Printer Friendly

The effect of soil texture and roots on the stable carbon isotope composition of soil organic carbon.

Introduction

The carbon isotope composition ([[delta].sup.13]C value) of soil organic carbon (SOC) has provided information on past changes in vegetation (Dzurec et al. 1985; Pressenda et al. 1996; Schwartz et al. 1996), on the rates at which carbon from live biomass is cycled through the SOC pool following a change in land-use (Martin et al. 1990; Desjardin et al. 1994; Neill et al. 1996), and, in combination with atmospheric C[O.sub.2] [[delta].sup.13]C measurements, on the distribution of terrestrial C[O.sub.2] sources and sinks (Ciais et al. 1995; Fung et al. 1997; Battle et al. 2000).

Important considerations in such studies are the effect of local environmental conditions on the [[delta].sup.13]C value of carbon input to the soil from live vegetation, as well as the effects of soil physical parameters and microbial processes on the evolution of the 1[[delta].sup.13]C value of SOC with time. The [[delta].sup.13]C value of SOC changes as it is progressively either metabolised and respired, stabilised in the `slow' and `passive' SOC pools, or lost as dissolved organic carbon.

In forested areas that are dominated by [C.sub.3] plants the [[delta].sup.13]C values of standing biomass components are generally in the range -25 to -32 [per thousand]. Variations within this range are controlled by factors such as species (Ehleringer et al. 1987), latitude/altitude (Korner et al. 1991; Bird et al. 1994, 1996), soil water deficit (Stewart et al. 1995), irradiance (Ehleringer et al. 1986, 1987), topographic position (Balesdent et al. 1993), and the degree of re-utilisation of respired C[O.sub.2] (Schleser and Jayasekera 1985; Van der Merwe and Medina 1989; Broadmeadows et al. 1992). In addition, the various parts of an individual plant and the various compounds which make up different tissue types within a single plant can vary by several parts per thousand (e.g. Leavitt and Long 1986; Benner et al. 1987; Brugnoli and Farquhar 2000).

Surface litter [[delta].sup.13]C values are approximately equivalent to the [[delta].sup.13]C value of carbon deposited on the soil surface from live biomass (e.g. Balesdent et al. 1993). With increasing depth in the soil, SOC [[delta].sup.13]C values become modified from this initial value, with well-drained forest soils most commonly exhibiting a rapid rise of 1-2 [per thousand] in the upper soil layers, followed thereafter by a further, slower increase of 1-2 [per thousand] at depth (Martin et al. 1990; Balesdent et al. 1993; Desjardin et al. 1994; Ehleringer et al. 2000; Garten et al. 2000).

Two major processes are thought to compete in determining the [[delta].sup.13]C value of the carbon remaining in the soil as microbial metabolism proceeds. The selective utilisation of nutrient and energy rich high-[sup.13]C compounds such as carbohydrates, sugars, and proteins will increase the relative proportion of low-[sup.13]C compounds such as lignin in the remaining SOC (Benner et al. 1987). Conversely, kinetic fractionation accompanying microbial metabolism favours the partitioning of [sup.13]C into microbial biomass with the preferential respiration of [sup.12]C, thus increasing the [[delta].sup.13]C value of the remaining SOC (Macko and Estep 1984; Blair et al. 1985; Santruckova et al. 2000). Litter quality (C:N ratio) is thought to exert a major control on the degree to which isotope fractionation accompanies microbial decomposition (Agren et al. 1996).

A further external factor that has been postulated to affect the distribution of SOC [[delta].sup.13]C values with depth is the decrease in the [[delta].sup.13]C value of atmospheric C[O.sub.2] since industrialisation (Friedli et al. 1986). This means that `old' SOC will be variably enriched in [sup.13]C by up to 1.5 [per thousand] relative to carbon newly introduced to the SOC pool, a phenomenon known as the `terrestrial Suess effect' (Bird et al. 1996; Fung et al. 1997).

Despite these advances in understanding, it remains difficult to reconcile the observation from many studies that lignin is enriched by the early degradation of plant litter (e.g. Berg et al. 1984, 1993) with the observation that SOC [[delta].sup.13]C values increase with depth in many soil profiles.

Many previous studies have noted that the [[delta].sup.13]C value of SOC associated with fine mineral particles is higher than in coarser size fractions (e.g. Balesdent et al. 1987; Bonde et al. 1992; Desjardins et al. 1994; Balesdent and Mariotti 1996; Bird and Pousai 1997; Cayet and Lichtfouse 2001; Bird et al. 2002a, 2002b). Therefore, it seems logical that the bulk [[delta].sup.13]C value of any soil should be dependent upon the proportion of fine mineral material in the soil. The fundamental role that soil texture plays in determining soil carbon storage is well known (Schimel et al. 1994; Jobbagy and Jackson 2000; Bird et al. 2002a, 2002b) but the postulate that soil texture exerts significant control on bulk SOC [[delta].sup.13]C values has yet to receive any attention.

It was the purpose of this study to examine the role that soil texture may play in controlling the distribution of both SOC and [sup.13]C, by comparing soil profiles developed on deep sand and on a clay-rich substrate from locations with similar climate and topography. During the course of the study, it became apparent that a further overlooked potential control on the [[delta].sup.13]C value of soil organic carbon might be the [[delta].sup.13]C value of root-derived carbon. Accordingly, a preliminary attempt was made to assess this possibility through additional analyses and literature review.

Location and experimental methods

Site descriptions

The 2 study sites are located on Cape York Peninsula in Queensland, north-eastern Australia (Fig. 1). The `sand' site (12[degrees]39.29'S, 143[degrees]24.29'E, ~40 m a.s.l.) was located south-west of Cape Weymouth within Iron Range National Park, on the crest of a low Quaternary sand mass that comprised part of a prograding shoreline formed following a fall of sea-level during the Pleistocene. The `clay' site (15[degrees] 46.07'S, 145[degrees] 16.36'E, ~200 m a.s.l.) was on low-grade metamorphic strata of Carboniferous age. The climate at both sites is monsoonal with precipitation between January and March making up about 75% of the annual total of 1800-2500 mm. The mean annual temperature at the sand site is 25[degrees]C and at the clay site is 24[degrees]C.

