3.1. Global vulnerability of SOC to climate
The land biospheric reservoir (soil and living biomass) constitutes a net sink of atmospheric CO2 at global scale. The size of this sink has increased alongside with increasing emissions, from 1.8 PgC in the 1980s to 2.6 PgC in the 1990s and 3 PgC over 2000–2008 (Le Quéré et al., 2009). It is currently considered that land biosphere carbon sink absorbs about 30% of the anthropogenic emissions from fossil fuel combustion and deforestation every year (Canadell et al., 2007; IPCC, 2007). This sink is explained by the increased net primary productivity (NPP) driven by a variety of environmental and management factors (Nemani et al., 2003; Piao et al., 2005; Magnani et al., 2007; Ciais et al., 2008; Lewis et al., 2009). Among the environmental factors, rising CO2, warming in ‘temperature-limited’ ecosystems, increased rainfall in dry ecosystems, extra-nitrogen deposition in ‘nitrogen-limited’ ecosystems and the combination of these factors are the main causes of increased NPP (Piao et al., 2006). This increase in NPP induces higher soil C stock by increasing C input to soil. However, although soil contains three times more C than the vegetation, it is a two times smaller sink than living biomass (Reichstein, 2008). There are two reasons for that. First, in managed ecosystems, biomass is partly harvested and consequently the soil does not ‘benefit’ from all the increased primary productivity. Second, in some ecosystems, increased primary productivity adds C into biomass pools where it could be stored for decades, and the soil has therefore not ‘seen’ the increase in primary productivity yet. This is especially true for the woody biomass pools of forests. Yet, global increase of soil C stock is mostly due to increased foliage and fine root productivity, which are neither harvested nor long lasting biomass pools (Nemani et al., 2003; Piao et al., 2006).
However, the ability of soil to behave as a C sink in the future is highly uncertain because SOC decomposition is sensitive to climate and increase with temperature. Generally, soil C respiration rate is multiplied by roughly a factor of two for a 10°C warming (Raich and Schlesinger, 1992; Davidson and Janssens, 2006; Reichstein et al., 2005) if microbial activity is not limited by substrate availability or soil moisture. The potential loss of soil C under the future warming forecasted by climate models is six times larger than the current soil C sink (Reichstein, 2008). Large uncertainties are associated with this estimate, both in the amount and spatial distribution of warming which depends on economic development scenarios and modelled climate processes, and in the response of soil C to climate change.
In the future, the net balance between extra inputs and extra losses is difficult to assess. Indeed, neither SOC input nor output fluxes are linear function of climate, and these functions are uncertain. It is generally assumed that decomposition increases exponentially with temperature (Lloyd and Taylor, 1994) and productivity is a saturating function of climate. Conceptually, this means that, above a certain degree of warming, the increase of respiration will not be offset by an equivalent productivity increase, and soils will lose C. Moreover, rising atmospheric CO2 concentration and warming may no more induce productivity increase in ecosystems where nutrient availability is limited and/or where drought accompanies the warming (Angert et al., 2005). This latter response is observed in simulations of the climate–carbon coupled system. For instance, in the widely cited study of Cox et al. (2000), Amazon forest productivity drops, causing a large and uncontrolled loss of soil C in the future century, acting as a positive feedback to global warming. Contrary to the results of this study, most global models of the coupled climate–SOC system predict a net gain of SOC during the 21st century even in the tropics, except for two models including the one used by Cox et al. (see Friedlingstein et al., 2006).
Moreover, in face of transient climate change, an ecosystem may lose soil C rapidly in response to warming while its productivity increases slowly over time. In that case, although the asymptotic value of soil C is theoretically proportional to the new productivity and may thus eventually increase above initial levels, a transient phase of soil C loss lasting several decades, followed by recovery and soil C build-up can be produced.
