Stable and radioactive carbon in forest soils of Chhattisgarh, Central India: Implications for tropical soil carbon dynamics and stable carbon isotope evolution

Journal of Asian Earth Sciences 123 (2016) 47–57

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Stable and radioactive carbon in forest soils of Chhattisgarh, Central India: Implications for tropical soil carbon dynamics and stable carbon isotope evolution A.H. Laskar ⇑, M.G. Yadava, R. Ramesh Geosciences Division, Physical Research Laboratory, Ahmedabad 380009, India

a r t i c l e

i n f o

Article history: Received 18 March 2015 Received in revised form 25 March 2016 Accepted 26 March 2016 Available online 31 March 2016 Keywords: Radiocarbon Soil organic matter Stable isotope Mean residence time Rayleigh fractionation

a b s t r a c t Soils from two sites viz. Kotumsar and Tirathgarh, located
5 km apart in a tropical reserve forest (18°520 N, 81°560 E) in central India, have been explored for soil organic carbon (SOC) content, its mean residence time (MRT) and the evolution of stable carbon isotopic composition (d13C). SOC stocks in the upper 30 cm of soil layers are
5.3 kg/m2 and
3.0 kg/m2; in the upper 110 m are
10.7 kg/m2 and
7.8 kg/m2 at Kotumsar and Tirathgarh, respectively. SOC decreases with increasing depth. Bomb carbon signature is observed in the upper
10 cm. Organic matters in the top soil layers (0–10 cm) have MRTs of the order of a century which increases gradually with depths, reaching 3500–5000 yrs at
100 cm. d13C values of SOC increase with depth, the carbon isotopic fractionation is obtained to be 1.2‰ and 3‰ for soils at Kotumsar and Tirathgarh, respectively, confirmed using Rayleigh isotopic fractionation model. The evolution of d13C in soils was also studied using a modified Rayleigh fractionation model incorporating a continuous input into the reservoir: the depth profiles of d13C for SOC show that the input organic matter from surface into the deeper soil layers is either insignificant or highly labile and decomposes quite fast in the top layers, thus making little contribution to the residual biomasses of the deeper layers. This is an attempt to understand the distillation processes that take place in SOC, assess the extent of decomposition by microbes and effect of percolation of fresh organic matter into dipper soil layers which are important for stable isotope based paleoclimate and paleovegetation reconstruction and understanding the dynamics of organic carbon in soils. Ó 2016 Elsevier Ltd. All rights reserved.

1. Introduction Soil is the largest terrestrial organic carbon reservoir with an estimated pool size of
1325 Gt (1 Gt = 1015 g) in the top 1 m and
3000 Gt in total including permafrost and tropical peat lands (Köchy et al., 2015), which is much more than that present in the atmosphere and living bio-mass combined (Schlesinger, 1997). This revised estimate of the soil organic carbon (SOC) is slightly different from that calculated previously (Eswaran et al., 1993; Batjes, 1996; Jobbagy and Jackson, 2000; Lal, 2004; Stockmann et al., 2013). A small increase in the rate of oxidation of soil organic carbon as a result of increasing temperature could result in a dramatic increase in atmospheric CO2 (Kirschbaum, 2000; Davidson and Janssens, 2006). Therefore, it plays an important role in the global climate dynamics and overall carbon cycle. The

⇑ Corresponding author at: Research Center for Environmental Changes, Academia Sinica, Taipei 115, Taiwan. E-mail address: [email protected] (A.H. Laskar). http://dx.doi.org/10.1016/j.jseaes.2016.03.021 1367-9120/Ó 2016 Elsevier Ltd. All rights reserved.

current increasing trend in the atmospheric CO2 may lead to an imbalance between uptake and loss of carbon in terrestrial ecosystems burdening the atmosphere further with more CO2. Tropical soils hold
25% of the global soil carbon stocks and are believed to have lesser residence time than temperate and boreal soils, though the rate of cycling of SOC in different soil pools is yet to be quantified adequately (Schimel et al., 1994; Thompson et al., 1996; Jobbagy and Jackson, 2000; Torn et al., 2009). While detailed quantitative studies of isotope based carbon dynamics in tropical soils from India are limited (Becker-Heidmann and Scharpenseel, 1989; Caner et al., 2007; Laskar et al., 2012), only a few studies have been made so far to assess the organic carbon stocks and the possible effect of climate change on tropical soils (Jenny and Raychaudhuri, 1960; Gupta and Rao, 1994; Bhattacharyya et al., 2000, 2009). To more fully understand the tropical soil carbon dynamics, it is therefore very important to extend the isotope based studies to the tropical soils with different climate and geography. Atmospheric carbon dioxide is converted to organic carbon through photosynthesis where it can be buried in the soil as

