Editing IPCC:AR6/SRCCL/Chapter-3 (section)

== 3.3 Desertification feedbacks to climate ==

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While climate change can drive desertification (Section 3.1.4.1), the process of desertification can also alter the local climate, providing a feedback (Sivakumar 2007 <sup>[[#fn:r454|454]]</sup> ). These feedbacks can alter the carbon cycle, and hence the level of atmospheric CO <sub>2</sub> and its related global climate change, or they can alter the surface energy and water budgets, directly impacting the local climate. While these feedbacks occur in all climate zones (Chapter 2), here we focus on their effects in dryland regions and assess the literature concerning the major desertification feedbacks to climate. The main feedback pathways discussed throughout Section 3.3 are summarised in Figure 3.8.

Drylands are characterised by limited soil moisture compared to humid regions. Thus, the sensible heat (heat that causes the atmospheric temperature to rise) accounts for more of the surface net radiation than latent heat (evaporation) in these regions (Wang and Dickinson 2013 <sup>[[#fn:r455|455]]</sup> ). This tight coupling between the surface energy balance and the soil moisture in semi-arid and dry sub-humid zones makes these regions susceptible to land–atmosphere feedback loops that can amplify changes to the water cycle (Seneviratne et al. 2010 <sup>[[#fn:r456|456]]</sup> ). Changes to the land surface caused by desertification can change the surface energy budget, altering the soil moisture and triggering these feedbacks.

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'''Figure 3.8'''

<span id="schematic-of-main-pathways-through-which-desertification-can-feed-back-on-climate-as-discussed-in-section-3.4.-note-red-arrows-indicate-a-positive-effect.-blue-arrows-indicate-a-negative-effect.-grey-arrows-indicate-an-indeterminate-effect-potentially-both-positive-and-negative.-solid-arrows-are-direct-while-dashed-arrows-are-indirect."></span>
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'''Schematic of main pathways through which desertification can feed back on climate, as discussed in Section 3.4. Note: Red arrows indicate a positive effect. Blue arrows indicate a negative effect. Grey arrows indicate an indeterminate effect (potentially both positive and negative). Solid arrows are direct while dashed arrows are indirect.'''
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[[File:82aed23d08e618f988a53dc1661b871f Figure-3.8-1024x731.jpg]]

Schematic of main pathways through which desertification can feed back on climate, as discussed in Section 3.4. Note: Red arrows indicate a positive effect. Blue arrows indicate a negative effect. Grey arrows indicate an indeterminate effect (potentially both positive and negative). Solid arrows are direct while dashed arrows are indirect.

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=== 3.3.1 Sand and dust aerosols ===

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Sand and mineral dust are frequently mobilised from sparsely vegetated drylands forming ‘sand storms’ or ‘dust storms’ (UNEP et al. 2016 <sup>[[#fn:r457|457]]</sup> ). The African continent is the most important source of desert dust; perhaps 50% of atmospheric dust comes from the Sahara (Middleton 2017 <sup>[[#fn:r458|458]]</sup> ). Ginoux et al. (2012) <sup>[[#fn:r459|459]]</sup> estimated that 25% of global dust emissions have anthropogenic origins, often in drylands. These events can play an important role in the local energy balance. Through reducing vegetation cover and drying the surface conditions, desertification can increase the frequency of these events. Biological or structural soil crusts have been shown to effectively stabilise dryland soils. Thus their loss due to intense land use and/ or climate change can be expected to cause an increase in sand and dust storms ( ''high confidence'' ) (Rajot et al. 2003 <sup>[[#fn:r460|460]]</sup> ; Field et al. 2010 <sup>[[#fn:r461|461]]</sup> ; Rodriguez-Caballero et al. 2018 <sup>[[#fn:r462|462]]</sup> ). These sand and dust aerosols impact the regional climate in several ways (Choobari et al. 2014 <sup>[[#fn:r463|463]]</sup> ). The direct effect is the interception, reflection and absorption of solar radiation in the atmosphere, reducing the energy available at the land surface and increasing the temperature of the atmosphere in layers with sand and dust present (Kaufman et al. 2002 <sup>[[#fn:r464|464]]</sup> ; Middleton 2017 <sup>[[#fn:r465|465]]</sup> ; Kok et al. 2018 <sup>[[#fn:r466|466]]</sup> ). The heating of the dust layer can alter the relative humidity and atmospheric stability, which can change cloud lifetimes and water content. This has been referred to as the semi-direct effect (Huang et al. 2017 <sup>[[#fn:r467|467]]</sup> ). Aerosols also have an indirect effect on climate through their role as cloud condensation nuclei, changing cloud radiative properties as well as the evolution and development of precipitation (Kaufman et al. 2002 <sup>[[#fn:r468|468]]</sup> ). While these indirect effects are more variable than the direct effects, depending on the types and amounts of aerosols present, the general tendency is toward an increase in the number, but a reduction in the size of cloud droplets, increasing the cloud reflectivity and decreasing the chances of precipitation. These effects are referred to as aerosol-radiation and aerosol–cloud interactions (Boucher et al. 2013 <sup>[[#fn:r469|469]]</sup> ).

