fn:r815|815]]</sup> ) and the characterisation of aerosol particles and their ability to act as ice and cloud condensation nuclei (Price et al. 2018 <sup>[[#fn:r816|816]]</sup> ). Size distribution at emission is another key parameter controlling dust interactions with radiation. Most models now use the parametrisation of Kok (2011) <sup>[[#fn:r817|817]]</sup> based on the theory of brittle material. It was shown that most models underestimate the size of the global dust cycle (Kok 2011 <sup>[[#fn:r818|818]]</sup> ).

Characterisation of spatial and temporal distribution of dust emissions is essential for weather prediction and climate projections ( ''high confidence'' ). Although there is a growing confidence in characterising the seasonality and peak of dust emissions (i.e., spring–summer (Wang et al. 2015 <sup>[[#fn:r819|819]]</sup> )) and how the meteorological and soil conditions control dust sources, an understanding of long-term future dust dynamics, inter-annual dust variability and how they will affect future climate still requires substantial work. Dust is also important at high latitude, where it has an impact on snow-covered surface albedo and weather (Bullard et al. 2016 <sup>[[#fn:r820|820]]</sup> ).

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<span id="effects-of-past-climate-change-on-dust-emissions-and-feedbacks"></span>
==== 2.4.1.2 Effects of past climate change on dust emissions and feedbacks ====

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A limited number of model-based studies found that dust emissions have increased significantly since the late 19th century: by 25% from the preindustrial period to the present day (e.g., from 729 Tg yr <sup>–1</sup> to 912 Tg yr <sup>–1</sup> ) with about 50% of the increase driven by climate change and about 40% driven by land use cover change, such as conversion of natural land to agriculture ( ''low confidence'' ) (Stanelle et al. 2014 <sup>[[#fn:r821|821]]</sup> ). These changes resulted in a clear sky radiative forcing at the top of the atmosphere of –0.14 W m <sup>–2</sup> (Stanelle et al. 2014 <sup>[[#fn:r822|822]]</sup> ). The authors found that, in North Africa, most dust is of natural origin, with a recent 15% increase in dust emissions attributed to climate change. In North America two-thirds of dust emissions take place on agricultural lands and both climate change and land-use change jointly drive the increase; between the pre-industrial period and the present day, the overall effect of changes in dust was –0.14 W m <sup>–2</sup> cooling of clear sky net radiative forcing on top of the atmosphere, with –0.05 W m <sup>–2</sup> from land use and –0.083 W m <sup>–2</sup> from changes in climate.

The comparison of observations for vertically integrated mass of atmospheric dust per unit area (i.e., dust mass path (DMP)) obtained from the remotely sensed data and the DMP from CMIP5 models reveal that the model-simulate range of DMP was much lower than the estimates (Evan et al. 2014 <sup>[[#fn:r823|823]]</sup> ). ESMs typically do not reproduce inter- annual and longer timescales variability seen in observations (Evan et al. 2016 <sup>[[#fn:r824|824]]</sup> ). Analyses of the CMIP5 models (Evan 2018 <sup>[[#fn:r825|825]]</sup> ; Evan et al. 2014 <sup>[[#fn:r826|826]]</sup> ) reveal that all climate models systematically underestimate dust emissions, the amount of dust in the atmosphere and its inter- annual variability ( ''medium confidence'' ).

One commonly suggested reason for the lack of dust variability in climate models is the models’ inability to simulate the effects of land surface changes on dust emission (Stanelle et al. 2014 <sup>[[#fn:r827|827]]</sup> ). Models that account for changes in land surface show more agreement with the satellite observations both in terms of aerosol optical depth and DMP (Kok et al. 2014 <sup>[[#fn:r828|828]]</sup> ). New prognostic dust emissions models are now able to account for both changes in surface winds and vegetation characteristics (e.g., leaf area index and stem area index) and soil water, ice and snow cover (Evans et al. 2016 <sup>[[#fn:r829|829]]</sup> ). As a result, new modelling studies (e.g., Evans et al. 2016 <sup>[[#fn:r830|830]]</sup> ) indicate that, in regions where soil and vegetation respond strongly to ENSO events, such as in Australia, inclusion of dynamic vegetation characteristics into dust emission parameterisations improves comparisons between the modelled and observed relationship with long-term climate variability (e.g., ENSO) and dust levels (Evans et al. 2016 <sup>[[#fn:r831|831]]</sup> ). Thus, there has been progress in incorporating the effects of vegetation, soil moisture, surface wind and vegetation on dust emission source functions, but the number of studies demonstrating such improvement remains small ( ''limited evidence, medium agreement'' ).

