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

=== 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> ).

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==== 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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