Editing IPCC:AR6/WGI/Chapter-5 (section)

==== 5.4.9.1 Assessment of Biogeochemical Tipping Points ====

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===== 5.4.9.1.1 Forest dieback =====

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Published examples of abrupt biogeochemical changes in models include tropical rain forest dieback ( [[#Cox--2004|Cox et al., 2004]] ; [[#Jones--2009|Jones et al., 2009]] ; [[#Brando--2014|Brando et al., 2014]] ; [[#Le%20Page--2017|Le Page et al., 2017]] ; [[#Zemp--2017|Zemp et al., 2017]] ), and temperate and boreal forest dieback ( [[#Joos--2001|Joos et al., 2001]] ; [[#Lucht--2006|Lucht et al., 2006]] ; [[#Scheffer--2012|Scheffer et al., 2012]] ; [[#Lasslop--2016|Lasslop et al., 2016]] ; [[#5.4.3|Section 5.4.3]] ). Such transitions may be related to: (i) large-scale changes in mean climate conditions crossing particular climate thresholds ( [[#Joos--2001|Joos et al., 2001]] ; [[#Cox--2004|Cox et al., 2004]] ; [[#Lucht--2006|Lucht et al., 2006]] ; [[#Hirota--2011|Hirota et al., 2011]] ; [[#Scheffer--2012|Scheffer et al., 2012]] ; [[#Le%20Page--2017|Le Page et al., 2017]] ; [[#Zemp--2017|Zemp et al., 2017]] ); (ii) temperature and precipitation extremes ( [[#Staver--2011|Staver et al., 2011]] ; [[#Higgins--2012|Higgins and Scheiter, 2012]] ; [[#Scheffer--2012|Scheffer et al., 2012]] ; [[#Pavlov--2015|Pavlov, 2015]] ; [[#Zemp--2017|Zemp et al., 2017]] ); or (iii) possible enhancement and intermittency in fire activity ( [[#Staver--2011|Staver et al., 2011]] ; [[#Higgins--2012|Higgins and Scheiter, 2012]] ; [[#Lasslop--2016|Lasslop et al., 2016]] ; [[#Brando--2020|Brando et al., 2020]] ). Simulated changes in forest cover are a combination of the effects of CO <sub>2</sub> on photosynthesis and water-use efficiency ( [[#5.4.1|Section 5.4.1]] ), and the effects of climate change on photosynthesis, respiration and disturbance ( [[#5.4.3|Section 5.4.3]] ). In ESMs, direct CO <sub>2</sub> effects tend to enhance forest growth, but the impacts of climate change vary between being predominantly negative in the tropics and predominantly positive in the boreal zone (Figure 5.27).

Most ESMs project continuing carbon accumulation in tropical forests as a result of direct CO <sub>2</sub> effects overwhelming the negative effects of climate change ( [[#Huntingford--2013|Huntingford et al., 2013]] ; [[#Drijfhout--2015|Drijfhout et al., 2015]] ; [[#Boulton--2017|Boulton et al., 2017]] ). In the real world, forests may be less vulnerable to climate changes than those modelled in ESMs because of the greater plant trait diversity, which confers additional resilience ( [[#Reyer--2015|Reyer et al., 2015]] ; [[#Levine--2016|Levine et al., 2016]] ; [[#Sakschewski--2016|Sakschewski et al., 2016]] ), and because of possible acclimation of vegetation to warming ( [[#Good--2011|Good et al., 2011]] , 2013; [[#Lloret--2012|Lloret et al., 2012]] ; [[#Mercado--2018|Mercado et al., 2018]] ). On the contrary, forests may be more vulnerable in the real world due to indirect climate change effects such as insect outbreaks and diseases not considered here ( [[#5.4.3.2|Section 5.4.3.2]] ) or model limitations in representing the effects disturbances such as wildfire and droughts. In general, forests are most vulnerable when climate change is combined with increased rates of direct deforestation ( [[#Nobre--2016|Nobre et al., 2016]] ; [[#Le%20Page--2017|Le Page et al., 2017]] ).

To estimate an upper limit on the impact of Amazon forest dieback on atmospheric CO <sub>2</sub> , we consider the ''very unlikely'' limiting case of negligible direct-CO <sub>2</sub> effects ( [[#5.4.1|Section 5.4.1]] ). Emergent constraint approaches ( [[#5.4.6|Section 5.4.6]] ) may be used to estimate an overall loss of tropical land carbon due to climate change alone, of around 50 PgC per °C of tropical warming ( [[#Cox--2013|Cox et al., 2013]] ; [[#Wenzel--2014|Wenzel et al., 2014]] ). This implies an upper limit to the release of tropical land carbon of &lt;200 PgC over the 21 <sup>st</sup> century (assuming tropical warming of &lt;4°C '','' and no CO <sub>2</sub> -fertilization), which translates to dCO <sub>2</sub> /dt &lt;0.5 ppm yr <sup>–1</sup> .