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==== 2.4.1.3 Future changes of dust emissions ====

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There is no agreement about the direction of future changes in dust emissions. Atmospheric dust loading is projected to increase over the southern edge of the Sahara in association with surface wind and precipitation changes (Pu and Ginoux, 2018 <sup>[[#fn:r832|832]]</sup> ), while Evan et al. (2016) <sup>[[#fn:r833|833]]</sup> project a decline in African dust emissions. Dust optical depth (DOD) is also projected to increase over the central Arabian peninsula in all seasons, and to decrease over northern China from March-April-May to September-October-November (Pu and Ginoux 2018 <sup>[[#fn:r834|834]]</sup> ). Climate models project rising drought risks over the south-western and central US in the 21st century. The projected drier regions largely overlay the major dust sources in the US. However, whether dust activity in the US will increase in the future is not clear, due to the large uncertainty in dust modelling (Pu and Ginoux 2017 <sup>[[#fn:r835|835]]</sup> ). Future trends of dust emissions will depend on changes in precipitation patterns and atmospheric circulation ( ''limited evidence, high agreement'' ). However, implication of changes in human activities, including mitigation (e.g., bioenergy production) and adaption (e.g., irrigation) are not characterised in the current literature.

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=== 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> ).

<div id="section-2-4-2-3-future-changes-of-carbonaceous-aerosol-emissions"></div>

<span id="future-changes-of-carbonaceous-aerosol-emissions"></span>
==== 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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=== 2.4.3 Biogenic volatile organic compounds ===

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BVOCs are emitted in large amounts by forests (Guenther et al. 2012 <sup>[[#fn:r910|910]]</sup> ). They include isoprene, terpenes, alkanes, alkenes, alcohols, esters, carbonyls and acids (Peñuelas and Staudt 2010 <sup>[[#fn:r911|911]]</sup> ; Guenther et al. 1995 <sup>[[#fn:r912|912]]</sup> , 2012 <sup>[[#fn:r913|913]]</sup> ). Their emissions represent a carbon loss to the ecosystem, which can be up to 10% of the carbon fixed by photosynthesis under stressful conditions (Bracho-Nunez et al. 2011 <sup>[[#fn:r914|914]]</sup> ). The global average emission for vegetated surfaces is 0.7g C m <sup>–2</sup> yr <sup>–1</sup> but can exceed 100 g C m <sup>–2</sup> yr <sup>–1</sup> in some tropical ecosystems (Peñuelas and Llusià 2003 <sup>[[#fn:r915|915]]</sup> ).

<div id="section-2-4-3-1-bvoc-precursors-of-short-lived-climate-forcers-from-land"></div>

<span id="bvoc-precursors-of-short-lived-climate-forcers-from-land"></span>
==== 2.4.3.1 BVOC precursors of short-lived climate forcers from land ====

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BVOCs are rapidly oxidised in the atmosphere to form less volatile compounds that can condense and form SOA. In boreal and tropical forests, carbonaceous aerosols originate from BVOC oxidation, of which isoprene and terpenes are the most important precursors (Claeys et al. 2004 <sup>[[#fn:r916|916]]</sup> ; Hu et al. 2015 <sup>[[#fn:r917|917]]</sup> ; De Sá et al. 2017 <sup>[[#fn:r918|918]]</sup> ; de Sá et al. 2018 <sup>[[#fn:r919|919]]</sup> ; Liu et al. 2016b <sup>[[#fn:r920|920]]</sup> ). See the following sub-section for more detail.

BVOCs are the most important precursors of SOA. The transformation process of BVOCs affects the aerosol size distribution both by contributing to new particle formation and to the growth of larger pre-existing particles. SOA affects the scattering of radiation by the particles themselves (direct aerosol effect), but also changes the amount of CCN and the lifetime and optical properties of clouds (indirect aerosol effect).