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A.H. Laskar et al. / Journal of Asian Earth Sciences 123 (2016) 47–57

litter, roots and plant and animal residues, altered into soil organic matter by microbial decomposition and part of which is respired back to the atmosphere. Previous studies have demonstrated that the stable carbon isotopic composition (d13C) of SOC increases with depth and is accompanied by a decrease in SOC (Garten et al., 2000; Poage and Feng, 2004). A number of possible mechanisms have been invoked to explain this observation. These include progressive 13C decrease in the atmospheric CO2 due to the extensive use of fossil fuels (the Suess effect) during the industrial era (Ehleringer et al., 2000), preferential decomposition of specific organic compounds (Deines, 1980; Feng, 2002), carbon isotope fractionation associated with microbial decomposition (Ehleringer et al., 2000; Accoe et al., 2002) and the change in the extent of fractionation during photosynthesis due to elevated atmospheric CO2 concentration (Treydte et al., 2009). Therefore, quantification of fractionation during decomposition is important as stable isotopes of SOC have been widely used for paleoclimate and paleovegetation studies (Bowman et al., 2004; Laskar et al., 2010, 2013). If microbial decomposition significantly modifies the d13C value of SOC through time, then the use of SOC as a proxy for paleoclimatic reconstruction must account for these potential changes. The study of isotopic fractionation of carbon during microbial decomposition in soils can also be very useful to shed light on the processes that govern the organic carbon storage and nutrient availability in sub-soil layers. The most widely used method to study the evolution of isotopic composition in a reservoir, due to partial removal of materials from it which associated with an isotopic fractionation, is the Rayleigh distillation model. However, its applicability is limited to closed systems, i.e., there is no input into the reservoir. Several authors applied the Rayleigh model directly to study SOC fractionation (Accoe et al., 2002; Wynn et al., 2005, 2006) assuming (i) the soil to be a closed reservoir and (ii) a very small fraction of the SOC is successively and continuously removed as CO2 in discrete steps via decomposition (with 12C getting removed preferentially). Wynn et al. (2006) found the model to be valid for the finer fraction (<63 lm). Accoe et al. (2002) applied the Rayleigh model to the bulk SOC under permanent grassland and fitted the model for the top
30 cm of the soil profile. But the validity of this model in open systems such as soils is debatable because of the continuous input of fresh organic matter from the top surface and their vertical percolation into the deeper soil horizons. Here, we attempt to modify the Rayleigh fraction model by including an input term and apply this model to two soil profiles located in a reserve forest in the Chhattisgarh state, Central India, unexplored thus far. We also present the values of concentration, d13C and mean residence time (MRT) of SOC in the two soil profiles. The specific objectives of the present study are: (i) to estimate the organic carbon content and factors that control the soil carbon inventory, (ii) to estimate the MRT of organic carbon in a typical tropical forest soil, (iii) to develop a model that successfully monitors the evolution of carbon isotopic composition (d13C) in SOC with depth. To address the vertical variation of carbon isotopic composition in a soil system, a modified Rayleigh model (Ramesh and Singh, 2010) has been applied to the present experimental data in which the following two assumptions are made: (i) there is a continuous addition of the fresh organic matter supplied from the surface which moves down the profile without any isotopic fractionation during transport and (ii) the isotopic fractionation occurs only during microbial decomposition processes, causing a difference between d13C of the respired CO2 and residual biomass.

2. Materials and methods 2.1. Study area, sampling and analysis The soil samples were collected in 2006 from Kotumsar and Tirathgarh (18°520 N, 81°560 E, 553 m above m.s.l.), two sites inside the reserve forests of Kangar Valley National Park of Jagdalpur district, Chhatisgarh, Central India (Fig. 1). The two sites are located within
5 km distance of each other, covered with dense tropical evergreen forests and human interference was almost absent. The samples were collected by making a vertical trench at each site. The long-term mean values of annual precipitation, temperature and relative humidity are 1533 mm,
25.7 °C and
63%, respectively (meteorological data from 1951 to 1980 CE: India Meteorological Department, 1999). The present day surface vegetation (Fig. 2a) is mainly medium to big size trees (C3 type) with average d13C values for the surface litter being 29.1 ± 1.0‰ and 28.2 ± 1.5‰ at Kotumsar and Tirathgarh, respectively. The soils are brown earth with thick litter layers. Structurally, the soils are clay loams with moderate organic carbon content, identified as alfisols. The human disturbance inside the forest is almost absent and therefore, the soil erosion and weathering are expected to be minimal. Samples were collected from trenches of depths 110 cm at Kotumsar and 170 cm at Tirathgarh. Fig. 2b shows the trench made at Tirathgarh. The bedrock was observed at depths of 110 and 170 cm at Kotumsar and Tirathgarh, respectively. The soil samples were sealed in plastic bags and brought to laboratory for analysis. Before analysis, soils were sieved through a 2 mm mesh to remove plant detritus, roots and gravels. Small rootlets were removed physically. The soils were treated with 5% HCl to remove carbonate if present. Stable and radioactive carbon isotopes were measured up to depths of 110 and 106 cm at Kotumsar and Tirathgarh, respectively, while particle size analysis was carried out up to the deepest layers of the two profiles. The 14C contents of soil were estimated using liquid scintillation spectrometry. About 100 g of a decarbonated and rootlet-free soil sample was combusted at a temperature of
900 °C in presence of pure oxygen to produce CO2. The CO2 was purified cryogenically and expanded in a known volume with a pressure gauge connected to it to measure the concentration. The error associated with the volumetric concentration measurement was less than 2%. An aliquot of the pure CO2 was measured for d13C for fractionation correction. Benzene was synthesized from the CO2 following standard procedure, and the radioactivity of the benzene was measured using a liquid scintillation counter 1220-Quantulus at Physical Research Laboratory, Ahmedabad (Yadava and Ramesh, 1999). Radioactivity of benzene was reported as percent Modern Carbon: pMC = A/Ao  100 where, A and Ao are the specific activity of sample and Oxalic acid II, respectively (Stuiver and Polach, 1977). For d13C measurements,
1 g CuO and some silver foils were added to
1 g of powdered soil sample, mixed and poured in a quartz tube, evacuated, sealed and heated at
900 °C for about an hour. The CO2 produced was purified cryogenically and analyzed using a dual-inlet stable isotope ratio the mass spectrometer (Europa-Scientific GEO 20–20) at Physical Research Laboratory, Ahmedabad. The carbon isotopic ratio is expressed as d13C in parts per thousand (‰) relative to the Vienna Pee Dee Belemnite (VPDB) standard. The accuracy of measurement was better than ±0.1‰ for d13C, checked using two international standards Oxalic Acid II and ANU Sucrose with d13C values of 17.8‰ (Nehmi, 1980) and 10.45‰ (Coplen et al., 2006), respectively relative to VPDB (also discussed in Laskar, 2012). The reproducibility in d13C was 0.1‰ estimated by analyzing five ANU Sucrose samples during measurement of the samples. The total organic carbon (kg/m2) was estimated using the procedure described by Amundson (2001).

