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Estimating Soil Carbon Turnover Using Radiocarbon Data

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Estimating Soil Carbon Turnover Using Radiocarbon Data 1 Estimating soil carbon turnover using radiocarbon data: 2 European Russia case-study 3 4 5 6 7 Victor Brovkin1,*, Alexander Cherkinsky2, Sergey Goryachkin3 8 9 10 11 1Potsdam Institute for Climate Impact Research, P.O.Box 601203, 14412 Potsdam, Germany 12 2Center for Applied Isotope Studies, University of Georgia, 120 Riverbend Rd., Athens, GA 13 30602, USA 14 3Institute of Geography, Russian Academy of Sciences, Staromonetny, 29 Moscow 109017 15 Russia 16 17 *Corresponding author, tel. +49 331 2882592, fax +49 331 2882620, e-mail: 18 [email protected] 19 20 21 1 1 2 Abstract 3 4 Turnover rates of soil carbon for 20 soil types typical for 3.7 million km2 area of European 5 Russia were estimated based on 14C data. The rates are corrected for bomb radiocarbon 6 which strongly affects the topsoil 14C balance. The approach is applied for carbon stored in 7 organic and mineral layers of the upper 1 m of soil profile. Turnover rates of carbon in the 8 upper 20 cm are relatively high for forest soils (0.16-0.78% yr-1), intermediate for tundra 9 soils (0.25% yr-1), and low for grassland soils (0.02-0.08% yr-1) with exception for southern 10 Chernozems (0.32% yr-1). In the soil layer at 20-100 cm depth the turnover rates were much 11 lower for all soil types (0.01-0.06% yr-1) except for peat bog soils of southern taiga (0.14% 12 yr-1). Combined with a map of soil type distribution and a dataset of several hundred soil 13 carbon profiles, the method provides annual fluxes for slowest components of soil carbon 14 assuming that the later is in equilibrium with climate and vegetation cover. Estimated carbon 15 flux from the soil is highest for forest soils (12 to 147 gC/(m2⋅yr)), intermediate for tundra 16 soils (33 gC/(m2⋅yr)), and lowest for grassland soils (1-26 gC/(m2⋅yr)). The approach does 17 not distinguish active and recalcitrant carbon fractions and this explains low turnover rates 18 in the top layer. Since changes in soil types will follow changes in climate and land cover, 19 we suggest that pedogenesis is an important factor influencing future dynamics of soil 20 carbon fluxes. Up to now, an effect of soil type changes, as well a clear evidence from 14C 21 measurements that most of soil organic carbon has millennial time scale are basically st 22 neglected in the global carbon cycle models used for projections of atmospheric CO2 in 21 23 century and beyond. 24 2 1 2 Introduction 3 4 Soil carbon is the main component of terrestrial carbon cycle. Estimates of storages of soil 5 organic matter (SOM) in the upper 1 meter layer vary in the range of 1,500 to 2,000 GtC 6 (Post et al., 1982, 1997; Batjes et al., 1996, Prentice et al., 2001) depending on a way to 7 account for organic carbon storage in wetlands. Soil layers deeper than 1 meter contain 8 several hundred PgC in form of peat in northern ecosystems (e.g. Gorham, 1991) and 9 organic carbon in moist tropical forest soils (Trumbore et al., 1995). Additionally, about 10 950 PgC are stored in inorganic (carbonate) form, predominantly in drylands (Lal, 2004). 11 Ample storages of soil carbon outweigh by a factor of three to six the plant biomass 12 estimated in the range of 470 to 660 PgC (Prentice et al., 2001). In case of continued 13 greenhouse gas emissions, global mean air temperature is projected to increase up to 6.4°C 14 during the 21st century (IPCC-2007, SPM). Consequent drastic changes in plant 15 productivity, soil thermal and hydrological balance will strongly affect terrestrial carbon 16 storage and, through the land-atmosphere CO2 exchange, the atmospheric CO2 17 concentration. The feedback loop between CO2 and climate most likely has amplified 18 climate change in the past (Scheffer et al., 2006, Torn and Harte, 2006) and could 19 substantially increase global warming in the future (Cox et al., 2000, Kirschbaum, 2000). 