High amounts of SOA are observed over forest areas, in particular in boreal and tropical regions where they have been found to mostly originate from BVOC emissions (Manish et al. 2017 <sup>[[#fn:r921|921]]</sup> ). In particular, isoprene epoxydiol-derived SOA (IEPOX-SOA) is being identified in recent studies in North America and Amazonian forest as a major component in the oxidation of isoprene (Allan et al. 2014 <sup>[[#fn:r922|922]]</sup> ; Schulz et al. 2018 <sup>[[#fn:r923|923]]</sup> ; De Sá et al. 2017 <sup>[[#fn:r924|924]]</sup> ). In tropical regions, BVOCs can be convected up to the upper atmosphere, where their volatility is reduced and where they become SOA. In some cases those particles are transported back to the lower atmosphere (Schulz et al. 2018 <sup>[[#fn:r925|925]]</sup> ; Wang et al. 2016a <sup>[[#fn:r926|926]]</sup> ; Andreae et al. 2018 <sup>[[#fn:r927|927]]</sup> ). In the upper troposphere in the Amazon, SOA are important CCN and are responsible for the vigorous hydrological cycle (Pöhlker et al. 2018 <sup>[[#fn:r928|928]]</sup> ). This strong link between BVOC emissions by plants and the hydrological cycle has been discussed in a number of studies (Fuentes et al. 2000 <sup>[[#fn:r929|929]]</sup> ; Schmale et al. 2018 <sup>[[#fn:r930|930]]</sup> ; Pöhlker et al. 2018 <sup>[[#fn:r931|931]]</sup> , 2016 <sup>[[#fn:r932|932]]</sup> ).

Changing BVOC emissions also affect the oxidant concentrations in the atmosphere. Their impact on the concentration of ozone depends on the NOx concentrations. In polluted regions, high BVOC emissions lead to increased production of ozone, followed by the formation of more OH and a reduction in the methane lifetime. In more pristine regions (NOx-limited), increasing BVOC emissions instead lead to decreasing OH and ozone concentrations, resulting in a longer methane lifetime. The net effect of BVOCs then can change over time if NOx emissions are changing.

BVOCs’ possible climate effects have received little attention because it was thought that their short lifetime would preclude them from having any significant direct influence on climate (Unger 2014a <sup>[[#fn:r933|933]]</sup> ; Sporre et al. 2019 <sup>[[#fn:r934|934]]</sup> ). Higher temperatures and increased CO <sub>2</sub> concentrations are (separately) expected to increase the emissions of BVOCs (Jardine et al. 2011 <sup>[[#fn:r935|935]]</sup> , 2015 <sup>[[#fn:r936|936]]</sup> ; Fuentes et al. 2016 <sup>[[#fn:r937|937]]</sup> ). This has been proposed to initiate negative climate feedback mechanisms through increased formation of SOA (Arneth et al. 2010 <sup>[[#fn:r938|938]]</sup> ; Kulmala 2004 <sup>[[#fn:r939|939]]</sup> ; Unger et al. 2017 <sup>[[#fn:r940|940]]</sup> ). More SOA can make clouds more reflective, which can provide a cooling effect. Furthermore, the increase in SOA formation has also been proposed to lead to increased aerosol scattering, resulting in an increase in diffuse radiation. This could boost GPP and further increase BVOC emissions (Kulmala et al. 2014 <sup>[[#fn:r941|941]]</sup> ; Cirino et al. 2014 <sup>[[#fn:r942|942]]</sup> ; Sena et al. 2016 <sup>[[#fn:r943|943]]</sup> ; Schafer et al. 2002 <sup>[[#fn:r944|944]]</sup> ; Ometto et al. 2005 <sup>[[#fn:r945|945]]</sup> ; Oliveira et al. 2007 <sup>[[#fn:r946|946]]</sup> ). This important feedback is starting to emerge (Sporre et al. 2019 <sup>[[#fn:r947|947]]</sup> ; Kulmala 2004 <sup>[[#fn:r948|948]]</sup> ; Arneth et al. 2017 <sup>[[#fn:r949|949]]</sup> ). However, there is evidence that this influence might be significant at different spatial scales, from local to global, through aerosol formation and through direct and indirect greenhouse effects (l ''imited evidence, medium agreement'' ). Most tropical forest BVOCs are primarily emitted from tree foliage, but soil microbes can also be a major source of some compounds including sesquiterpenes (Bourtsoukidis et al. 2018 <sup>[[#fn:r950|950]]</sup> ).