[FIGURE 1 OMITTED]

The sand site was sampled in July 1997. The area was covered by semi-deciduous notophyll vine forest with a canopy height of approximately 15 m, a sparse understorey, and very sparse groundcover. A soil pit was dug to 1.2 m and augered to 2.5 m on the low dune crest. Slopes in the surrounding area were uniformly very shallow. Coarse wood and leaf litter was sparsely distributed over the soil surface. The soil comprises a white, podsolised quartz sand with the upper 60 cm slightly discoloured grey by organic matter (Australian Soil Classification: Aeric Podosol). Roots were very abundant in the upper 30 cm, and roots up to 3 cm in diameter were common down to 60 cm depth, with some fine roots present deeper in the profile. Below 90 cm the sand was loose and dry. Soil pH values (1:1 water:soil) decrease from 5.7 at the surface to 4.5 at 100 cm.

The clay site was sampled in April 2000. The area was located on the crest of a low ridge in an area covered by mesophyll vine forest with a canopy height of 25 m, a dense understorey, and a very sparse groundcover of small shrubs and ferns. Slopes in the vicinity of the site were low. A soil pit was dug to 1 m in fine red earth profile, beneath relatively abundant coarse wood and leaf debri. The soil graded from a heavy, grey, silty-clay in the upper 20 cm to brick red silty-clay at depth (Australian Soil Classification: Red Dermosol). Coarse roots were abundant to 40 cm depth, and below 50 cm blocks of partly weathered saprolite were sporadically encountered to 100 cm. The pH of the soil increased slightly from 5.8 at the surface to 6.2 at depth.

Sample preparation

Soil samples were taken horizontally into the exposed soil profile every 10 cm using a core tube with a length of 50 mm and an internal diameter of 34 mm. Surface 0-5 cm samples were taken vertically down from the soil surface after clearing away coarse surface litter. Samples were sealed in plastic bags immediately after collection to prevent loss of moisture.

Soil (40-60 g) from each sample was oven-dried at 70[degrees]C. Water content and bulk density were calculated by weight loss after drying. A comparatively low ultrasonic energy of 300 J/mL was used to disaggregate samples from both sites, as it was considered important not to `artificially' create fine particles in samples from the sand site. Schmidt et al. (1999) found no redistribution between particle sizes for energies of 30-590 J/mL, whereas Roscoe et al. (2000) found that at energies exceeding 260-275 J/mL unstable aggregates were completely disrupted while some stable aggregates may not have been completely disrupted. In the context of this study, it is therefore possible that some stable aggregates in samples from the clay site were not fully disaggregated, but if this were the case it would simply serve to mute the trends discussed below, with no impact on the overall conclusions.

Following ultrasonic disaggregation, the soil samples were separated into 4 fractions (<63 [micro]m, 63-500 [micro]m, 500-2000 [micro]m, >2000 [micro]m) and the <5 [micro]m fraction was separated from the 5-63 [micro]m fraction by differential settling in water. Recoveries from this procedure were >95% by weight in all cases. After drying and determination of the proportion of sediment present in each size fraction, aliquots were prepared for carbon abundance measurement and [sup.13]C-isotope analysis. Fractions >63 [micro]m were crushed before further analysis.

Some soil carbon-isotope studies attempt to quantify the isotopic composition of inputs of organic carbon to the soil by collecting samples of leaves and wood from the local area. However, the range in standing biomass [[delta].sup.13]C values (even in [C.sub.3]-dominated ecosystems) induced by different species, growth position, growth form, maturity, tissue type, local microclimate, and climate history is likely to be ~5 [per thousand] or more around a single sample site. To be of use to soil carbon studies this large range of values then must be appropriately weighted by each component's (unknown) contribution towards the total input to the soil. Therefore, it is not possible to accurately constrain the long-term [[delta].sup.13]C value of carbon added to the soil surface through the analysis of modern standing vegetation. Thus, for the purposes of this study sampling was limited to a comparison of [[delta].sup.13]C values of leaf, wood, and roots at the 2 sites, whereas the average [[delta].sup.13]C value of material added to the soil surface is considered to be most reliably recorded by the coarsest soil organic carbon in the uppermost soil layers.

Incubation experiments

A long-term incubation experiment was carried out in order to determine the effect of clay minerals on the amount and isotopic composition of clay minerals added to aliquots of soil from the sand site. A large sample of surface soil material (approximately 0-5 cm) from the area surrounding the soil pit at the sand site was used for this experiment. About 100 g of soil was subjected to repeated ultrasonic disaggregation and wet sieving at 63 [micro]m in order to completely remove all fine material from the sample. The >63 [micro]m fraction, thus cleaned, was then dried and split into ~5-g aliquots. A pure kaolin clay (from Weipa, Queensland) was added in amounts up to 27% by weight to some aliquots of >63 [micro]m sand soil material, while others were left free of clay. All the aliquots were then moistened with 9-10% by weight of water and incubated at 28[degrees]C for up to 1 year. After the incubation, the sample was gently sieved at 63 [micro]m and the <5 [micro]m fraction separated by settling in water. The separated fractions were dried, crushed, and subjected to carbon isotope analysis.

Isotope analysis

A weighed aliquot of every sample for carbon-isotope analysis was added to a quartz tube with CuO and Ag and the sample tube was then evacuated and flame-sealed. Organic carbon in the sample was oxidised to C[O.sub.2] by combustion at 900[degrees]C for 5 h. C[O.sub.2] was collected and purified cryogenically in a vacuum extraction line, with the quantity of C[O.sub.2] measured manometrically before collection into a break-seal tube for subsequent mass spectrometric analysis (Sofer et al. 1980). [[delta].sup.13]C values were determined using a Finnigan MAT-251 mass spectrometer and the results are reported as parts per thousand ([per thousand]) deviations from the V-PDB standard with an uncertainty of [+ or -] 0.1 [per thousand].

Results

Sand site

The sand site exhibited a range of dry bulk densities from 1.06 to 1.76 g/c[m.sup.3] with no regular trend down the profile (Table 1). In excess of 90% of the mineral particles at all depths were present in the 500-2000 [micro]m fraction with 3-8% present in the 63-500 [micro]m fraction, and <0.5% present in the <63 [micro]m fractions. The proportion of total carbon in the 500-2000 [micro]m fraction decreased from 87% at the surface to 60% at 50 cm depth, and SOC in the <63 [micro]m fractions increased from totals of ~5% at the surface to 20-30% at 100 cm.