A.H. Laskar et al. / Journal of Asian Earth Sciences 123 (2016) 47–57

Fig. 1. Map of Chhattisgarh showing the study area, inset shows the map of India.

Fig. 2. (a) View of the dense forest at Kotumsar and (b) a vertical trench made at Tirathgarh for sample collection.

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A.H. Laskar et al. / Journal of Asian Earth Sciences 123 (2016) 47–57

Separation of clay (size: <2 lm), silt (size: 2–63 lm) and sand (size: >63 lm) were carried out using the sieve-pipette method (Carver, 1971).

The steady state solution of the Eq. (2) in terms of 14C content is given by

k¼ 2.2. MRT of soil organic carbon The MRT or turn over time of soil carbon is defined as the average time that elapsed between the incorporation of carbon in the vegetation via photosynthesis and its release back to the atmosphere by respiration. Residence time of organic carbon in leaf and roots, the main contributor of the SOC varies from a few months to a few years which is almost insignificant compared to the MRT of SOC (Zhang et al., 2010). The MRT of SOC is different from the true soil age as the latter is an open system with continuous input of fresh organic matter and loss due to soil respiration, in addition to radiocarbon decay. At the most, MRT can be considered to be the minimum age of the soil (Wang et al., 1996). For active SOC (bomb carbon present: pMC > 100%), the MRT is calculated using the relation described by Hsieh (1993) given by

Pp Aa ¼

i¼b

exp ½ðp  iÞ=MRT   14 Ci exp ½ðp  iÞ=8268 Pp i¼b exp ½ðp  iÞ=MRT 

ð1Þ

where Aa is the specific 14C activity of an active SOC pool, p is the year of sample collection, 14Ci is the specific 14C activity in the atmosphere in the ith year, MRT is the mean residence time of organic carbon in soil and 8268 years is the mean life of 14C. The equation is derived considering the soil to be in a steady state in terms of assimilation and release of organic carbon i.e., the input and decay rates of organic carbon in the soil are equal. The average atmospheric radiocarbon content for the Northern Hemisphere (from 1950 up to 2006 CE) was taken from Hua et al. (2013), and pre 1950 values were assumed to be 100% modern (i.e. constant) except during 1890 to 1950 CE, the value was taken to be 98%, considering Suess effect which dropped atmospheric 14C level by
2% during the industrial era due to extensive fossil fuel combustion (Tans et al., 1979). A slight change in the atmospheric 14C content prior to 1950 CE does not change the estimated MRT significantly. The above method loses its sensitivity for older soils i.e., for soils without active carbon pool or bomb carbon (Hsieh, 1993). A soil layer is assumed to have bomb carbon if pMC > 100%. Bomb carbon can be present even for pMC < 100%, but this method is not applicable in such cases. For soils with pMC < 100% carbon, MRT was calculated following Wang et al. (1996). In brief, the variation of organic carbon with time for a soil in any horizon is the result of net production of organic matter through litter, roots and bio-diffusion and loss due to microbial decomposition given by the relation

dC ¼ /  kC dt

ð2Þ

where / is the net production in moles/cm3/yr and kC is the rate of carbon loss with k the dissociation constant. As soil is a heterogeneous mixture of different components with different decay rates, the assumptions that the / and k are constants may not be valid always. However, for simplicity, they are assumed to be the average value of a long duration. Another assumption is that the 14C content of the input organic carbon is the same as the atmospheric CO2, implying that the time spent by carbon atoms from the fixing by photosynthesis to the conversion of SOC is neglected. This time is of the order of a few months to few years for leaf litters and roots, the main contributor of the SOC. These inputs go directly into the given pool, i.e., the input in any given soil horizon is assumed to have contemporary 14C values and the transfer of carbon from one pool to another is neglected. The equation is solved for steady state i.e., the rate at which organic matter added to a given soil layer is the same as that released.

ðd14 Com þ 1000Þk d14 C iom  d14 C om

ð3Þ

where d14Com and d14 Ciom are the d14C values of the organic matter present at a given depth and input organic matter (litter), respectively and k is the 14C decay constant. d14C is defined by h 14 12 i CÞsam d14 C ¼ ðð14C=  1  1000, measured using a liquid scintillation C=12 CÞ std

counter. Standard here is the Oxalic Acid II. The reciprocal of the decay constant (1/k) is considered as the MRT of SOC. 2.3. Depth evolution of the soil d13C through a modified Rayleigh model 2.3.1. Rayleigh model Rayleigh isotopic fractionation model (henceforth referred to as the classical Rayleigh Model to distinguish it from the modified one) is applicable for a reservoir from which material is successively removed along with an associated isotopic fractionation. In the case of carbon isotopic evolution in soils, the model is applicable if the change in d13C in soils is due to discrimination against 13C during decomposition and not by any other means. Let N be the total number of molecules in the system and R, the ratio of the rare to the abundant isotope. The Rayleigh equation describing the evolution of isotopic composition of the reservoir is given by (Gonfiantini, 1981):