There is ''high confidence'' that there is a negative relationship between vegetation green-up and the occurrence of dust storms (Engelstaedter et al. 2003 <sup>[[#fn:r470|470]]</sup> ; Fan et al. 2015 <sup>[[#fn:r471|471]]</sup> ; Yu et al. 2015 <sup>[[#fn:r472|472]]</sup> ; Zou and Zhai 2004 <sup>[[#fn:r473|473]]</sup> ). Changes in groundwater can affect vegetation and the generation of atmospheric dust (Elmore et al. 2008 <sup>[[#fn:r474|474]]</sup> ). This can occur through groundwater processes such as the vertical movement of salt to the surface causing salinisation, supply of near-surface soil moisture, and sustenance of groundwater dependent vegetation. Groundwater dependent ecosystems have been identified in many dryland regions around the world (Decker et al. 2013 <sup>[[#fn:r475|475]]</sup> ; Lamontagne et al. 2005 <sup>[[#fn:r476|476]]</sup> ; Patten et al. 2008 <sup>[[#fn:r477|477]]</sup> ). In these locations declining groundwater levels can decrease vegetation cover. Cook et al. (2009) <sup>[[#fn:r478|478]]</sup> found that dust aerosols intensified the ‘Dust Bowl’ drought in North America during the 1930s.

By decreasing the amount of green cover and hence increasing the occurrence of sand and dust storms, desertification will increase the amount of shortwave cooling associated with the direct effect ( ''high confidence'' ). There is ''medium confidence'' that the semi-direct and indirect effects of this dust would tend to decrease precipitation and hence provide a positive feedback to desertification (Huang et al. 2009 <sup>[[#fn:r479|479]]</sup> ; Konare et al. 2008 <sup>[[#fn:r480|480]]</sup> ; Rosenfeld et al. 2001 <sup>[[#fn:r481|481]]</sup> ; Solmon et al. 2012 <sup>[[#fn:r482|482]]</sup> ; Zhao et al. 2015 <sup>[[#fn:r483|483]]</sup> ). However, the combined effect of dust has also been found to increase precipitation in some areas (Islam and Almazroui 2012 <sup>[[#fn:r484|484]]</sup> ; Lau et al. 2009 <sup>[[#fn:r485|485]]</sup> ; Sun et al. 2012 <sup>[[#fn:r486|486]]</sup> ). The overall combined effect of dust aerosols on desertification remains uncertain with low agreement between studies that find positive (Huang et al. 2014 <sup>[[#fn:r487|487]]</sup> ), negative (Miller et al. 2004 <sup>[[#fn:r488|488]]</sup> ) or no feedback on desertification (Zhao et al. 2015 <sup>[[#fn:r489|489]]</sup> ).

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==== 3.3.1.1 Off-site feedbacks ====

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Aerosols can act as a vehicle for the long-range transport of nutrients to oceans (Jickells et al. 2005 <sup>[[#fn:r490|490]]</sup> ; Okin et al. 2011 <sup>[[#fn:r491|491]]</sup> ) and terrestrial land surfaces (Das et al. 2013 <sup>[[#fn:r492|492]]</sup> ). In several locations, notably the Atlantic Ocean, the west of northern Africa, and the Pacific Ocean east of northern China, a considerable amount of mineral dust aerosols, sourced from nearby drylands, reaches the oceans. It was estimated that 60% of dust transported off Africa is deposited in the Atlantic Ocean (Kaufman et al. 2005 <sup>[[#fn:r493|493]]</sup> ), while 50% of the dust generated in Asia reaches the Pacific Ocean or further (Uno et al. 2009 <sup>[[#fn:r494|494]]</sup> ; Zhang et al. 1997 <sup>[[#fn:r495|495]]</sup> ). The Sahara is also a major source of dust for the Mediterranean basin (Varga et al. 2014 <sup>[[#fn:r496|496]]</sup> ). The direct effect of atmospheric dust over the ocean was found to be a cooling of the ocean surface ( ''limited evidence, high agreement'' ) (Evan and Mukhopadhyay 2010 <sup>[[#fn:r497|497]]</sup> ; Evan et al. 2009 <sup>[[#fn:r498|498]]</sup> ) with the tropical North Atlantic mixed layer cooling by over 1°C (Evan et al. 2009 <sup>[[#fn:r499|499]]</sup> ).