Carbon densities decreased from 19.6 mg/c[m.sup.3] at the surface to <1.6 mg/c[m.sup.3] below 50 cm depth. These carbon densities translated into SOC inventories of 380 mg/c[m.sup.2] from 0-30 cm and 557 mg/c[m.sup.2] over 100 cm. The [[delta].sup.13]C value of all size fractions increased by about 1 [per thousand] in the first 10 cm, from values around-27.5 [per thousand] in the 0-5 cm interval. Beneath 10 cm, [[delta].sup.13]C values increased gradually down the profile by ~0.5 [per thousand] to 100 cm and a further -0.5 [per thousand] (to -25.5 [per thousand]) at 250 cm. There were no large differences between the isotopic compositions of the various size fractions, but the 5-63 [micro]m fraction was consistently 0.1-0.4 [per thousand] depleted in [sup.13]C relative to the other size fractions below 10 cm depth.

Clay site

In samples from clay site, soil densities increasd with depth from 0.72 g/c[m.sup.3] at the surface to 1.85 g/c[m.sup.3] at 100 cm (Table1). Over 60% of the mineral material was present in the <63 [micro]m fractions at the surface, rising to over 80% at the base of the profile. This distribution was mirrored in the distribution of SOC, with the finest (<5 [micro]m) fraction containing over 60% of the total SOC below 50 cm depth. Carbon densities decreased from 62.6 mg/c[m.sup.3] at the surface to 2.1 mg/c[m.sup.3] at 100 cm. These carbon densities indicate total SOC amounts of 1319 mg/c[m.sup.3] in the top 30 cm and 1725 mg/c[m.sup.3] over 100 cm.

The [[delta].sup.13]C value of bulk SOC increased dramatically with depth from -26.7 [per thousand] at the surface to values higher than -23 [per thousand] below 50 cm depth. In addition, [[delta].sup.13]C values increased with decreasing particle size, but values were extremely variable below 30 cm in the >63 [micro]m size fractions.

Discussion

Figure 2 shows the dramatic difference in the size distribution of mineral particles between the 2 study sites, which, given the similarities in climate and topographic position between the two, must constitute the likely source of differences in SOC distribution and isotopic composition observed at the 2 sites. The dramatic difference in soil texture between the sites is likely to have also resulted in differences in water-holding capacity, nutrient status, bioturbation rates, illuviation rates, oxygen diffusivity, litter quality, root penetration resistance, and net primary productivity. Some of these will affect the nature and quantity of organic matter delivered to the soil, while others will affect physical and chemical processes within the soil, but all can be considered to be dominantly controlled by soil texture when climate is invariant (Schimel et al. 1994; Jobbagy and Jackson 2000).

[FIGURE 2 OMITTED]

SOC inventories and particle size distributions

Figure 3 shows that carbon densities decrease regularly with increasing depth at both sites but that the decrease occurs more rapidly at the clay site. Whereas 76% of the total SOC to 100 cm is present in the soil above 30 cm at the clay site, the comparable figure is 68% at the sand site. This difference could reflect the greater permeability of the sand site, leading to a greater potential for the illuviation of SOC down the profile as well as a tendency for deeper rooting at the sand site in response to lower water-holding capacity, but both are within the usual bounds for tropical forest (Jobbagy and Jackson 2000).

[FIGURE 3 OMITTED]

Comparative inventories at the sand and clay sites over 0-30 cm are 380 and 1319 mg/c[m.sup.2], respectively, while over the full 100-cm interval, the inventories are 557 and 1725 mg/c[m.sup.2]. Thus, the inventories at the clay-rich site are 3.0-3.5 times the inventories at the sandy site and this is very likely to be predominantly an effect of the gross difference in soil texture between the sites.

Figure 4 shows that the quantities of SOC in the coarse (>63 [micro]m) fractions in the upper soil layers at the 2 sites are similar (both are 21.1 mg/c[m.sup.2] in the 0-5 cm interval), while over the full 100 cm interval, the quantities of SOC in the >63 [micro]m fractions at both sites are also similar at 500 [+ or -] 4 mg/c[m.sup.2]. In contrast, in the 0-5 cm interval the amount of <63 [micro]m SOC is 41.5 mg/c[m.sup.2] at the clay site but it is 50 times less than this value at the sand site (0.89 mg/c[m.sup.2]), and these large relative differences are maintained down through the profiles at the 2 sites. Thus, the very large differences in the total amount of SOC between the 2 sites is due primarily to the large proportion of SOC present in the <63 [micro]m fractions at the clay site, that is, to the difference in soil texture between the sites between the sites.

[FIGURE 4 OMITTED]

Figure 5 shows the relative contribution of SOC to the different size fractions at the sand and clay sites. In the surface layers, SOC in the <63 [micro]m fractions makes up <5% of the total at the sand site, whereas at the clay site the value is 66%. The proportion of <63 [micro]m SOC increases somewhat to 28% at the base of the sand profile, but increases dramatically to 93% at the base of the clay profile. Over 80% of the total SOC present at the base of the clay profile is present in the finest <5 [micro]m fraction.

[FIGURE 5 OMITTED]

These results point to the importance of fine mineral material in stabilising and protecting fine SOC produced by microbial degradation (Turchenek and Oades 1979; Anderson and Paul 1984; Mayer 1994). In the sand site, the lack of fine material means that there is little chance for the products of microbial degradation to be preserved by interaction with fine mineral surfaces. Hence, fine SOC is fully re-mineralised, progressively lost as dissolved organic carbon, or transported by illuviation to layers below the 100 cm interval. Analysis of samples obtained from up to 250 cm depth at the sand site suggest that appreciable SOC densities pertain below 100 cm, at about half the carbon density observed at 100 cm.

Carbon-isotope distributions

Figure 6 compares the [[delta].sup.13]C values for the bulk soil and all size fractions at the clay and sand sites. Both profiles exhibit some of the same features, in particular, both exhibit a rapid rise of about 1 [per thousand] between the 0-5 cm interval and 10 cm, followed by a slower rise to higher values at increasing depth. At both sites the [[delta].sup.13]C value of SOC in the coarsest 500-2000 gm fraction in the 0-5 cm interval is approximately -27.5 [per thousand], and as discussed previously, this material is most likely to best represent the long-term average [[delta].sup.13]C value of organic carbon input to the soil from above-ground biomass. There would therefore appear to be little difference in the [[delta].sup.13]C value of organic carbon input to the soil from above-ground biomass at the two sites.

At the sand site, there is comparatively little variation in [[delta].sup.13]C values between size fractions, and all size fractions exhibit only a small increase in [[delta].sup.13]C value beneath 5 cm in depth (-0.5 [per thousand] from 10-100 cm). A study of 8000 cores from coarse-textured soils from widely different climates estimated that the difference in [[delta].sup.13]C value between the 0-5 cm and 5-30 cm interval averaged 0.9 [+ or -] 0.6 [per thousand], suggesting that the results from this study are in agreement with other data from coarse-textured soils (Bird et al. 2001).