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==== 2.4.3.2 Historical changes of BVOCs and contribution to climate change ====

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Climate warming over the past 30 years, together with the longer growing season experienced in boreal and temperate environments, have increased BVOC global emissions since the preindustrial times ( ''limited evidence, medium agreement'' ) (Peñuelas 2009 <sup>[[#fn:r951|951]]</sup> ; Sanderson et al. 2003 <sup>[[#fn:r952|952]]</sup> ; Pacifico et al. 2012 <sup>[[#fn:r953|953]]</sup> ). This was opposed by lower BVOC emissions caused by the historical conversion of natural vegetation and forests to cropland ( ''limited evidence, medium agreement'' ) (Unger 2013 <sup>[[#fn:r954|954]]</sup> , 2014a <sup>[[#fn:r955|955]]</sup> ; Fu and Liao 2014 <sup>[[#fn:r956|956]]</sup> ). The consequences of historical anthropogenic land cover change were a decrease in the global formation of SOA (–13%) (Scott et al. 2017 <sup>[[#fn:r957|957]]</sup> ) and tropospheric burden (–13%) (Heald and Geddes 2016 <sup>[[#fn:r958|958]]</sup> ). This has resulted in a positive radiative forcing (and thus warming) from 1850–2000 of 0.017 W m <sup>–2</sup> (Heald and Geddes 2016 <sup>[[#fn:r959|959]]</sup> ), 0.025 W m <sup>–2</sup> (Scott et al. 2017 <sup>[[#fn:r960|960]]</sup> ) and 0.09 W m <sup>–2</sup> (Unger 2014b <sup>[[#fn:r961|961]]</sup> ) through the direct aerosol effect. In present-day conditions, global SOA production from all sources spans between 13 and 121 Tg yr <sup>–1</sup> (Tsigaridis et al. 2014 <sup>[[#fn:r962|962]]</sup> ). The indirect aerosol effect (change in cloud condensation nuclei), resulting from land use induced changes in BVOC emissions, adds an additional positive radiative forcing of 0.008 W m <sup>–2</sup> (Scott et al. 2017 <sup>[[#fn:r963|963]]</sup> ). More studies with different model setups are needed to fully assess this indirect aerosol effect associated with land use change from the preindustrial to present. CMIP6 global emissions pathways (Hoesly et al. 2018 <sup>[[#fn:r964|964]]</sup> ; Gidden et al. 2018 <sup>[[#fn:r965|965]]</sup> ) estimates global VOCs emissions in 2015 at 230 MtVOC yr <sup>–1</sup> . They also estimated that, from 2000–2015, emissions were up from 200–230 MtVOC yr <sup>–1</sup> .

There is ( ''limited evidence, medium agreement'' ) that historical changes in BVOC emissions have also impacted on tropospheric ozone. At most surface locations where land use has changed, the NOx concentrations are sufficiently high for the decrease in BVOC emissions to lead to decreasing ozone concentrations (Scott et al. 2017 <sup>[[#fn:r966|966]]</sup> ). However, in more pristine regions (with low NOx concentrations), the imposed conversion to agriculture has increased ozone through decreased BVOC emissions and their subsequent decrease in OH (Scott et al. 2017 <sup>[[#fn:r967|967]]</sup> ; Heald and Geddes 2016 <sup>[[#fn:r968|968]]</sup> ). In parallel, the enhanced soil NOx emissions from agricultural land can increase the ozone concentrations in NOx limited regions (Heald and Geddes 2016 <sup>[[#fn:r969|969]]</sup> ).