20 21 Global coupled climate – carbon cycle models are the best tools currently available for 22 assessment of changes in global carbon balance in the future. Within the Coupled Climate– 23 Carbon Cycle Model Intercomparison Project (C4MIP), eleven coupled climate–carbon 24 cycle simulations of different complexity performed simulations over the twenty-first 25 century (Friedlingstein et al., 2006). All but one model simulated a reduction of SOM 26 turnover time, and some models show a strong negative impact of climate change on 27 turnover time (up to 1 yr decrease per 1°C global warming). This assessment is very 28 preliminary because the SOM balance is one of the most crudely represented processes in 29 the global carbon models. One of the biggest uncertainties is a decomposition of the inert 30 (stable or recalcitrant) organic carbon. It is likely that biological processes consume the 31 recalcitrant SOM as well but little is currently known about these processes (Prentice et al., 3 1 2001). An effect of soil temperature change on SOM decomposition rate is doubtless, but its 2 magnitude in long-term dynamics is currently a matter of debate (Knorr et al., 2005, 3 Reichstein et al., 2005, Fang et al., 2006). 4 5 Here, we focus on time scales of the SOM decomposition based on radiocarbon 6 measurements using the soil formation model formulated by Cherkinsky and Brovkin 7 (1993). Similar type of model has been applied by Trumbore (1995), Perruchoud (1996) and 8 Gaudinsky et al. (2000) for evaluation of soil carbon cycling in tropical and temperate 9 forests. Hahn and Buchmann (2004) accounted for two pools (active and passive) based on 10 prescribing a threshold in 14C activity for active carbon in the soil. This method is more 11 advanced than the bulk carbon models mentioned before, but it does require a specification 12 of organic input by aboveground and belowground litter. Since these data were not available 13 for us, we applied hereafter the model by Cherkinsky and Brovkin (1993) of unfractionated 14 SOM and combine it with database on soil carbon storage in different soil types of European 15 Russia. 16 17 18 Methods 19 20 Radiocarbon (14 C) analysis 21 22 The method is based on the fact that due to β-decay the specific carbon activity = 23 I(t) C14 (t) C12 24 of organic matter remaining after the organism’s death obeys the exponential decay: −λ 25 It()= Ae t , (1) 26 where λ is the decay-rate of 14 C (14 C half-time is 5730 years), A is the specific activity of 27 atmospheric carbon and Ct12 (), C14 (t) are the contents of carbon isotopes in organic matter 28 (the content of stable 12 C isotope does not change with time). 29 In accordance with equation (1) −1 I(t) 30 T = ln (2) λ A 4 1 is 14 C age of analyzed organic matter. 2 3 Equation (2) is widely used to estimate the period of time since the organism’s death. This 4 equation implicitly assume a stable 14 C concentration in the atmosphere, which is 5 approximately true at least for the last several thousand years. Equations (1) and (2) refer to 6 a "closed carbon system", which presumes no carbon exchange with the environment. In 7 contrast, the soil carbon represents an open system. The 14 C age calculated from equation 8 (2) can thus not be interpreted as absolute age in the context of SOM (Scharpenseel, 1971) 9 but has the meaning of a mean residence time of soil organic carbon. Its reciprocal is a 10 turnover rate of soil carbon, m: 11 m = 1 T . (3) 12 13 Soil organic matter is a heterogeneous system and its turnover rate depends on the fraction 14 of soil carbon, depth of soil sample, soil type and many other factors (Schimel et al., 1994). 15 16 Model of soil organic profile formation 17 18 The model of a monogenetic soil organic profile formation under stable conditions of 19 pedogenesis is based on the following assumptions: 20 • the soil organic profile develops from the surface downward as a result of the increased 21 involvement of rocks in humus-forming processes; 22 • the underlying layers are formed later than the overlying ones; 23 • rates of organic detritus input and of humus turnover are constant. 24 Under these assumptions, carbon accumulation in the soil can be represented by dC() t 12 =−−pAmCt()1 () dt 12 25 , (4) dC() t 14 =−pA mC() t −λ C () t dt 14 14 26 where p(1− A ) and p A define the amounts of input of 12 C and 14 C respectively, and 27 C12 (t),C14 (t) are the contents of carbon isotopes in soil. Assuming as a first approximation 28 constant inflow and turnover of carbon, we obtain from (4): 5 pA()1− Ct()= ()1− e−mt 12 m 1 , (5) pA Ct()= (1− e−+()mtλ ) 14 m + λ 2 resulting in a specific activity m A − −mt Ct14 () 1 e 3 It()== . (6) +−λ − −+()mtλ Ct12 () (mA )(1 )1 e − 4 Since A ≈ 10 12 , we can replace 1-A with 1 in the equations (5-6). 5 For equilibrium conditions we find from (5): p p A m A 6 C **==,,C I *= , (7) 12m 14 m + λ m + λ 7 where I * is the specific carbon activity for equilibrium case, and A is the specific carbon 8 activity for the input flux of plant detritus.
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