The 5-63 [micro]m fraction at the sand site has the lowest [[delta].sup.13]C value at any depth (average -0.26 [+ or -] 0.14 per thousand] relative to SOC in the >63 [micro]m fractions). As there is little mineral material, most of the SOC in this size fraction is likely to be fine particulate SOC. The low [[delta].sup.13]C values in this size fraction suggest that it may be the degraded, resistant residue after microbial decomposition of the coarser material, comparatively rich in low-[sup.13]C compounds such as lignin (Benner et al. 1987).

The <5 [micro]m fraction at the sand site exhibits [[delta].sup.13]C values similar to SOC in the >63 [micro]m fractions (average +0.02 [+ or -] 0.26 [per thousand] relative to SOC in the >63 [micro]m fractions). This similarity reflects the overall paucity of fine mineral material and thus the low preservation potential of microbially metabolised carbon. Nevertheless, the [[delta].sup.13]C values in the <5 [micro]m fraction are higher than observed in the comparable 5-63 [micro]m fractions due to a larger relative proportion of high-[sup.13]C carbon of microbial origin associated with the small amount of fine mineral material present.

In contrast to the sand site, the [[delta].sup.3]C values of SOC in all size fractions at the clay site continue to increase dramatically below the surface layers, but with some large, erratic changes in [[delta].sup.13]C value in the >63 [micro]m fractions (these fractions only make up 2-5% of the total SOC below 15 cm depth). Also in contrast to the sand profile, the [[delta].sup.13]C value of the 5-63 [micro]m fraction in the clay profile is similar to the [[delta].sup.13]C value of the >63 [micro]m fractions (+0.03 [+ or -] 1.0 [per thousand] relative to the >63 [micro]m fractions) and the [[delta].sup.13]C value of the <5 [micro]m fraction is on average 1.87 [per thousand] higher than SOC in the >63 [micro]m fractions.

Some very high [[delta].sup.13]C values for some fractions at around 60 cm depth (up to -21.6 [per thousand]), as well as the large erratic shifts from -22.1 to -26.4 [per thousand] in the coarse fractions, suggest the possibility that a shift in vegetation has occurred at the clay site at some time in the past. The site was chosen because it is located in a wet, tropical region, which should have been forested for most of the Holocene (Kershaw 1978). Nevertheless, it seems that some carbon remains from a more open vegetation type established over the site at some earlier time that included some [C.sub.4] grasses.

This conclusion complicates the interpretation of soil texture effects on [sup.13]C distribution at the clay site, but higher [[delta].sup.13]C values for SOC in fine than in coarser particle sizes have been widely reported for other soils under pure [C.sub.3] vegetation (Table 2). Therefore, the observed high [[delta].sup.13]C values in the fine particle sizes relative to the coarse particle sizes at the clay site are not likely to be solely, or even dominantly, an artefact of vegetation change.

The major differences between the sand and clay profiles lie in the [[delta].sup.13]C value of carbon in the <63 [micro]m fractions, in particular, a small depletion in [sup.13]C in the 5-63 [micro]m fraction at the sand site and a large enrichment in [sup.13]C in the <5 [micro]m fraction at the clay site. These relative differences are maintained throughout the entire profiles at both locations. It is likely that the presence of fine mineral material at the clay site serves to protect and stabilise the products of microbial degradation, which have a higher [[delta].sup.13]C value than bulk SOC because (i) microbes fractionate carbon isotopes between biomass and respired C[O.sub.2] during metabolism, and/or (ii) microbes preferentially utilise high-[sup.13]C compounds.

The net effect of these isotopic differences between the sand and clay sites is that although the [[delta].sup.13]C value of carbon input to the soil is similar at both locations, the bulk [[delta].sup.13]C value of SOC in the 0-5 cm interval of the sand site is-27.3 [per thousand], while at the clay site it is -26.7 [per thousand], this difference being attributable to the `extra' high-[sup.13]C carbon present in the fine fractions at the clay site. The same comparisons over 0-30 cm are -26.7 and -25.1 [per thousand], respectively. Over the full 0-100 cm interval the comparable values are -26.5 [per thousand] and -24.5 [per thousand], although as noted above there may be a small [C.sub.4]-derived component at depth at the clay site. Thus, the effect of fine minerals at the clay site is to increase the [[delta].sup.13]C value of bulk SOC by 0.6-2.0 [per thousand] relative to the sand site.

In order to verify the hypothesis that fine minerals promote the preservation of high-[sup.13]C, microbially processed carbon in the fine fraction of soil, aliquots of >63 [micro]m sand material were incubated with and without added clay for up to a year. A summary of the results of these experiments is presented in Table 3. Both the >63 [micro]m and <5 [micro]m fractions of the original sample had a `natural' [[delta].sup.13]C value of-26.8 [per thousand], but after a year of incubation, there were significant differences attributable to the presence or absence of clay in the sample. In the cases where no clay was added, only 2.2 [+ or -] 1 mg/g of original carbon had accumulated in the <5 [micro]m fraction, and the [[delta].sup.13]C value of this carbon was only 0.2 [per thousand] higher than that of the original `indigenous' <5 [micro]m carbon in the sample. In contrast, in samples that had `added clay', 8 times as much carbon (16.3 [+ or -] 5 mg/g) had accumulated in the <5 [micro]m fraction and this carbon had a [[delta].sup.13]C value 0.7[per thousand] higher than that of the original `indigenous' <5 [micro]m carbon in the sample, similar to the difference observed in natural samples (Table 2).

There were no significant differences in either respiration rate or the isotopic composition of the respired C[O.sub.2] between the `clay' and `no clay' samples (M. I. Bird, unpublished data). Although a full carbon balance is not possible, because only a few per cent of the total carbon in the samples was metabolised over the course of the experiment, the data provide some further support for the hypothesis that clay minerals can protect a fraction of microbially processed, higher [sup.13]C carbon from further degradation.

Changes in the [[delta].sup.13]C value of SOC above 10 cm depth Many studies have demonstrated that there is a rapid enrichment of 0.5-1.0 [per thousand] between the surface and 5-10 cm depth, and this change is usually ascribed to the effects of microbial degradation. The problem with such an interpretation is that SOC is most abundant and least degraded at the soil surface, and it has been demonstrated that the microbial population preferentially metabolises nutrient and energy-rich high-[sup.13]C compounds, leaving the remaining SOC enriched in low-[sup.13]C compounds such as lignin (Berg et al. 1984, 1993; Benner et al. 1987). Indeed, the [delta][sup.13]C values of the <63 [micro]m fractions in the 0-5 cm internal of the sand site are 0.3 to 0.6 [per thousand] lower than the bulk SOC, suggesting that the effect of microbial degradation in the absence of fine mineral material is to deplete the residual, unmetabolised carbon in [sup.13]C.