For soil layers where the C input are virtually zero, such as subsurface peat and permafrost C, the process of decomposition dominates the vulnerability of soil C (Zimov et al., 2006; Khvorostyanov et al., 2008a,b). Permafrost-affected soils span a wide range of soil environments, and thus the huge ‘permafrost soil carbon pool’ includes carbon stocks which have built up over a variety of timescales and with differing levels of vulnerability to global warming (Fig. 1). At their surface, permafrost-affected soils have a seasonally thawing active layer, in which SOC balance is, as in temperate soils, sensitive to both litter inputs and respiration fluxes over short timescales. However, the defining feature of permafrost soils is that below the surface there exist soil layers that do not thaw, and in these layers carbon may accumulate for long periods of time, leading to the massive quantities of carbon stored there. Shallow permafrost layers often contain carbon that is transported from the active layer through cryoturbation, and including this process in terrestrial models leads to a large increase in modelled C storage at high latitudes (Koven et al., 2009). The uncertainties of both the magnitude and vulnerability of these carbon pools increase with depth. Shallow permafrost carbon is clearly highly vulnerable, and field evidence suggests that carbon loss due to thawing of shallow permafrost is already taking place in Alaska (e.g. Schuur et al., 2008). The uncertainty on the quantity of carbon in deeper permafrost stocks, such as Yedoma, is much higher, as is the uncertainty on its vulnerability. Indeed, the thawing of deep permafrost within the next 100–200 years would require additional processes such as expansion of thermokarst lakes formed by meltwater (Walter et al., 2007) or release of heat by soil microorganisms (Khvorostyanov et al., 2008a,b), which are only beginning to be added to models. It is nonetheless interesting to note that most global carbon cycle models (Friedlingstein et al., 2006) show a net gain in high-latitude terrestrial carbon due to warming, while initial results using a permafrost-enabled version of ORCHIDEE (Koven C., Ringeval B., Ciais P., Friedlingstein P., personal communication) show carbon losses due to warming.
Figure 1. Schematic of permafrost carbon cycle. Shallow permafrost carbon is likely to be the most vulnerable to deepening of active layers with warming, whereas deeper permafrost carbon may be vulnerable as well due to processes such as thermokarst or microbial heating. Roots are not figured but may intervene in both active layer and shallow permafrost.
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3.2. Global vulnerability of SOC to land-use changes
Land cover and land management also strongly impact soil C balance acting on both decomposition and input. Land use changes that began with human agriculture several millennia ago, have escalated during the last centuries. These have consisted mostly of expanding cropland and pastureland and, more recently of urbanization. Today land use change is concentrated in the tropics, where forest clearing is taking place (Houghton, 2003). Tropical forest clearing induced a direct loss of C in biomass of 1.4 PgC yr−1 over 2000–2008, but clearing also causes an indirect loss of soil C (Houghton, 2003) that may last for decades after land use change (Pongratz et al., 2009). The fate of soil C in newly deforested areas converted to cropland or pastureland will depend on agricultural practice. Intensive cropping will quickly empty the former soil C pools within two to three decades (Arrouays et al., 1995; Jolivet et al., 1997; Reeves et al., 1997), whereas sustainable pasture management may stabilize them to a value close to the former tropical forest (Trumbore et al., 1995). In the following, we thus examine the effects of land use change on soil C pools during the last century, and into the future.
To appreciate the timescales involved in the evolution of soil C balance after a land use change, Fig. 2 shows the measured evolution of soil C in so-called ‘bare fallow’ field experiments. These soils, formerly under cropland or grassland were kept free of vegetation by human intervention during several decades, while SOC decomposition continues. Figure 2 shows that even after 40 years of bare soil conditions, the soil is continuing to lose C, apparently without reaching a new equilibrium. The shape of SOC decay in bare fallow plots cannot be reproduced with a simple 1-pool exponential decay model. Moreover, the SOC decay differs according to the former land use (Fig. 2). For example, the Rothamsted bare fallow implemented on grassland shows an initial large C loss that is not observed in the Askov bare fallow implemented on arable land. This suggests that soil organic matter is composed of two or more pools, with some pools having a decay time of 50–100 years, or more. Moreover, the size of the more stable pool may depend on previous land-use (Balesdent et al., 1988).
Figure 2. Evolution of soil C content in bare fallow experiments in the first 23 cm of a silty-loam soil at Rothamsted (England) and in the first 20 cm of a loamy-sand soil at Askov (Denmark). Data from Askov (Askov-FL1-B4) were taken from Petersen et al. (2005). Data from Rothamsted (High field bare fallow) were redrawn from Johnston et al. (2009).
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Consequently, any change in land use or ecosystem disturbance is expected to have a long lasting effect on the soil C balance. Former land use will continue to exert an influence on today's soil C budget and pre-condition its future evolution as well. Because of the cascade of long timescales involved, one needs to take a long historical perspective to simulate land use and climate change impacts on soil C balance.