Another impact of the historical decrease in BVOC emissions is the reduction in the atmospheric lifetime of methane ( ''limited evidence, medium agreement'' ), which results in a negative radiative forcing that ranges from –0.007 W m <sup>–2</sup> (Scott et al. 2017 <sup>[[#fn:r970|970]]</sup> ) to –0.07 W m <sup>–2</sup> (Unger 2014b <sup>[[#fn:r971|971]]</sup> ). However, knowledge of the degree that BVOC emissions impact on oxidant concentrations, in particular OH (and thus methane concentrations), is still limited and therefore these numbers are very uncertain (Heald and Spracklen 2015 <sup>[[#fn:r972|972]]</sup> ; Scott et al. 2017 <sup>[[#fn:r973|973]]</sup> ). The effect of land use change on BVOC emissions are highly heterogeneous (Rosenkranz et al. 2015 <sup>[[#fn:r974|974]]</sup> ) and though the global values of forcing described above are small, the local or regional values can be higher, and even of opposite sign, than the global values.

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==== 2.4.3.3 Future changes of BVOCs ====

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Studies suggest that increasing temperature will change BVOC emissions through change in species composition and rate of BVOC production. A further 2°C–3°C rise in the mean global temperature could increase BVOC global emissions by an additional 30–45% (Peñuelas and Llusià 2003 <sup>[[#fn:r975|975]]</sup> ). In two modelling studies, the impact on climate from rising BVOC emissions was found to become even larger with decreasing anthropogenic aerosol emissions (Kulmala et al. 2013 <sup>[[#fn:r976|976]]</sup> ; Sporre et al. 2019 <sup>[[#fn:r977|977]]</sup> ). A negative feedback on temperature, arising from the BVOC-induced increase in the first indirect aerosol effect, has been estimated by two studies to be in the order of –0.01 W m <sup>–2</sup> K (Scott et al. 2018b <sup>[[#fn:r978|978]]</sup> ; Paasonen et al. 2013 <sup>[[#fn:r979|979]]</sup> ). Enhanced aerosol scattering from increasing BVOC emissions has been estimated to contribute to a global gain in BVOC emissions of 7% (Rap et al. 2018 <sup>[[#fn:r980|980]]</sup> ). In a warming planet, BVOC emissions are expected to increase but magnitude of this increase is unknown and will depend on future land use change, in addition to climate ( ''limited evidence, medium agreement'' ).

There is a very limited number of studies investigating the climate impacts of BVOCs using future land use scenarios (Ashworth et al. 2012 <sup>[[#fn:r981|981]]</sup> ; Pacifico et al. 2012 <sup>[[#fn:r982|982]]</sup> ). Scott et al. (2018a) <sup>[[#fn:r983|983]]</sup> found that a future deforestation according to the land use scenario in RCP8.5 leads to a 4% decrease in BVOC emissions at the end of the century. This resulted in a direct aerosol forcing of +0.006 W m <sup>–2</sup> (decreased reflection by particles in the atmosphere) and a first indirect aerosol forcing of –0.001 W m <sup>–2</sup> (change in the amount of CCN). Studies not including future land use scenarios but investigating the climate feedbacks leading to increasing future BVOC emissions, have found a direct aerosol effect of –0.06 W m <sup>–2</sup> (Sporre et al. 2019 <sup>[[#fn:r984|984]]</sup> ) and an indirect aerosol effect of –0.45 W m <sup>–2</sup> (Makkonen et al. 2012 <sup>[[#fn:r985|985]]</sup> ; Sporre et al. 2019 <sup>[[#fn:r2134|2134]]</sup> ). The stronger aerosol effects from the feedback compared to the land use are, at least partly, explained by a much larger change in the BVOC emissions.

A positive climate feedback could happen in a future scenario with increasing BVOC emissions, where higher ozone and methane concentrations could lead to an enhanced warming which could further increase BVOC emissions (Arneth et al. 2010 <sup>[[#fn:r986|986]]</sup> ). This possible feedback is mediated by NOx levels. One recent study including dynamic vegetation, land use change, CO <sub>2</sub> and climate change found no increase and even a slight decrease in global BVOC emissions at the end of the century (Hantson et al. 2017 <sup>[[#fn:r987|987]]</sup> ). There is a lack of understanding concerning the processes governing the BVOC emissions, the oxidation processes in the atmosphere, the role of the BVOC oxidation products in new particle formation and particle growth, as well as general uncertainties in aerosol–cloud interactions. There is a need for continued research into these processes, but the current knowledge indicates that changing BVOC emissions need to be taken into consideration when assessing the future climate and how land use will affect it. In summary, the magnitude and sign of net effect of BVOC emissions on the radiation budget and surface temperature is highly uncertain.

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