A further possibility for the rapid initial rise in [delta][sup.13]C value that has yet to receive any attention relates to root-derived carbon. Below-ground primary production is a significant component of total primary production in forests and root/shoot ratios in tropical forest ecosystems generally range from 0.19 to 0.34 (Jackson et al. 1996). Therefore, while above-ground carbon inputs to the soil are dominated by leaf material (Ross 1993), considerable direct below-ground carbon inputs are derived from root material.

Table 4 summarises published data on leaf and root isotope compositions for a range of species. For [C.sub.3] plants, the average difference in [delta][sup.13]C value between above-ground biomass (leaves) and below-ground biomass (roots) is + 1.2 [+ or -] 0.8 [per thousand] (1 [sigma]; n = 12). Garten et al. (2000) examined the isotopic composition of litterfall and roots from 6 temperate forest sites and found that roots had [delta][sup.13]C values that were on average +1.1 [+ or -] 0.5 [per thousand] higher than litterfall from the same plot. Natelhoffer and Fry (1988), on the other hand, found the difference in [delta][sup.13]C value between 2 samples of `fine roots' (<0.5 mm) and 6 samples of leaf litter from 2 oak forest plots was -0.5 [+ or -] 0.4 [per thousand]. Leavitt and Long (1986) have demonstrated that woody tissues such as those in structural roots have [delta][sup.13]C values that are 1-2 [per thousand] higher than leaves from the same plant, but there is as yet no systematic information on differences in [delta][sup.13]C value that might exist between fine roots, coarse roots, and root exudates.

Although further work will be required to determine whether there are systematic differences between the [delta][sup.13]C values amongst the components of below-ground biomass, the available data overwhelmingly suggest that it has a significantly higher [delta][sup.13]C value than above-ground biomass. This difference in isotopic composition should result in a rapid increase in [delta][sup.13]C value from the surface litter layers to the upper soil layers where root densities are highest. The current study was not designed to assess in detail the isotopic composition of live biomass, and sampling of above- and below-ground biomass from the same plants was not undertaken. Nevertheless, from the limited number of samples analysed from both the sand and clay sites, coarse roots had the highest [delta][sup.13]C values (-26.2 to -27.0 [per thousand]; n = 9), whereas wood, leaves, and fine roots had variably lower [delta][sup.13]C values, ranging from -27.7 to -32.6 [per thousand] (n = 15). Root distributions were not quantitatively determined, but field observations suggest that, as for other tropical forests (Jackson et al. 1996), root biomass decreased below the surface to 40-50 cm depth at both sites. No coarse roots were found below 50cm at either site, although some fine roots persisted to the base of both profiles.

The results from this study, as well as previously published results, are therefore consistent with the possibility that a systematic difference in isotopic composition of carbon delivered to the soil from above- v. below-ground sources may account for part of the initial rapid rise in [delta][sup.13]C value with depth observed in many soils. Further detailed studies of live forest biomass carbon-isotope allocation both above and below ground (coarse root, fine root, and root exudates) will be required to properly explore this suggestion.

The observed rapid increase in the upper soil layers must also include a component due to the changing [delta][sup.13]C value of atmospheric carbon dioxide since industrialisation (Friedli et al. 1986). However, in the moist tropical climate pertaining at both sites in this study, SOC residence times in the surface soil are likely to be less than a decade (Bird et al. 1996). Thus the `terrestrial Suess effect' is unlikely to have contributed more than 0.1-0.3 [per thousand] to the change observed in the top 10 cm at both sites.

Changes in the [delta][sup.13]C value of SOC below 10 cm depth

The full isotopic effect, resulting from the changing [delta][sup.13]C value of the atmosphere (~ 1.5 [per thousand]) since industrialisation, must be present over an unknown depth interval in both soils, assuming that at some depth most carbon is older than 150 years (O'Brian and Stout 1978; Townsend et al. 1995; Pressenda et al. 1996). At the sand site the full magnitude of the change in the bulk [delta][sup.13]C value of the SOC from 10 cm to 250 cm is >1.5 [per thousand], suggesting that there is little need even to invoke microbial alteration to explain the trend in the deeper soil layers. The slow rise in [delta][sup.13]C value with increasing depth can be explained simply by the progressive mixing of `young' SOC at the surface with `old' SOC at depth. The observation that the [delta][sup.13]C value of SOC in all size fractions increases with depth supports this hypothesis. The decreasing abundance of roots at increasing depth may provide a mechanism for mixing `young' carbon into the deep soil in progressively lower amounts, as well as the possible illuviation of younger carbon down the profile.

Interpretation of the observed increase in [delta][sup.13]C value with depth at the clay site is complicated by the likelihood of past vegetation change. It is, nevertheless, possible to speculate that a further effect of fine minerals is that [high-.sup.13]C SOC, previously metabolised by microbes and partly stabilised by association with fine minerals, can itself be slowly metabolised, with the resultant metabolites themselves re-stabilised by fine minerals. Over time, this process might lead to a further increase in the [delta][sup.13]C value of bulk SOC in the deep soil, beyond that expected from the changing [delta][sup.13]C value of the atmosphere. This might explain the fact that the differences of >1.6 [per thousand] between bulk SOC in the upper soil and the deep soil are commonly reported in soils where appreciable quantities of fine minerals are present.

Conclusions

This study provides support for the contention that a major factor influencing both the inventory of SOC and the [delta][sup.13]C value bulk SOC in a soil profile is the texture of the soil, and in particular the abundance of fine mineral particles. Over 3 times as much SOC is present in the clay profile over 100 cm as is present in the sand profile, and this appears to be entirely attributable to the large amount of SOC present in the fine size fractions at the clay site. The differences in the bulk [delta][sup.13]C value of SOC at the two sites are also due to the large amount of high-[sup.13]C SOC present in the fine size fractions at the clay site.