3.3. Historical land cover, agricultural management and climate change effects on SOC
To gain insights on past land-use change effects on SOC regionally, the results of a global biosphere model forced by land-use and climate changes since 1900 are analysed. This model is a process-based global ecosystem model called ORCHIDEE (ORganizing Carbon and Hydrology In Dynamic EcosystEms model; Krinner et al., 2005). ORCHIDEE was used to quantify the effects of changes in climate, atmospheric CO2 concentration and changes in land cover area on terrestrial C cycle for the period from 1901 to 2002. SOC stocks were considered to be at steady-state equilibrium in 1901. This is a very crude assumption because human-induced land-use change has occurred before the 20th century particularly in Europe, in the Former Soviet Union and in East Asia (Pongratz et al., 2008). This human-induced land-use change is not negligible because it was estimated that about 5 millions of km2 were transformed to cropland or pasture between 800 and 1700 AD, leading SOC stocks to a significant disequilibrium in 1900 (Pongratz et al., 2009). For example, they estimated that, whereas vegetation lost about 100 Gt C from 800 to 1900, global SOC stocks increased by about 30 PgC due to additional plant material added to the soil pools from the converted natural vegetation. The steady-state assumption underlying the ORCHIDEE simulation results analysed below implies that the effect of land-cover change on SOC stocks is likely to be overestimated. More details about the model set up and architecture are given in Appendix.
3.4. Historical global land cover change and SOC balance
Both land cover and CO2+ climate changes have a strong impact in SOC balance. Running simulations taking into account CO2+ climate changes only or CO2+ climate + land cover changes allow determining the relative importance of land cover change. The results of these two simulations at global scale are presented on Fig. 3a. If only CO2 and climate drivers are accounted for, NPP and thereby biomass increases (Fig. 3a). Soil C shows a slight decreasing trend until 1960, due to warming induced acceleration of SOC decomposition in Europe and in temperate North America. Elsewhere, the CO2 fertilization effect on NPP incorporated in the model increases soil C input in excess to warming increased decomposition. Therefore, the soil C balance is slightly positive over the 20th century.
Figure 3. Evolution (a) of global soil and vegetation C stocks simulated by ORCHIDEE–LUC in response to combined effects of atmospheric CO2 concentration and climate, and to combined effects of atmospheric CO2 concentration, climate and land cover change and (b) of global soil C stocks simulated by the eleven models of the C4MIP intercomparison (Friedlingstein et al., 2006) from 1901 to 2002.
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By contrast, when annual land cover changes are incorporated in the study, the soil C stock decreases globally between 1900 and 1960. This land use induced decrease is later reversed, ending up only in a small variation of soil C, ΔSOC =−7.3 PgC by year 2000. The net change represents an insignificant 0.3% decrease of the initial stocks in 1901, assumed in equilibrium. The essential point here is that land-use change has an effect on soil C change between 1901 and 2000 that is equivalent in magnitude to the effect of climate and CO2 change, but of opposite sign. These results are in agreement with the results of a modelling study by McGuire et al. (2001) who did not distinguish soil carbon pools from the total land biospheric reservoir. Hence, McGuire et al. (2001) estimate with four terrestrial biosphere models that terrestrial ecosystems transitioned during the 20th century from releasing carbon to the atmosphere (∼8.8 PgC from 1920 to 1957) because of land-use changes to storing carbon (∼14.3 PgC from 1958 to 1992) mainly because of the increase in atmospheric CO2 concentration.
Because the change in NPP consecutive to forest clearing is not as large as the direct loss of biomass (ΔVEG), the impact of historical land cover change is more drastic on the biomass pools change than on the SOC pools change. Therefore, in our simulation, compared to the effect of CO2 and climate alone, including land cover change modifies ΔVEG globally from + 140 to −51 PgC and ΔSOC only from +27 to −7.3 PgC. This land cover effect on global SOC stocks (−34.3 PgC) is almost half the range of global SOC changes simulated by the eleven C4MIP models for the 20th century and represents ∼1/3 (MPI) to ∼100 times (UMD) the estimated effect of climate and atmospheric CO2 changes (Figs 3a and b). Moreover, ORCHIDEE–LUC shows the only negative SOC balance. This result highlights (1) the importance of land-use change comparatively to climate change and (2) the need for consideration of past land-use change in models that were parameterized to be consistent with the current contemporary global land carbon budget. On the other hand, the model does not reproduce secondary forest ecosystem growth unless a given grid point is subject to deforestation and later to reforestation. Thus, the negative impact of deforestation on the SOC balance is probably overestimated (Shevliakova et al., 2009).
When looking at the combined effects of land use, climate and CO2 changes on the SOC balance at regional scale, we can see in Fig. 4a different behaviour between tropical regions on the one hand, and temperate and boreal regions on the other hand. In temperate and boreal regions, the effect of land cover change on the SOC stocks is larger than in the tropics, even though the loss of biomass is greater in the tropics.