Boreal forest dieback is not expected to change the atmospheric CO <sub>2</sub> concentration substantially because forest loss at the south is partly compensated by: (i) temperate forest invasion into previously boreal areas; and (ii) boreal forest gain at the north ( [[#Friend--2014|Friend et al., 2014]] ; [[#Kicklighter--2014|Kicklighter et al., 2014]] ; [[#Schaphoff--2016|Schaphoff et al., 2016]] ) ( ''medium confidence'' ). An upper estimate of this magnitude, based on statistical modelling of climate change alone, is of 27 Pg vegetation carbon loss in the southern boreal forest, which is roughly balanced by gains in the northern zone ( [[#Koven--2013|Koven, 2013]] ). Carbon release from vegetation and soil due to wildfires in boreal regions ( [[#Eliseev--2014b|Eliseev et al., 2014b]] ; [[#Turetsky--2015|Turetsky et al., 2015]] ; X.J. [[#Walker--2019|]] [[#Walker--2019|Walker et al., 2019]] ) is also not expected to change this estimate substantially because of its small present-day value of about 0.2 PgC yr <sup>–1</sup> ( [[#van%20der%20Werf--2017|van der Werf et al., 2017]] ), and because of ''likely'' increases in precipitation in boreal regions ( [[IPCC:Wg1:Chapter:Chapter-4#4.5.1|Section 4.5.1]] ).

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===== 5.4.9.1.2 Biogenic emissions following permafrost thaw =====

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There is large uncertainty in release of GHGs from permafrost in the 21st century. The largest of these estimates implies tens to hundreds of gigatons of carbon released in the form of CO <sub>2</sub> (Box 5.1) and CH <sub>4</sub> emissions up to 100 TgCH <sub>4</sub> yr <sup>–1</sup> (Box 5.1). A carbon dioxide release of such magnitude would lead to an increase in the CO <sub>2</sub> accumulation rate in the atmosphere of ≤1 ppm yr <sup>–1</sup> . These emissions develop at a multi-decadal time scale. Assuming a CH <sub>4</sub> lifetime in the atmosphere of the order of 10 years and the associated feedback parameter of 1.34 ± 0.04 (Section 6.2.2.1), this would increase the atmospheric CH <sub>4</sub> content by about 500 ppb over the century, corresponding to a rate of ≤10 ppb yr <sup>–1</sup> . Irrespective of its origin, additional CH <sub>4</sub> accumulation of such a magnitude is not expected to modify the temperature response to anthropogenic emissions by more than a few tenths of a °C ( [[#Gedney--2004|Gedney et al., 2004]] ; [[#Eliseev--2008|Eliseev et al., 2008]] ; [[#Denisov--2013|Denisov et al., 2013]] ). Emissions from permafrost thawing are assessed in Box 5.1.

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===== 5.4.9.1.3 Methane release from clathrates =====

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The total global clathrate reservoir is estimated to contain 1500–2000 PgC ( [[#Archer--2009|Archer et al., 2009]] ; [[#Ruppel--2017|Ruppel and Kessler, 2017]] ), held predominantly in ocean sediments, with only an estimated 20 PgC in and under permafrost ( [[#Ruppel--2015|Ruppel, 2015]] ). The present-day CH <sub>4</sub> release from shelf clathrates is &lt;10 TgCH <sub>4</sub> yr <sup>–1</sup> ( [[#Kretschmer--2015|Kretschmer et al., 2015]] ; [[#Saunois--2020|Saunois et al., 2020]] ). Despite polar amplification (Chapter 7), substantial releases from the permafrost-embedded subsea clathrates is ''very unlikely'' ( [[#Minshull--2016|Minshull et al., 2016]] ; [[#Malakhova--2017|Malakhova and Eliseev, 2017]] , 2020). This is consistent with an overall small release of CH <sub>4</sub> from the shelf clathrates during the last deglacial transition, despite large reorganizations in climate state ( [[#Bock--2017|Bock et al., 2017]] ; [[#Petrenko--2017|Petrenko et al., 2017]] ; [[#Dyonisius--2020|Dyonisius et al., 2020]] ). The long time scales associated with clathrate destabilization makes it ''unlikely'' that CH <sub>4</sub> release from the ocean to the atmosphere will deviate markedly from the present-day value through the 21st century ( [[#Hunter--2013|Hunter et al., 2013]] ), corresponding to no more than additional 20 ppb of atmospheric CH <sub>4</sub> (i.e., &lt;0.2 ppb yr <sup>–1</sup> ). Another possible source of CH <sub>4</sub> is gas clathrates in deeper terrestrial permafrost and below it ( [[#Buldovicz--2018|Buldovicz et al., 2018]] ; [[#Chuvilin--2018|Chuvilin et al., 2018]] ), which may have caused recent craters in the north of Russia ( [[#Arzhanov--2016|Arzhanov et al., 2016]] , 2020; [[#Arzhanov--2017|Arzhanov and Mokhov, 2017]] ; [[#Kizyakov--2017|Kizyakov et al., 2017]] , 2018). Land clathrates are formed at depths greater than 200 m ( [[#Ruppel--2017|Ruppel and Kessler, 2017]] ; [[#Malakhova--2020|Malakhova and Eliseev, 2020]] ), which precludes a substantial response to global warming over the next few centuries and associated emissions.

Thus, it is ''very unlikely'' that CH <sub>4</sub> emissions from clathrates will substantially warm the climate system over the next few centuries.

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