A comparison of trends in [delta][sup.13]C value between particle sizes and with depth at the 2 sites allows inferences to be drawn regarding the processes that control the commonly observed increase in [delta][sup.13]C value with depth in many soil profiles. It is postulated that an initial rise of ~1-1.5 [per thousand] in the surface layers of many soils in all particle sizes may be due to an increase in the relative contribution of root-derived carbon beneath the soil surface. A survey of the literature and limited data from this study strongly support the hypothesis that below-ground (root-derived) carbon has an average [delta][sup.13]C value that is higher than the average [delta][sup.13]C value of the above-ground biomass (mostly leaves) that provides the surface inputs of carbon to the soil. Further systematic study of carbon-isotope allocation to above-and below-ground biomass in forest ecosystems will be required to unequivocally confirm this result, although there is some evidence that such a difference could relate to malic acid synthesis by roots (Popp et al. 1982).

The results from field studies and laboratory incubations presented above suggest that in the presence of fine minerals, a proportion of high-[sup.13]C, microbially processed products is stabilised by adsorption onto fine mineral surfaces or the through formation of organo-mineral complexes (Turchenek and Oades 1979; Anderson and Paul 1984; Mayer 1994). Microbial metabolism of SOC also fractionates [sup.13]C preferentially into microbial biomass (Macko and Estep 1984; Blair et al. 1985), and therefore ultimately, into compounds with relatively high [delta][sup.13]C values that can be stabilised by organo-mineral interactions. In the absence of fine minerals, the products of microbial degradation are themselves rapidly remineralised.

The results have implications for the estimation of the carbon-isotope composition of terrestrial at the regional/global scale, as the bulk [delta][sup.13]C value of SOC cannot simply be inferred from the [delta][sup.13]C value of standing vegetation. Additional factors, such as variations in soil texture and climate-dependant variations in the average age of SOC, must also be considered.
Table 1. Basic characteristics, particle-size distribution,
SOC size distribution, and carbon-isotope composition for
the soil profiles at the sand and clay sites

 Bulk Carbon Carbon
 density Carbon density inventory
 [H.sub.2]O (g/ content (mg/ (mg/
 (%) [cm.sub.3]) (%) [cm.sub.3]) [cm.sub.2])

 Sand site

 0-5 6.2 1.25 1.57 19.58 97.9
 5-15 6.0 1.06 1.32 13.93 139.3
 15-25 3.9 1.33 0.79 10.49 104.9
 25-35 4.3 1.33 0.57 7.63 76.3
 35-45 3.8 1.57 0.34 5.40 54.0
 45-55 3.8 1.76 0.14 2.51 25.1
 55-65 3.8 1.49 0.11 1.57 15.7
 65-75 3.4 1.56 0.10 1.52 15.2
 75-85 3.6 1.55 0.09 1.35 13.5
 85-95 3.9 1.48 0.07 1.04 10.4
 95-100 3.3 1.55 0.05 0.85 4.2
105-115 2.9 1.57 0.06
115-125 3.3 1.30 0.05
150-200 1.0 1.29 0.03
200-250 0.9 1.38 0.02

 Clay site

 0-5 35.0 0.72 8.64 62.63 313.2
 5-15 24.5 1.07 5.38 57.67 576.7
 15-25 24.2 1.15 2.86 32.91 329.1
 25-35 22.4 1.29 1.54 19.91 199.1
 35-45 20.3 1.40 0.88 12.24 122.4
 45-55 19.4 1.47 0.42 6.21 62.1
 55-65 18.3 1.37 0.29 3.99 39.9
 65-75 18.0 1.50 0.18 2.76 27.6
 75-85 15.5 1.59 O.17 2.71 27.1
 85-95 17.7 1.65 0.11 1.76 17.6
 95-100 17.7 1.85 0.11 2.10 10.5

 Particle size distribution (%)

 <5 5-63 63-500 500-2000
 [micro]m [micro]m [micro]m [micro]m

 Sand site

 0-5 0.05 0.16 4.5 95.3
 5-15 0.04 0.34 8.1 91.5
 15-25 0.02 0.21 7.9 91.9
 25-35 0.04 0.22 6.3 93.5
 35-45 0.05 0.25 3.2 96.5
 45-55 0.05 0.27 4.7 95.0
 55-65 0.06 0.32 5.5 94.1
 65-75 0.07 0.42 4.9 94.6
 75-85 0.07 0.43 5.3 94.2
 85-95 0.08 0.38 4.2 95.3
 95-100 0.07 0.26 3.5 96.1
105-115 0.06 0.34 4.5 95.1
115-125 0.06 0.33 7.4 92.2
150-200 0.06 0.33 5.4 94.2
200-250 0.05 0.33 4.1 95.5

 Clay site

 0-5 25.42 38.36 32.63 3.59
 5-15 27.47 39.20 27.47 5.87
 15-25 32.63 43.71 21.56 2.10
 25-35 36.18 33.99 27.85 1.97
 35-45 40.94 39.70 16.87 2.48
 45-55 34.41 31.38 30.36 3.85
 55-65 34.53 36.92 27.81 0.74
 65-75 39.40 44.20 16.08 0.32
 75-85 32.71 40.40 26.61 0.29
 85-95 37.45 46.54 15.64 0.37
 95-100 39.68 45.09 14.88 0.36

 Carbon size distribution (%)

 <5 5-63 63-500 500-2000
 [micro]m [micro]m [micro]m [micro]m

 Sand site

 0-5 1.1 3.4 8.0 87.5
 5-15 1.2 6.3 10.1 82.4
 15-25 0.8 5.0 24.2 70.0
 25-35 2.0 3.4 13.6 81.0
 35-45 4.1 6.8 10.0 79.1
 45-55 7.0 10.4 13.3 69.3
 55-65 8.4 18.7 13.6 59.2
 65-75 6.9 16.4 11.1 65.6
 75-85 11.9 15.7 11.4 61.0
 85-95 15.7 18.7 12.5 53.2
 95-100 10.7 17.2 10.8 61.3
105-115 6.5 9.0 9.7 74.8
115-125 8.3 12.6 15.6 63.6
150-200 5.0 9.9 10.5 74.6
200-250 4.7 7.7 8.3 79.2

 Clay site

 0-5 28.1 38.2 25.6 8.1
 5-15 35.7 25.3 20.5 18.5
 15-25 59.7 20.9 15.3 4.2
 25-35 66.5 21.5 11.2 0.7
 35-45 61.0 6.3 30.8 1.8
 45-55 56.3 8.4 30.5 4.8
 55-65 61.0 14.9 23.1 1.0
 65-75 86.8 10.8 2.4 0
 75-85 67.2 14.0 18.8 0
 85-95 84.1 14.4 1.5 0
 95-100 80.5 12.6 6.9 0

 [delta][sup.13]C ([per thousand])

 Bulk
 <5 5-63 63-500 500-2000 <2000
 [micro]m [micro]m [micro]m [micro]m [micro]m