Figure 4. Evolution of regional soil and vegetation C stocks simulated by ORCHIDEE–LUC in response to combined effects of atmospheric CO2 concentration and climate, and to combined effects of atmospheric CO2 concentration, climate and land cover change from 1901 to 2002.
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More specifically, in the Former Soviet Union (FSU) and North America, forest clearing took place during the first half of the 20th century. This forest clearing has impacted the SOC balance with inertia, owing to the long residence times of SOC (Figs 4c and d). In the FSU, the modelled difference between SOC stocks with climate and CO2 changes only (S1) and SOC stocks with land-use changes (S2) is still increasing with time. In North America, this difference is rather constant after 1970, indicating that the indirect effect of land cover change on SOC stocks is stabilised. In Europe and China (Figs 4e and f), no major change in forest area took place during the 20th century (most of deforestation being earlier to that date). The SOC difference between S2 and S1 is thus negligible, indicating that land cover change did not impact SOC stocks. However, we will see below that SOC stocks in China and Europe has likely declined during the 20th Century due to intensification of agriculture. In China, large plantations created in the 1980s may contribute to an increase in SOC (Piao et al., 2009a,b), which is not well captured by the coarse scale global land cover maps used to force our model.
In our model simulation, climate and CO2 have increased SOC stocks everywhere but in Africa (because of drought that decreased NPP, see Fig. 4h) and in Canada (because increased NPP did not compensate warming increased heterotrophic respiration, see Fig. 4g). In contrast, land cover change has diminished SOC stocks everywhere. On average, the effect of land cover change on SOC balance between 1901 and 2000 is opposite to the one of climate and CO2. Regions where land cover induced SOC losses dominate over climate and CO2-induced SOC accumulation are the USA, the FSU and Europe. These regions experienced a net loss of SOC since 1901. Oppositely, a net SOC gain since 1901 was modelled for tropical Asia, China and South America. Despite deforestation in tropical Asia and South America, the increase in NPP induced by higher CO2 and rainfall in undisturbed forests (Lewis et al., 2009) has driven the pan-tropical SOC balance positive up to the present.
3.5. Agricultural intensity and SOC balance
On top of climate, CO2 and land cover, land management intensity also impacts SOC stocks, especially in cultivated soils. Several studies reported that changes in the fate of cropping residues (straw and stubble), cultivation and tillage practices have a strong impact on SOC stocks (Paustian et al., 1998, 2000; West and Post, 2002). In addition, arable soil erosion (which depends on SOC content, cultivation and tillage practices) causes a significant loss of SOC in cultivated soils. However, at global-scale, accumulation of eroded soil particles in floodplains may compensate for this loss at the field scale (van Oost et al., 2007).
In North America, this trend may have been reversed recently, as it has been considered that changes in agricultural practices have increased SOC sequestration during the late 30 years. Indeed, Ogle et al. (2003) estimated that adoption of reduced use of bare fallow, conservation tillage practices on cropland, and conversion of annual cropland to grass and trees in the Conservation Reserve Program (CRP) have resulted in a net gain of 10.8 Tg C yr−1 during the period from 1982 to 1997. Most of this gain was due to setting-aside lands in the CRP and only 1.3 Tg C yr−1 was attributed to change in US agricultural land-use and management. Using the ISAM-2 model, Jain and Yang (2005) estimated a 868 Tg SOC increase (i.e. 43.4 Tg C yr−1 from 1980 to 2000) in North American soils due to no-till practices. Productivity increases induced by irrigation may also have played a significant role in SOC stock increase in the USA. The SOC stock increase induced by irrigation was estimated at 21.3 Tg C in the Great Plains since 1950s using the CENTURY model (Parton et al., 2005).
In China, a continuous decline in SOC levels for a wide scope of Chinese agricultural soils has been observed since the 1950s (Qiu et al., 2009). This was mainly explained by the reduction of crop residue incorporation and manure amendments, which have been the major source of SOC for most Chinese farmlands (Qiu et al., 2009). As in North America, this trend seems to have been reversed recently. New policies have been launched during the past decades to encourage farmers to return more organic matter to the soil and some agricultural fields have shown SOC sequestration (Huang and Sun, 2006; Yan et al., 2007). In addition, intensive agriculture practices, such as application of chemical fertilizer, increased irrigation in arid areas, expansion of straw incorporation, and shallow plowing have also led to an increase in the C sequestration of agricultural ecosystems (Huang and Sun, 2006). Preliminary estimates suggested that along with the increase of agricultural productivity and less removal of biogenic material for societal energy purposes, about 0.025–0.037 PgC yr−1 is potentially fixed in soil (Lal, 2004), which is slightly higher than current estimation of Huang and Sun (2006) and Xie et al. (2007). Huang and Sun (2006) analysed changes of organic carbon stocks in China's agricultural soils during the past two decades, and conclude that the mean carbon sink of agricultural soils was 0.015–0.020 PgC yr−1. Overall, Piao et al. (2009a,b) estimated in a meta-analysis that topsoil in croplands is a net sink of 0.026 ± 0.011 PgC yr−1 since the early 1980s.