It has been suggested that dust may act as a source of nutrients for the upper ocean biota, enhancing the biological activity and related carbon sink ( ''medium'' ''evidence, low agreement'' ) (Lenes et al. 2001 <sup>[[#fn:r500|500]]</sup> ; Shaw et al. 2008 <sup>[[#fn:r501|501]]</sup> ; Neuer et al. 2004 <sup>[[#fn:r502|502]]</sup> ). The overall response depends on the environmental controls on the ocean biota, the type of aerosols including their chemical constituents, and the chemical environment in which they dissolve (Boyd et al. 2010 <sup>[[#fn:r503|503]]</sup> ).

Dust deposited on snow can increase the amount of absorbed solar radiation leading to more rapid melting (Painter et al. 2018 <sup>[[#fn:r504|504]]</sup> ), impacting a region’s hydrological cycle ( ''high confidence'' ). Dust deposition on snow and ice has been found in many regions of the globe (e.g., Painter et al. 2018; Kaspari et al. 2014 <sup>[[#fn:r505|505]]</sup> ; Qian et al. 2015 <sup>[[#fn:r506|506]]</sup> ; Painter et al. 2013 <sup>[[#fn:r507|507]]</sup> ), however quantification of the effect globally and estimation of future changes in the extent of this effect remain knowledge gaps.

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=== 3.3.2 Changes in surface albedo ===

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Increasing surface albedo in dryland regions will impact the local climate, decreasing surface temperature and precipitation, and provide a positive feedback on the albedo ( ''high confidence'' ) (Charney et al. 1975 <sup>[[#fn:r508|508]]</sup> ). This albedo feedback can occur in desert regions worldwide (Zeng and Yoon 2009 <sup>[[#fn:r509|509]]</sup> ). Similar albedo feedbacks have also been found in regional studies over the Middle East (Zaitchik et al. 2007 <sup>[[#fn:r510|510]]</sup> ), Australia (Evans et al. 2017 <sup>[[#fn:r511|511]]</sup> ; Meng et al. 2014a <sup>[[#fn:r512|512]]</sup> , b <sup>[[#fn:r513|513]]</sup> ), South America (Lee and Berbery 2012 <sup>[[#fn:r514|514]]</sup> ) and the USA (Zaitchik et al. 2013 <sup>[[#fn:r515|515]]</sup> ).

Recent work has also found albedo in dryland regions can be associated with soil surface communities of lichens, mosses and cyanobacteria (Rodriguez-Caballero et al. 2018 <sup>[[#fn:r516|516]]</sup> ). These communities compose the soil crust in these ecosystems and due to the sparse vegetation cover, directly influence the albedo. These communities are sensitive to climate changes, with field experiments indicating albedo changes greater than 30% are possible. Thus, changes in these communities could trigger surface albedo feedback processes ( ''limited evidence, high agreement'' ) (Rutherford et al. 2017 <sup>[[#fn:r517|517]]</sup> ).

A further pertinent feedback relationship exists between changes in land-cover, albedo, carbon stocks and associated GHG emissions, particularly in drylands with low levels of cloud cover. One of the first studies to focus on the subject was Rotenberg and Yakir (2010) <sup>[[#fn:r518|518]]</sup> , who used the concept of ‘radiative forcing’ to compare the relative climatic effect of a change in albedo with a change in atmospheric GHGs due to the presence of forest within drylands. Based on this analysis, it was estimated that the change in surface albedo due to the degradation of semi-arid areas has decreased radiative forcing in these areas by an amount equivalent to approximately 20% of global anthropogenic GHG emissions between 1970 and 2005.

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=== 3.3.3 Changes in vegetation and greenhouse gas fluxes ===

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Terrestrial ecosystems have the ability to alter atmospheric GHGs through a number of processes (Schlesinger et al. 1990 <sup>[[#fn:r519|519]]</sup> ). This may be through a change in plant and soil carbon stocks, either sequestering atmospheric CO <sub>2</sub> during growth or releasing carbon during combustion and respiration, or through processes such as enteric fermentation of domestic and wild ruminants that leads to the release of methane and nitrous oxide (Sivakumar 2007 <sup>[[#fn:r520|520]]</sup> ). It is estimated that 241–470 GtC is stored in dryland soils (top 1 m) (Lal 2004 <sup>[[#fn:r521|521]]</sup> ; Plaza et al. 2018 <sup>[[#fn:r522|522]]</sup> ). When evaluating the effect of desertification, the net balance of all the processes and associated GHG fluxes needs to be considered.