 Sand site

 0-5 -27.5 -27.6 -27.1 -27.3 -27.3
 5-15 -26.5 -26.6 -26.8 -26.5 -26.6
 15-25 -26.2 -26.6 -26.3 -26.4 -26.4
 25-35 -26.3 -26.5 -26.4 -26.2 -26.2
 35-45 -25.9 -26.5 -26.4 -26.1 -26.2
 45-55 -25.9 -26.4 -26.2 -26.1 -26.2
 55-65 -26.0 -26.2 -26.1 -26.1 -26.1
 65-75 -26.0 -26.2 -26.1 -26.0 -26.0
 75-85 -26.1 -26.3 -25.9 -26.0 -26.1
 85-95 -26.0 -26.3 -25.8 -25.9 -26.0
 95-100 -25.9 -26.3 -26.0 -25.8 -25.9
105-115 -25.7 -25.8 -25.5 -25.7 -25.7
115-125 -25.8 -25.9 -25.4 -25.9 -25.8
150-200 -25.7 -25.6 -25.4 -25.2 -25.3
200-250 -25.2 -25.2 -25.5 -25.3 -25.3

 Clay site

 0-5 -25.2 -27.3 -27.1 -27.6 -26.7
 5-15 -23.9 -25.4 -25.4 -25.7 -24.9
 15-25 -23.3 -25.3 -25.8 -25.9 -24.2
 25-35 -22.6 -25.4 -26.0 -26.1 -23.6
 35-45 -21.6 -24.4 -26.4 -24.7 -23.3
 45-55 -21.6 -24.1 -22.8 -24.0 -22.3
 55-65 -21.2 -23.8 -22.1 -25.1 -21.8
 65-75 -21.6 -23.5 -25.6 - -21.9
 75-85 -21.9 -23.6 -23.5 - -22.4
 85-95 -22.3 -23.5 -26.4 - -22.5
95-100 -22.3 -24.1 -26.5 - -22.8

Table 2. Compilation from the literature of differences in
carbon-isotope composition between the coarsest and finest
particle sizes analysed (positive difference indicates fine
size fractions have higher [delta][sup.13]C values than
coarse size fractions)

Soil type, total depth of sampling, and size fractions
analysed also shown

Soil Coarse Fine
 ([micro]m) ([micro]m)

Hapludalf (0-40 cm) 200-2000 <0.2
Typic Haplorthox (0-12 cm) 200-2000 <2
Typic Kandiudults (0-40 (cm) 200-2000 <0.2
Various (0-2 cm) 500-2000 <2
Hapludalf (0-40 cm) 200-2000 <50
Various (`topsoil') 200-2000 <0.2
Various (0-5 cm) 500-2000 <2

Soil Difference
 ([per thousand])

Hapludalf (0-40 cm) +1.0 [+ or -] 0.3 (2)
Typic Haplorthox (0-12 cm) +1.2 [+ or -] 0.0 (2)
Typic Kandiudults (0-40 (cm) +2.2 [+ or -] 0.6 (4)
Various (0-2 cm) +0.9 [+ or -] 0.5 (6)
Hapludalf (0-40 cm) +1.4 (1)
Various (`topsoil') +1.3 (7)
Various (0-5 cm) +0.8 [+ or -] 0.5 (12)

Soil Reference

Hapludalf (0-40 cm) Balesdent et al. (1987)
Typic Haplorthox (0-12 cm) Bonde et al. (1992)
Typic Kandiudults (0-40 (cm) Desjardins et al. (1994)
Various (0-2 cm) Bird and Pousai (1997)
Hapludalf (0-40 cm) Cayet and Lichtfouse (2001)
Various (`topsoil') Balesdent and Mariotti (1996)
Various (0-5 cm) Bird et al. (2002b)

Table 3. Carbon-isotope composition of >63 [micro]m and <5 [micro]m
material from the original samples of surface soil from the sand
site, and the comparable results after up to one year's
incubation, both with and without added clay (see text)

Also shown is the relative amount of original carbon transferred
from the >63 [micro]m fraction to the <5 [micro]m fraction over
the incubation period. Results are corrected for the small amount
of organic carbon present in the high purity kaolin clay that was
to the `clay added' aliquots. In all cases the original `bulk'
sample contained 410 [+ or -] 20 mg carbon with a carbon-isotope
composition of -26.85 [+ or -] 0.06 ([per thousand])

 Original sample No clay added Clay added

[delta][sup.13]C
 >63 [micro]m -26.83 [+ or -] -27.03 [+ or -] -27.02 [+ or -]
 ([per thousand]) 0.10 (4) 0.03 (4) 0.02 (8)
[delta][sup.13]C
 >5 [micro]m -26.84 [+ or -] -26.66 [+ or -] -26.09 [+ or -]
 ([per thousand]) 0.08 (2) 0.15 (4) 0.28 (8)
Transfer to <5
 [micro]m (mg/g) -- 2.2 [+ or -] 1.0 16.3 [+ or -] 5

Table 4. Compilation from the literature of differences in
carbon-isotope composition between above-ground and below-ground
biomass (A/B) or leaves and roots (L/R) of the same plant for a
range of [C.sub.3] and [C.sub.4] species

Positive difference indicates that below-ground biomass (or roots)
has a higher [delta][sup.13]C value than above-ground biomass
(or leaves)

 Difference
Plant Pathway Comparison ([per thousand])

Desmodium ovifolium C3 A/B +1.5 [+ or -] 0.2
Agropyron C3 A/B +1.2 [+ or -] 0.05
Poa C3 A/B +0.7 [+ or -] 0.1
Agrostis C3 A/B -0.1 [+ or -] 0.1
Rice C3 L/R +1.1 [+ or -] 0.4
Tomato C3 L/R +1.0 [+ or -] 0.6
Sunflower C3 L/R +0.5 [+ or -] 0.2
Peanut C3 L/R +0.4
Beet C3 L/R +1.3
Phragmites australis C3 L/R +2.5 [+ or -] 0.3
Picea abies C3 L/R +1.8 [+ or -] 1.0
Citrus volkameriana C3 L/R +2.4 [+ or -] 0.3
Citrus aurantium C3 L/R +2.4 [+ or -] 0.3
Schizachyrium C4 A/B +0.4 [+ or -] 0.05
Brachiaria humidicola C4 L/R -0.5
Zea mays C4 L/R +1.5 [+ or -] 0.5
 +0.4

Plant Reference

Desmodium ovifolium Schweizer et al. (1999)
Agropyron Wedin et al. (1995)
Poa Wedin et al. (1995)
Agrostis Wedin et al. (1995)
Rice Summarised in Brugnoli and Farquhar (2000)
Tomato ,,
Sunflower ,,
Peanut ,,
Beet ,,
Phragmites australis ,,
Picea abies ,,
Citrus volkameriana Syvertson et al. (1997)
Citrus aurantium Syvertson et al. (1997)
Schizachyrium Wedin et al. (1995)
Brachiaria humidicola Schweizer et al. (1999)
Zea mays Balesdent et al. (1987)
 Schonwitz et al. (1986)


Acknowledgments

We thank L.B. Law for assistance with sample collection, and J. Cowley and K. McAllister for assistance with aspects of the sample preparation. The provision of visiting fellowships by the Max Planck Institute fur Biogeochemie and the National University of Singapore to MB greatly facilitated the completion of this manuscript.