In Europe, SOC change was estimated using three models including specific crop management and practice ORCHIDEE–STICS, RothC, and LPJml. In both ORCHIDEE–STICS and RothC, the simulated soil C dynamics included the effect of past changes in agricultural technology on input and tillage, as well as the effects of climate and atmospheric CO2 changes. The ORCHIDEE–STICS model was initialized with (reconstructed) ancestral farming practice and crop varieties in 1900 (Gervois et al., 2008). This model provides a net change in SOC, ΔSOC = 0.01 ± 0.06 tC ha−1 yr−1 between 1901 and 2000. This insignificant gain is mostly explained by agricultural intensification and agricultural technology changes, with a small negative contribution of recent climate change over Southern Europe (Gervois et al., 2008). Indeed, in the Iberian Peninsula and a few southern Mediterranean regions, the increase in heterotrophic respiration and a precipitation-induced reduction in soil C inputs slightly overcome positive SOC stock increment induced by agricultural technology changes. At the European scale, a strong loss of SOC is modelled when agriculture became mechanized in the 1950s. But this loss trend was reversed in the 1970s and SOC increased thereafter. Arable lands are today C neutral (+0.16 tC ha−1 yr−1 with a range of 0–0.3 tC ha−1 yr−1). This modelled small sink is not supported by regional inventories, available in UK, France, Belgium and Finland, and in some regions of Germany, which rather suggest net SOC losses (Mäkelä-Kurtto and Sippola, 2002; Sleutel et al., 2003; Bellamy et al., 2005; Smith et al., 2005a,b). However, if the net C sink provided by erosion (Lal, 2003; van Oost et al., 2007) is taken into account, SOC losses measured in the inventory become comparable to the model results. This confirms small net changes in the SOC balance of arable lands in Europe.
Over 1990–1999, RothC and LPJml provide Net Biome Productivities (NBP) of −7.6 and 1.3 g C m−2 yr−1, respectively. It is noteworthy that these results are extremely dependent on assumptions about the management options employed and how these have changed over recent decades. In summary, the three process models integrated over Europe and the inventories available at regional-scale agree on the relatively small absolute magnitude of the net SOC balance of arable lands in Europe. However, each model's sensitivity tests point to the strong sensitivity of NBP to the assumed choice and past history of management options.
3.6. Forest management and SOC balance
The impact of forest management change on soil carbon has not been assessed for the past century. ORCHIDEE–FM, a novel version of ORCHIDEE includes a forest management module. Recent results emphasized the importance of forest management on SOC stocks (Bellassen et al., 2010a,b). The evolution of carbon stocks in a managed forest, as simulated by ORCHIDEE–FM, is reported on Fig. 5a. Harvesting wood increases soil carbon in the short term as 50–100 tC ha−1 of slash wood is added to the litter pool as coarse woody debris. This input leads to a peak in soil carbon respiration of about +50% after harvest, which is consistent with measurements (Savage and Davidson, 2001; Pypker and Fredeen, 2002; Pregitzer and Euskirchen, 2004). During decomposition, a part of the coarse woody debris is buried in deep soil horizons, driving a smaller and slower increase in soil carbon per se (Nabuurs et al., 2008). However, in the long term the impact of forest management on SOC stock is negative: as the stems of dead trees are no longer left to feed soil carbon, the long-term litter input is lower under management than in an unmanaged forest. On average, ORCHIDEE–FM simulates a soil carbon stock reduction of about 15 tC ha−1 under management (Fig. 5a) compared to no management (Fig. 5b).
Figure 5. Simulated carbon stocks during a rotation period in a temperate broadleaf forest stand (Bellassen et al., 2010b) (a) with management and (b) without management. AB wood is yearly average in aboveground biomass. CWD is coarse woody debris. Other litter stands for dead leaves, woody roots and fine roots. Soil carbon does not include litter.
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