Desertification usually leads to a loss in productivity and a decline in above – and below-ground carbon stocks (Abril et al. 2005 <sup>[[#fn:r523|523]]</sup> ; Asner et al. 2003 <sup>[[#fn:r524|524]]</sup> ). Drivers such as overgrazing lead to a decrease in both plant and SOC pools (Abdalla et al. 2018 <sup>[[#fn:r525|525]]</sup> ). While dryland ecosystems are often characterised by open vegetation, not all drylands have low biomass and carbon stocks in an intact state (Lechmere-Oertel et al. 2005 <sup>[[#fn:r526|526]]</sup> ; Maestre et al. 2012 <sup>[[#fn:r527|527]]</sup> ). Vegetation types such as the subtropical thicket of South Africa have over 70 tC ha– <sup>1</sup> in an intact state, greater than 60% of which is released into the atmosphere during degradation through overgrazing (Lechmere-Oertel et al. 2005 <sup>[[#fn:r528|528]]</sup> ; Powell 2009 <sup>[[#fn:r529|529]]</sup> ). In comparison, semi-arid grasslands and savannahs with similar rainfall, may have only 5–35 tC ha– <sup>1</sup> (Scholes and Walker 1993 <sup>[[#fn:r530|530]]</sup> ; Woomer et al. 2004 <sup>[[#fn:r531|531]]</sup> ).

At the same time, it is expected that a decline in plant productivity may lead to a decrease in fuel loads and a reduction in CO <sub>2</sub> , nitrous oxide and methane emissions from fire. In a similar manner, decreasing productivity may lead to a reduction in numbers of ruminant animals that in turn would decrease methane emissions. Few studies have focussed on changes in these sources of emissions due to desertification and it remains a field that requires further research.

In comparison to desertification through the suppression of primary production, the process of woody plant encroachment can result in significantly different climatic feedbacks. Increasing woody plant cover in open rangeland ecosystems leads to an increase in woody carbon stocks both above – and below-ground (Asner et al. 2003 <sup>[[#fn:r532|532]]</sup> ; Hughes et al. 2006 <sup>[[#fn:r533|533]]</sup> ; Petrie et al. 2015 <sup>[[#fn:r534|534]]</sup> ; Li et al. 2016 <sup>[[#fn:r535|535]]</sup> ). Within the drylands of Texas, USA, shrub encroachment led to a 32% increase in aboveground carbon stocks over a period of 69 years (3.8 tC ha– <sup>1</sup> to 5.0 tC ha– <sup>1</sup> ) (Asner et al. 2003 <sup>[[#fn:r536|536]]</sup> ). Encroachment by taller woody species can lead to significantly higher observed biomass and carbon stocks. For example, encroachment by ''Dichrostachys cinerea'' and several Vachellia species in the sub-humid savannahs of north-west South Africa led to an increase of 31–46 tC ha– <sup>1</sup> over a 50–65 year period (1936–2001) (Hudak et al. 2003 <sup>[[#fn:r537|537]]</sup> ). In terms of potential changes in SOC stocks, the effect may be dependent on annual rainfall and soil type. Woody cover generally leads to an increase in SOC stocks in drylands that have less than 800 mm of annual rainfall, while encroachment can lead to a loss of soil carbon in more humid ecosystems (Barger et al. 2011 <sup>[[#fn:r538|538]]</sup> ; Jackson et al. 2002 <sup>[[#fn:r539|539]]</sup> ).

The suppression of the grass layer through the process of woody encroachment may lead to a decrease in carbon stocks within this relatively small carbon pool (Magandana 2016 <sup>[[#fn:r540|540]]</sup> ). Conversely, increasing woody cover may lead to a decrease and even halt in surface fires and associated GHG emissions. In their analysis of drivers of fire in southern Africa, Archibald et al. (2009) <sup>[[#fn:r541|541]]</sup> note that there is a potential threshold around 40% canopy cover, above which surface grass fires are rare. Whilst there have been a number of studies on changes in carbon stocks due to desertification in North America, southern Africa and Australia, a global assessment of the net change in carbon stocks – as well as fire and ruminant GHG emissions due to woody plant encroachment – has not been done yet.

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