 R ¼ Ro

N No

a1

ð4Þ

where Ro is the initial ratio (13C/12C), No is the initial number of molecules in the system and a is the fractionation factor defined by a ¼

RrCO

2

Rrbm

with RrCO2 and Rrbm , the

13

C/12C ratios of respired CO2

and the residual biomass, respectively. In terms of d13C, the above equation can be written as:

d13 C ¼ ðd13 Co þ 103 Þf

a1

13

 103 13

where d Co is the initial d C and f ¼

ð5Þ N No

is the fraction of the

initially deposited material remaining at a given depth. Eq. (5) can approximately be written as (Clark and Fritz, 1997):

d13 C ¼ d13 Co þ e ln f

ð6Þ 3

where e = (a  1)  10 is called the enrichment factor. The way a is defined, it is always <1 for SOC decomposition and hence e is always negative (as is ln f for f < 1). 2.3.2. Extending Rayleigh model to an open system with one source Classical Rayleigh model, discussed above is valid for closed reservoirs from which material is removed continuously with a preference to a particular isotope over the others. In an open system such as soil, material is also added continuously; therefore this requires incorporation of a source term in the classical Rayleigh model. In such a case, the isotopic ratio (R) at any given depth is given by the following mass balance equation (Ramesh and Singh, 2010):  a þ1 R ¼ Ro f ðb1 Þ þ

h i b  a þ1 Rin 1  f ðb1 Þ aþb1

ð7Þ

where Ro and Rin are the 13C/12C ratios of initial organic matter and the added input, respectively and b(= Nin/Nout) is the ratio of the amount of material added (Nin) to that lost (Nout). When there is no input (b = 0), the above equation reduces to a classical Rayleigh Eq. (4). The above equation has been derived considering b – 1 as b = 1 leads to singularity which is discussed in the next paragraph

A.H. Laskar et al. / Journal of Asian Earth Sciences 123 (2016) 47–57

(for details, see Ramesh and Singh, 2010). In terms of d notation, the Eq. (7) can be written as:

i din b h  a þ1  a þ1 d ¼ do f ðb1 Þ þ 1  f ðb1 Þ aþb1   a þ1 b b ðb1  a þ1 3 Þþ  f ðb1 Þ  1 þ 10 f aþb1 aþb1

ð8Þ

For b = 0, the Eq. (8) reduces to the classical Rayleigh Eq. (5). When b = 1, i.e., the rate at which material added to the system is equal to that lost, the so called steady state, the mass balance equation becomes (Ramesh and Singh, 2010).

R ¼ Ro e

aN  in No

þ

Rin

a



aN   in 1  e N o

ð9Þ

Which when expressed in terms of d notation becomes Nin

d ¼ do ea No þ

din

a

      Nin Nin Nin 103 1  ea No þ 1  ea No  103 1  ea No

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and natural soils (Natelhoffer and Fry, 1988; Feng et al., 1999; Santruckova et al., 2000). The input added is assumed to suffer no fractionation (i.e. ai = 1) and its isotopic composition (di) is the same as that of the surface litter. The initial isotopic composition (do) is taken as the d13C of the SOC at the soil surface. It is to be noted that the model does not include the effects of preferential decomposition of specific organic compounds. It is found that such preferential decomposition is limited only to the early phase of the litter decay (Feng, 2002). The data used here are not sufficient to consider the preferential decomposition of different organic compounds. This model can also be applied to different organic compounds by parceling the bulk organic matter into discrete pools each with an individual set of model parameters. Another shortcoming of the model is that it cannot quantify the differential inputs in different soil layers. 3. Results and discussion

a

ð10Þ

2.4. Model input parameters The model presented here describes the evolution of d13C of SOC in a soil profile, assuming that d13C of the input organic matter, mainly in the form of dissolved organic carbon (DOC), supplied from the soil surface remains same, and with time the subsequent enrichment in the d13C of SOC along depth is purely due to isotopic fractionation during decomposition. The assumption that there is no change of the isotopic composition of the downward percolating organic matter is not always valid. However, we demonstrated (later in the present study) that the contribution of added organic matter, especially in the deeper soil layers is very little. Loss of carbon can also take place via leaching in the form of dissolved inorganic carbon (DIC), DOC and dissolved methane. SOC mainly leaches in the form of DOC, which is very less compared to the DIC especially, in the deeper soil horizons (Kindler et al., 2011) and hence it is neglected in the mass balance calculation. d13C enrichment of the residual biomass is calculated by fitting the model line with the experimental data points and is expressed by a carbon isotope enrichment factor e (e = (a  1)  1000, where a is the fractionation factor and is not tightly constrained). The reported values of e vary between 0‰ and 3‰ for laboratory

