3. Natural terrestrial methane sources in the Arctic
3.5 Quantification of carbon stocks and spatial extent
3.5.1
Size and characteristics of the Arctic soil carbon reservoir
This section addresses organic carbon stocks in Arctic soils where they matter for potential changes in methane release.
For methane to be produced in soils (and then emitted to the atmosphere) there needs to be an organic carbon source.
Estimating the amount of carbon stored in Arctic soils is therefore critical to projections of future methane emissions in a warming Arctic. Low temperatures, wet conditions and permafrost (soil or peat that is frozen for at least two consecutive years) are favorable for the accumulation and preservation of organic matter in soils as the decomposition rates are limited. In the northern hemisphere, permafrost soils occur on about 25% of the land area (23 million km2), either as a continuous cover (continuous permafrost, >90% of the area) or patchy (discontinuous, 50–90% of the surface) and sporadic permafrost (10–50% of the surface) (Brown et al. 2014).
Estimates of the amount of global organic carbon in Arctic soils have been revised upward recently, amounting to about 50% of the world’s global soil carbon (Tarnocai et al. 2009).
Decomposition of this carbon in a rapidly warming Arctic, and the resulting emissions of carbon dioxide or methane, is a potentially important feedback to climate warming. However, the extent to which this carbon is available for decomposition is dependent on its conservation in frozen ground and vulnerability to permafrost degradation, burial depth, and how easily the organic material is decomposed (McGuire et al. 2010; Van Huissteden and Dolman 2013). Cryoturbation (vertical movement of soil resulting from freeze-thaw Table 3.1 Summary of average ground-based observational flux estimates (120 published studies) on a per square meter basis (McGuire et al. 2012) from sites shown in Fig. 3.4. Emissions to the atmosphere in g CH4 m2 per summer or per year. ‘Summer’ is defined by the individual studies which cover various lengths. Site year is the total number of measurement years at the collected sites analyzed by McGuire et al. (2012) and used to obtain the mean flux.
Time period Wet tundra Dry/mesic tundra
Mean Site year estimates 95% confidence
interval Mean Site year estimates 95% confidence interval
Summer 12.3 38 7.2–17.3 1.1 25 0.4–1.9
Annual 19.5 22 11.3–27 3.1 24 0.4–5.7
processes) mixes carbon to deeper levels in the soil, thereby potentially removing it from layers of rapid decomposition (Kaiser et al. 2007; Koven et al. 2009). On the other hand, permafrost thaw may cause erosion processes and soil subsidence resulting in lake and pond formation and erosion, processes that can expose soil carbon to either anaerobic decomposition causing methane and carbon dioxide emission, aerobic decomposition resulting in carbon dioxide emission, or transport as dissolved and particulate organic carbon to rivers and streams to lakes and the sea (Van Huissteden and Dolman 2013; Vonk and Gustafsson 2013). The latter processes have been neglected in considerations of Arctic tundra carbon cycling but have attracted increased attention recently (Vonk and Gustafsson 2013; Cory et al. 2014).
Older inventories of Arctic soil carbon included only the top 100 cm of soil; however, as permafrost thaw may affect soil organic matter (SOM) at greater depths, Tarnocai et al. (2009) also included soil layers up to 300 cm depth. Tarnocai et al. (2009) also regionalized soils and organic deposits across the Arctic,
making a subdivision between peatland areas, alluvial deposits, and yedoma. The highest organic carbon contents are found in peat soils and peaty, cryoturbated mineral soils (32.2–69.6 kg/m2). Although yedoma soils have a low carbon content, due to the vast extent of these deposits, this adds up to a considerable carbon pool (Zimov et al. 2006; Schirrmeister et al. 2010, 2011).
The most recent estimates for Arctic soil carbon stocks converge on a range between 1400 and 1850 Pg C for all northern permafrost soils (750–1024 Pg C for peatlands, 200–450 Pg C for yedoma and 241–250 Pg C for alluvial deposits). However, the uncertainties associated with these estimates are large, in particular for carbon content at depth and thickness of deposits.
To reduce the uncertainties, the Northern Circumpolar Soil Database is being developed (Hugelius et al. 2013; Fig. 3.5).
Hugelius et al. (2014) arrived at lower estimates in particular for yedoma (178±143 Pg C) and alluvial deposits (31±39 Pg C) (Table 3.2); the yedoma estimate being in line with the sediment analysis data of Strauss et al. (2013).
Table 3.2 Recent estimates of Arctic soil carbon, Pg C.
Source 0–100 cm 0–300 cm >300 cm >300 cm
(delta/ alluvial) Total
Tarnocai et al. 2009 496 1024 407 241 1672
Schuur et al. 2008; McGuire et al. 2009 Not determined 750 400 250 1400–1850
Hugelius et al. 2014 472±34 1034±183,
1104±133 178 +140/-146 31±39 1300–1370 (uncertainty range: 930–1690) Fig. 3.5 Carbon distribution of northern permafrost soils derived from the Northern Circumpolar Soils Database (Hugelius et al. 2013, 2014).
Soil organic carbon content at 0-1 m depth, kg C/m2
0.1 - 10 10 - 25 25 - 50 50 - 100
> 100
Th e decomposability (lability) of SOM strongly determines the rate at which carbon stored in soil reservoirs can be transferred to the atmosphere upon thaw and shows considerable variability.