3.1. Soil organic carbon content and particle size distribution The distribution of clay, silt and sand at various depths are shown in Fig. 3. The clay content increases with depth between 0 and 50 cm at Kotumsar and 0 and 70 cm at Tirathgarh, and after attaining maximum values, it decreases further down in the profiles. The clay content is much less at Kotumsar compared to Tirathgarh while the reverse is the case for silt content; the sum of the clay and silt contents is almost the same in both the profiles. Sand content is considerably less than the cumulative silt and clay content in both the sites. The variation in organic carbon content is shown in the Fig. 4a. Between surface and
20 cm depth the organic carbon content decreases quite rapidly especially, at Kotumsar. The stock of total organic carbon estimated for soil layers between 0 and 30 cm, is
5.3 and
3.0 kg/m2 at Kotumsar and Tirathgarh, respectively. Similarly, for layers between 0 and 110 cm the value is
10.7 kg/m2 for Kotumsar and
7.8 kg/m2 for Tirathgarh (Table 2). These values are within the observed range (Batjes, 1996) of organic carbon stocks in Alfisols (Luvisols and Notosols) which are 2.0–10 kg m2 (up to 30 cm) and 5.1–18.2 kg m2 (up to 100 cm). As the two sites are close by (
5 km) and experience similar climate, the observed variation in the organic carbon content, mainly in the top 30 cm, is probably due to non-climatic factors. Higher organic carbon at Kotumsar soil might be due to the higher vegetation density, higher silt contents

Fig. 3. Depth profiles of the particle size distribution at (a) Kotumsar and (b) Tirathgarh.

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A.H. Laskar et al. / Journal of Asian Earth Sciences 123 (2016) 47–57

Fig. 4. (a) Organic carbon contents in weight%, (b) d13C of soil organic carbon, (c) percent Modern Carbon (pMC) and (d) mean residence time (MRT) of the soil organic carbon in the two profiles. Encircled points at the bottom of the Kotumsar profile are measured values for organic carbon in bedrocks.

and a different surface topography as these factors do control the soil formation and organic carbon content (Jenny, 1994; Jobbagy and Jackson, 2000; Silver et al., 2000; Tatsuhara and Kurashige, 2001; Tan et al., 2004; Rhoton et al., 2006; Trumbore, 2009). Carbon content in some tropical forest soils can be more in coarser fractions than in the fine clay probably due to more penetration of roots in the former (Silver et al., 2000). The slightly inclined surface is probably another factor for lower SOC content at Tirathgarh (Fig. 2B). The surface is flatter at Kotumsar resulting in higher retention of organic matter here. 3.2. d13C of SOC The d13C depth profiles are shown in Fig. 4b. The present day vegetation in the region is dominated by C3 plants: d13C values of surface vegetation are 29.1‰ and 28.2‰, respectively at Kotamsar and Tirathgarh (Table 1). The d13C values typically vary between 20‰ and 34‰ for the C3 type vegetation and 9‰ to 19‰ for the C4 type (Liu et al., 2003; Kohn, 2010). Depth profile of d13C shows 3–6‰ enrichment in the deeper layers, similar to observations reported from elsewhere (Natelhoffer and Fry, 1988; Becker-Heidmann and Scharpenseel, 1989; Accoe et al., 2002; Baisden et al., 2002; Wynn et al., 2006; Boström et al., 2007). This is attributed to isotopic fractionation during degradation by microorganisms. Other possibilities which could lead to the enrichment in 13C in the deeper soil layers are: (a) Suess effect: during the industrial era, because of the fossil fuel combustion, addition of 13C depleted CO2 has lowered the atmospheric d13C by
1.5‰ (McCarroll and Loader, 2004). This decrease in d13C of atmospheric CO2 could have a similar effect on SOC; younger soils in the shallow layers would have lighter preserved d13C values compared to older soils in deeper layers. (b) Change in the extent of fractionation by plants during photosynthesis due to elevated atmospheric CO2 concentration. Treydte et al. (2009) analyzed trees from higher elevations to check the impact of climate and CO2 on the d13C of recently laid rings in trees. After correcting for the Suess effect, they observed a net decrease of 0.012‰/ppmv in the d13C of the tree-rings. Therefore, increase in concentration of atmospheric CO2 by
100 ppmv during industrial era may lead to a decrease of 1.2‰ in the tree-ring d13C. But this observation may be site dependent as counter-argued in another study carried out in the northern boreal forests, where during atmospheric CO2 assimila-

tion; plants did not show any dependence on the prevailing atmospheric CO2 levels (Saurer et al., 2004). 3.3. MRT of SOC pMC and MRT of SOC at different depths are presented in Table 1. pMC is an indicator of the amount of residual radiocarbon present in a given soil horizon. Fig. 4c shows the variation of pMC with depth for the two soil profiles. Variation in MRT of organic Table 1 Stable carbon isotopic composition (d13C), organic carbon concentration, radiocarbon in percent Modern Carbon (pMC) and estimated mean residence time (MRT) of SOC in the two soil profiles. The results are based on single measurements at each depth and the errors are the ±1r analytical uncertainties. The errors in MRTs are the statistical errors associated with the 14C counting. Presence of bomb carbon is indicated by pMC > 100%. All the MRT values are rounded to the nearest ten. Depth (cm)

d13C (‰)

Organic carbon (wt.%)

pMC (%)

MRT (yrs)

Kotumsar PRL-3076 (litter) PRL-3067 PRL-3068 PRL-3069 PRL-3070 PRL-3071 PRL-3072 PRL-3077 PRL-3078 PRL-3073 PRL-3074 PRL-3079

– 0–2 2–4 6–8 8–10 10–12 18–20 28–30 38–40 46–48 58–62 100–110

29.06 27.67 27.62 26.05 26.06 25.18 25.21 25.49 25.89 25.78 25.34 25.11

– 3.19 2.72 1.79 1.56 1.41 0.77 0.79 0.63 0.73 0.58 0.52

102.91 ± 0.87 103.36 ± 0.92 109.37 ± 0.97 102.77 ± 0.91 97.33 ± 0.89 101.16 ± 0.90 92.21 ± 0.84 83.33 ± 0.90 79.54 ± 1.05 81.35 ± 0.86 77.70 ± 0.74 70.95 ± 0.80

–

Tirathgarh PRL-3056 (litter) PRL-3057 PRL-3058 PRL-3059 PRL-3060 PRL-3061 PRL-3062 PRL-3063 PRL-3064 PRL-3065 PRL-3066 PRL-3080 PRL-3081

– 0–2 4–6 8–10 16–18 28–30 38–40 50–52 60–62 70–74 78–82 90–94 102–106

28.20 26.59 25.45 24.96 24.67 24.73 24.90 24.60 24.32 23.08 22.02 21.47 21.25

– 1.14 1.05 0.90 0.70 0.64 0.43 0.46 0.38 0.45 0.39 0.33 0.31

104.52 ± 0.92 104.91 ± 0.94 102.80 ± 0.99 103.83 ± 0.93 92.01 ± 0.91 86.21 ± 0.96 83.59 ± 0.96 76.01 ± 1.06 75.12 ± 0.75 67.86 ± 0.67 64.57 ± 0.65 66.59 ± 0.87 65.88 ± 0.78

Sample code

180 ± 70 80 ± 70 200 ± 70 470 ± 80 270 ± 70 960 ± 80 1940 ± 90 2430 ± 110 2190 ± 90 2680 ± 80 3730 ± 100 – 140 ± 70 200 ± 80 170 ± 70 980 ± 80 1600 ± 90 1910 ± 100 2930 ± 120 3060 ± 80 4270 ± 80 4910 ± 80 4510 ± 110 4650 ± 100

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carbon with depth is shown in Fig. 4d. The surface soil horizons up to a depth of 10–15 cm have MRT of SOC of the order of a century showing fast cycling organic matter which increases with depth showing the dominance of slow cycling components in subsurface horizons. The upper soil layers up to a depth of
10 cm clearly show the presence of bomb carbon indicating that a significant fraction of organic carbon in these layers is accumulated after 1950s. In the lower horizons, MRT increases almost linearly with depth. There is a significant difference between the MRTs at the same depth in the two soil profiles in the deeper soil layers. MRT is almost 1000 yrs more for Tirathgarh at depths 50–100 cm (Fig. 4d). MRT depends on several factors e.g. (i) ambient precipitation and temperature, especially for the fast cycling components in the upper soil layers (Sorenson, 1981; Motavalli et al., 1994; Carvalhais et al., 2014; Trumbore, 1997), (ii) particle size distribution, especially the clay content (Oades, 1988; Torn et al., 1997; Jobbagy and Jackson, 2000; Telles et al., 2003; Laskar et al., 2012), (iii) the supply of fresh substrate from the surface that provides nutrients and energy to the decomposing organisms (Fontaine et al., 2007) and (iv) priming effect which is the extra decomposition of soil organic matter that occurs when microbes are stimulated by the addition of easily decomposable organic matter (Kuzyakov et al., 2000; Sayer et al., 2011; Foereid et al., 2014). In the present study, as the climate is common for the two sites, the first factor above is ruled out. A major factor for the higher MRT at a given depth of the soil profile at Tirathgarh compared to that at Kotumsar at the same depth is probably the higher clay content at the former site (Factor ii above). Clay particles provide reactive surfaces onto which organic matter can be adsorbed and such adsorption reactions stabilize the SOC against the microbial attack. Thus soils with higher clay content facilitate the formation of passive carbon pools with slow turnover times. Stabilization of SOC by clay particles are observed by several previous researchers (Schimel et al., 1985a, 1985b; Feller et al., 1991; Silver et al., 2000; Telles et al., 2003; Kleber et al., 2007; Torn et al., 1997, 2009). Oades (1989) discussed various interactions between SOC and clay particles. Also the supply of the fresh organic matter penetrating from the surface (factor iii above) is controlled by the soil texture. Higher amount of clay% may resist the downward transport of the modern organic material causing less priming effect (factor iv above) and hence higher residence time of SOC in the clay rich soil horizons. It seems that the clay content is one of the major factors that determine the MRT of SOC, but it needs more study especially the association and stabilization mechanisms of carbon with different clay minerals. 3.4. Evolution of d13C of SOC We consider that the SOC present at any depth is due to a balance between the fresh organic matter contributed by and percolated from the surface along with that present originally and the loss by microbial decomposition. Eq. (8) is fitted to the observed data, plotted between d13C as a function of the fraction of initial material left at different depths for different values of b varying from 0 to 1 and enrichment factor from 0 to 5‰ (Fig 5). The experimental data points fit best with the model plots for low