Schädel et al. (2014) analyzed the lability of SOM using aerobic incubation experiments on organic and mineral soil cores collected from Alaska and northern Siberia. Th e fraction of SOM that turns over in less than a year (‘fast pool’) was less than 5% for all soils. However, the ‘slow pool’ (defi ned here by a turnover time of 5–15 years) varied between organic and mineral soils, with organic soils showing the highest values and highest variability. Th e carbon/nitrogen ratio was a good proxy for the slow pool size. Treat et al. (2015) analyzed a large database of anaerobic incubations. Th ese showed diff erences in the anaerobic CO2:CH4 production ratio (lowest for tundra sites), and overall anaerobic carbon dioxide and methane production (greatest for organic soils and inundated soils, and least for deeper horizons). Methaneproduction was more than four times greater in soil from graminoid (grass) and shrub-dominated sites than in soils from forested sites, indicating that the vegetation community can infl uence methane fl uxes considerably, as also shown by fi eld observations (e.g. Ström et al. 2003; Turetsky et al. 2007). In the absence of any other changes a shift in graminoid species composition may change the CO2:CH4 production ratio (Ström et al. 2003). Pedersen et al. (2011) looked across a range of drained thermokarst lake basins of various ages in northern Alaska, and found evidence of substantial decomposable deposits of carbon in even the oldest lake beds (5500 years) and lower soil horizons.
3.5.2
Vulnerability of the Arctic soil carbon reservoir
Vulnerability of the Arctic soil carbon pool to climate change depends on its position with respect to the surface. Deeper (below 1 m) permafrost carbon is less vulnerable to short-term changes in temperature. A distinction can be made between gradual changes (such as changes in active layer thickness, soil wetness, and microbial process rates), and rapid ‘pulse disturbances’ (such as wildfi res and thermal disturbance of permafrost – thermokarst, Fig. 3.6) (Grosse et al. 2011). Th e latter strongly depend on the distribution of ice in the subsoil (Fig. 3.6) and geomorphological processes resulting from soil subsidence and erosion. Th e fate of the carbon that is lost from soils is either emitted as carbon dioxide and methane to the atmosphere, or transported as dissolved and particulate organic
carbon in rivers. Approximately two-thirds of this river-borne carbon is outgassed to the atmosphere during transport, but a considerable part may be sequestered again in lake, river, and marine sediments (Van Huissteden et al. 2013).
In an ‘expert opinion’ survey conducted in 2012, experts were asked to provide quantitative estimates of permafrost change in response to four scenarios of warming (Schuur et al. 2013). For the highest warming scenario (RCP8.5), experts hypothesized that carbon release from permafrost zone soils could be 19–45 Pg C by 2040, 162–288 Pg C by 2100, and 381–616 Pg C by 2300 in CO2 equivalent using 100-year methane global warming potential (GWP). Th e values become 50% larger using 20-year methane GWP, with a third to a half of expected climate forcing coming from methane even though methane accounted for only 2.3% of the expected carbon release. Experts projected that two-thirds of this release could be avoided under the lowest warming scenario (RCP2.6) (Schuur et al. 2013).
Both temperature and precipitation changes may affect permafrost soils. Ijima et al. (2010) reported increases in soil temperature and soil wetness after four years of high rainfall and snowfall in the central Lena basin in Siberia.
Th ese changes resulted locally in taiga forest die-back (Ijima et al. 2014) and decreases in boreal forest carbon sequestration.
On the other hand, water-limited plant communities in High Arctic environments may benefi t from soil moisture increases (Elberling et al. 2008). Increased active layer thickness enhances decomposition of older soil carbon as shown by climate manipulation experiments (e.g. Natali et al. 2014). For example, Dorrepaal et al. (2009) report that more than 69% of the increase in soil respiration was attributed to SOM near the base of the active layer in a warming experiment in a subarctic peatland.
While expansion of existing lakes may increase in a warmer and wetter climate, it is ultimately limited by fl uvial and subsurface drainage of lakes (Jones et al. 2011; Van Huissteden et al. 2011).
Th ere is a very limited body of reliable data on lake expansion, and even less so on how this may relate to changes in lake methane emissions, since it requires extensive multi-year high resolution remote sensing studies and has to take into account any non-climatic lake level changes (Plug et al. 2008;
Jones et al. 2011). In southern discontinuous permafrost areas, lake area tends to decrease by regrowth (fi lling in of lakes to become wetlands) and subsurface drainage (Roach et al. 2011).
Observation data from the Seward Peninsula (Alaska) indicate a net decrease of lake area resulting from the drainage of Fig. 3.6 How carbon is lost from Arctic soils with permafrost (fi re excluded). Th e bottom bar indicates various modes of soil carbon transfer. Area 1 represents an area with spatially homogeneous active layer thickness increase and Area 2 represents spatially heterogeneous permafrost thawing driven by differences in soil ice content.
Aft er Van Huissteden and Dolman (2013). Terms defi ned in Box 3.1.
E (erosion) A (aerobic) B (anaerobic) E A B
thaw/thermokarst lake
large lakes, while the number of smaller lakes and ponds is growing rapidly (Jones et al. 2011). In other areas there are also indications of rapidly increasing numbers of smaller lakes and ponds (Jorgenson et al. 2006). In western Siberia, a relationship was established between lake size and lake water carbon dioxide and methane concentration, with the smaller lakes showing the highest gas concentrations (Shirakova et al. 2013). This pattern indicates that small features below the resolution of current lake and wetland databases may be important controls on carbon transfers from permafrost soils to the atmosphere.
3.6