values of b < 0.2, checked using the least square method; this suggests that soils obey the classical Rayleigh fractionation model (b
0). For a better representation, the classical Rayleigh model described by Eq. (6) is fitted in Fig. 5b and d, and values of the parameters summarized in Table 2. Relatively less enrichment in 13 C with depth at Kotumsar, especially in the upper soil layers is probably due to the penetration of higher percentage of fresh organic matter from the surface. In the deeper soil layers, the contributions of the surface SOC is very small. A significant input of fresh organic matter from the top to the sub soil layers is expected to deviate the experimental data points from the Rayleigh fractionation curves. In the present climatic condition when the forest is dense and receives high precipitation, amount of surface carbon input cannot be neglected. It seems that the organic matter that percolates downward is mostly labile and decomposes quite fast. Further, due to the presence of high percentage of clay that is strongly bound to the old metabolized organic matter, the latter does not get rejuvenated with fresh organic matter input from the surface, having the d13C value of the contemporary vegetation. This is probably true for organic carbon contributed from the roots, especially the fine roots. Jobbagy and Jackson (2000), based on >2700 soil profiles, covering all the climate, geography and soil types concluded that there was no association between the relative contents of SOC and roots in the top 20 cm. Also clay binds organic matter strongly and favors the formation of aggregates, which protect against microbial decomposition (Balesdent et al., 2000; Giardina and Ryan, 2000; Rice, 2002; Bradford et al., 2008). We also checked the steady state evolution of the d13C in the two profiles. In a steady state, the rate at which material added is equal to that removed (b = 1, Eq. (10)). Fig. 6 shows the steady state d13C evolution for typical values of e = 1.2‰ and e = 3‰ (value of slopes in Fig. 5b and d) with d13C of the input organic matter of d13Cin = 28‰. Results show that the depth evolution of the d13C should follow simply a linear trend with only small changes in their values with depth. Such a behavior seems to be partly applicable to the top layers (0–4 cm) of the Kotumsar (Fig. 5a, Table 1) profile. Generally, the isotopic enrichment (e) depends on the rate of decomposition (Feng, 2002) and the environmental temperature (Garten et al., 2000); greater e values may be associated with higher rates of decomposition or higher soil temperatures. In the present scenario, the two soil profiles studied are part of a dense forest (average annual temperature
25.7 °C) with evergreen canopy and the same conditions could apply to the past, when human population was much less. Therefore, the role of temperature is likely to be limited here. The observed differences in e (1.2‰ and 3‰) is mainly due to different clay contents of the two soils which needs to be confirmed with more such studies. Accoe et al. (2002) estimated an enrichment factor of 2.3‰ in sandy-loam soil with 9% clay and 36% silt, under permanent grassland from Ghent University, Belgium where the mean annual temperature is
10 °C. An enrichment factor of 2‰ to 2.5‰ was observed in fine silty soils from the Yazoo River basin of Lower Mississippi (Wynn et al., 2006), with a mean annual temperature of
17 °C. Fig. 7 shows the evolution of d13C as a function of material left for the two widely different soil profiles mentioned above.

Table 2 Total organic carbon content and enrichment factors for d13C obtained using Rayleigh distillation model for the present two soil profiles and compared with two other widely different soil profiles from other locations. Site

Organic C content in top 110 cm (kg/m2)

Enrichment factor (e) ‰

Data source

Kotumsar Tirathgarh Ghent University, Belgium Lower slope, Yazoo Basin, Mississippi

10.7 7.8 – –

1.2 3.0 2.2 2.1

This study This study Accoe et al. (2002) Wynn et al. (2006)

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Fig. 5. Evolution of carbon isotopic composition as a function of the fraction of the material left (N/No) at different depths for different values of b (b – 1). b is the ratio of the amount of material added to that lost. The observed data fit best with the classical Rayleigh fractionation line (Eqs. (5) and (6) in the text). (a) and (b) are for the Kotumsar showing an observed enrichment factor of 1.2‰ while (c) and (d) are for the Tirathgarh, where enrichment factor is 3‰. Enrichment factors are the slopes of the Rayleigh fractionation lines (b) and (d).

Fig. 6. Steady state evolution of carbon isotopic composition (b = 1, Eq. (10)) for two values of the enrichment factors viz. e = 1.5‰ and 3‰, calculated assuming initial d13Cin = 28‰. Steady state refers to equal rate of addition and removal of material.

The data points fit well with the classical Rayleigh curve (b
0) as the authors predicted. The enrichment factors observed are 2.2‰ and 2.1‰ (Table 2), which are similar to that observed in the present case. Feng (2002) predicted using a theoretical model that the isotopic enrichment depends on the rate of decomposition. A higher rate of decomposition is associated with greater 13C enrichment as proposed by Garten et al. (2000). In the present case, we observe a significant variation in the enrichment factor for two soil profiles under similar climate. Therefore, our preliminary conclusion is that the decomposition and fractionation depends not only

on soil and climate parameters, but also on soil biology and clay minerals. More study is required to refine these predictions, especially the role of clay content and mineralogy and the roles of specific surface area on the distillation effect, through both laboratory experiment and field studies. The past variation of d13C in soils, sediments and paleosols is not only affected by the change of the surface vegetation (e.g., C3–C4 or vice versa) but also by the extent of microbial decomposition which needs to be considered before paleovegetation reconstruction. A possible way to estimate the change in the d13C due to microbial decomposition and transportation of new organic matter into the deeper soil horizons is to use the modified Rayleigh model as discussed here. In the present study the soil is considered as a single component whereas, in reality soil is a complex mixture of several components and hence there are scopes for refinement of the study applying the model into different soil components. A significant difference in the d13C values of SOC in different soil fractions has been observed (Marin-Spiotta et al., 2009), the extent of decomposition and discrimination by microbes in different SOC fractions could be different and hence each may have significantly different d13C depth profiles. This model could do with further refinement to account for these. The model can also be extended to study the temporal evolution of stable carbon isotopic composition in a soil profile. The time period during which the decomposition of SOC occurred can be obtained from the MRT estimated using radiocarbon as discussed here. An understanding of the distillation processes that take place in SOC will be of great benefit to the measurement and modeling efforts which describe the size and isotopic composition of global SOC pool. The distillation process may also extend our understanding of SOC storage on geological time scales. This model can also be extended to understand the role of pedogenic processes on the

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Fig. 7. Evolution of carbon isotopic composition as a function of the fraction of the material left (N/No) at different depths for different values of b (b – 1) for two widely different soils: (a) Ghent University, Belgium (0–30 cm) (Accoe et al., 2002) and (b) Lower slope, Yazoo Basin, Mississippi (0–20 cm) (Wynn et al., 2006). The data fit best with the classical Rayleigh fractionation line with enrichment factors of 2.2‰ and 2.1‰, respectively.

enrichments of other heavy isotopes such as 15N and 34S with depth and age of soil organic matter. Radiocarbon content in SOC shows that the average residence time of SOC is several hundreds to thousands of years. But the carbon turnover time for an ecosystem (calculated based on the carbon inventory divided by the flux of emitted CO2) is found to be
10 years in tropical grasslands and
500 years for Tundra and wetland environments with a global average of 32 yrs (Raich and Schlesinger, 1992) which recently re-estimated to be 15 years in the tropics and the global average is 23 years (Carvalhais et al., 2014). Turnover time, estimated using rate of change in the d13C due to plant cover change from C3 to C4 type or vice versa were found to be shortest when root derived d13C inputs are used (Marin-Spiotta et al., 2009). The large difference in the estimates between the present method and that estimated using flux of CO2 or by change in d13C due to change in vegetation type is due to the fact that respired CO2 flux is originated from the roots and microbial respirations and decomposition of modern labile organic matter whereas much of the carbon residing in deeper soils is stabilized and decomposes quite slowly (Trumbore, 1997). Radiocarbon based estimation of turnover of SOC, considering soil as a single reservoir may lead to a significant underestimation of the fluxes of organic matter through the soil organic matter pool, especially in the tropics (Trumbore, 1993). A multi-pool division of heterogeneous soil based on labile and refractory nature of organic matter and estimation of turnover of SOC in each pool is probably a more realistic approach as demonstrated with two components for tropical soils (Hall et al., 2015). Another alternative way to divide soil into labile and resistant pools is ramped pyrolysis (Rosenheim et al., 2008), a stepwise heating of organic matter to higher temperatures and estimation of radiocarbon content at each steps of heating. A systematic radiocarbon measurement of SOC in the topical soils using ramped pyrolysis may improve our understanding of the tropical soil carbon cycling especially its stability and sensitivity to warming.

4. Conclusions Soil organic carbon content is found to be moderate in the forest soils from Chhattisgarh, Central India. The mean residence time of the organic carbon varies from a century (in the top 10 cm) to
5 kyr (at 100 cm depth). Bomb carbon is found to have penetrated only up to the upper 10 cm of the soil. The residence time

in a given soil profile is mainly governed by the clay content of the soil, larger the clay content, higher is the residence time. A strong agreement of the soil organic carbon d13C data with the Rayleigh isotopic fractionation model indicates that the enrichment in 13C with depth is due to different extents of decomposition associated with a discrimination against 13C. A modified Rayleigh fractionation model predicts that the input of the fresh organic matter to the sub soil layers is very little relative to the size of the carbon pool, or the input organic matter decomposes rapidly in the surface soil layers without much affecting the isotopic composition of the SOC in deeper layers. The enrichment factors for d13C are 1.2‰ to 3‰ for Kotumsar and Tirathgarh, respectively, the difference is probably due to the interplay between the climate, biology and clay minerals. Acknowledgements We thank the Indian Space Research Organization Geosphere Biosphere Programme for funding. Helpful comments and suggestions from Yukio Isozaki and two anonymous reviewers are highly appreciated. References Accoe, F., Boeckx, P., Cleemput, O.V., Hofman, G., Zhang, Y., Li, R.H., Guanxiong, C., 2002. Evolution of d13C signature related to total carbon contents and carbon decomposition rate constants in a soil profile under grassland. Rapid Commun. Mass Spectrom. 16, 2184–2189. Amundson, R., 2001. The carbon budget in soils. Annu. Rev. Earth Planet. Sci. 29, 535–562. Baisden, W.T., Amundson, R., Brenner, D.L., 2002. A multiisotope C and N modeling analysis of soil organic matter turnover and transport as a function of soil depth in a California annual grassland soil chronosequence. Global Biogeochem. Cycles 16, 1135. http://dx.doi.org/10.1029/2001GB001823. Balesdent, J., Chenu, C., Balabane, M., 2000. Relationship of soil organic matter dynamics to physical protection and tillage. Soil Till. Res. 53, 215–230. Batjes, N.H., 1996. Total carbon and nitrogen in the soils of the world. Euro. J. Soil Sci. 47, 151–163. Becker-Heidmann, P., Scharpenseel, H.W., 1989. Carbon isotope dynamics in some tropical soils. Radiocarbon 31, 672–679. Bhattacharyya, T., Pal, D.K., Mandal, C., Velayutham, M., 2000. Organic carbon stock in Indian soils and their geographical distribution. Curr. Sci. 79, 655–660. Bhattacharyya, T., Ray, S.K., Pal, D.K., Chandran, P., Mandal, C., Wani, S.P., 2009. Soil carbon stocks in India – issues and priorities. J. Ind. Soc. Soil Sci. 57, 461–468. Boström, B., Comstedt, D., Ekblad, A., 2007. Isotope fractionation and 13C enrichment in soil profiles during the decomposition of soil organic matter. Oecologia 153 (1), 89–98. Bowman, D.M.J.S., Cook, G.D., Zoppi, U., 2004. Holocene boundary dynamics of a northern Australian monsoon rainforest patch inferred from isotopic analysis of

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