References

Agren GI, Bosatta E, Balesdent J (1996) Isotope discrimination during decomposition of organic matter: a theoretical analysis. Soil Science Society of America Journal 60, 1121-1126.

Anderson DW, Paul EA (1984) Organo-mineral complexes and their study by radiocarbon dating. Soil Science Society of America Journal 48, 298-301.

Balesdent J, Girardin C, Mariotti A (1993) Site-related [delta][sup.13]C of tree leaves and soil organic matter in a temperate forest. Ecology 74, 1713-1721.

Balesdent J, Marriotti A (1996) Measurement of soil organic matter turnover using [sup.13]C natural abundances. In `Mass spectrometry of soils'. (Eds TW Boutton, SI Yamasaki) pp. 83-111. (Marcel Dekker Inc.: New York)

Balesdent J, Mariotti A, Guillet B (1987) Natural [sup.13]C abundance as a tracer of soil organic matter dynamics. Soil Biology and Biochemistry 19, 25-30.

Battle M, Bender ML, Tans PP, White JCW, Ellis JT, Conway T, Francey RJ (2000) Global carbon sinks and their variability inferred from atmospheric [O.sub.2] and [delta][sup.13]C. Science 287, 2467-2470.

Benner R, Fogel ML, Sprague EK, Hodson RE (1987) Depletion of [sup.13]C in lignin and its implication for stable isotope studies. Nature 329, 708-710.

Berg B, Ekbohm G, McClaugherty C (1984) Lignin and holocellulose relations during long-term decomposition of some forest litters: Long-term decomposition in a Scots pine forest IV. Canadian Journal of Botany 62, 2540-2550.

Berg B, McClaugherty C, Johansson M-B (1993) Litter mass-loss rates in late stages of decomposition at some climatically and nutritionally different pine sites: Long-term decomposition in a Scots pine forest. VIII. Canadian Journal of Botany 71,680-692.

Bird MI, Chivas AR, Head J (1996) A latitudinal gradient in carbon turnover times in forest soils. Nature 381, 143-146.

Bird MI, Haberle SG, Chivas AR (1994) Effect of altitude on the carbon-isotope composition of forest and grassland soils from Papua New Guinea. Global Biogeochemical Cycles 8, 13-22.

Bird MI, Pousai P (1997) Variations of [delta][sup.13]C in the surface soil organic carbon pool. Global Biogeochemical Cycles 11, 313-322.

Bird MI, Santruckova H, Arneth A, Lloyd JJ, Gleixner G, Schulze ED (2002a) Inventories of carbon and isotopes on a latitude gradient through central Siberia. Tellu 54B (in press).

Bird MI, Santruckova H, Lloyd JJ, Lawson E (2002b) The isotopic composition of soil organic carbon on a latitude transect in western interior Canada. European Journal of Soil Science 53, 393-403.

Bird MI, Zhou Y, Carter J, Farquhar GH (2001) Global-scale variations in the carbon-isotope composition of soil organic carbon in coarse-textured soils. In `Proceedings of the 5th International Carbon Dioxide Conference'. Sendai, Japan, pp. 313.

Blair N, Leu A Munos E Olsen J, Kwong E, des Marais D (1985) Carbon isotopic fractionation in heterotrophic microbial metabolism. Applied Environmental Microbiology 50, 996-1001.

Bonde TA, Christensen BT, Cerri CC (1992) Dynamics of soil organic matter as reflected by natural [.sup.13]C abundance in particle size fractions of forested and cultivated oxisols. Soil Biology and Biochemistry 24, 275-277.

Broadmeadows MS J, Griffiths H, Maxwell C, Borland AM (1992) The carbon isotope ratio of plant organic matter reflects temporal and spatial variations of C[O.sub.2] within tropical forest formations in Trinidad. Oecologia 89, 435-441.

Brugnoli E, Farquhar GD (2000) Photosynthetic fractionation of carbon isotopes. In `Photosynthesis: physiology and metabolism'. (Eds RC Leegood, S von Caemmerer) pp. 399-434. (Kluwer: The Netherlands)

Cayat C, Lichtfouse E (2001) [delta][sup.13]C of plant-derived n-alkanes in soil particle-size fractions. Organic Geochemistry 32, 253-258.

Ciais P, Tans PP, White JCW, Trolier M, Francey R J, Berry JA, Randall DR, Sellers P J, Collatz JG, Schimel DS (1995) Partitioning of ocean and land uptake of C[O.sub.2] as inferred by delta-C-13 measurements from the NOAA Climate Monitoring and Diagnostics Laboratory Global Air Sampling Network. Journal of Geophysical Research 100, 5051-5070.

Desjardins T, Andreaux F, Volkoff B, Cerri CC (1994) Organic carbon and [sup.13]C contents in soils and soil size-fractions, and their changes due to deforestation and pasture installation in eastern Amazonia. Geoderma 61, 103-118,

Dzurec RS, Boutton TW, Caldwell MM, Smith BN (1985) Carbon isotope ratios of soil organic matterand their use in assessing community composition changes in Curlew Valley, Utah. Oecologia 66, 17-24.

Ehleringer JR, Buchmann N, Flanagan LB (2000) Carbon isotope ratios in belowground carbon cycle processes. Ecological Applications 10, 412-422.

Ehleringer JR, Field CB, Lin Z-F, Kuo C-Y (1986) Leaf carbon isotope and mineral composition in subtropical plants along an irradiance cline. Oecologia 70, 520-526.

Ehleringer JR, Lin Z-F, Field CB, Sun GC, Kuo C-Y (1987) Leaf carbon isotope ratios from a subtropical monsoon forest. Oecologia 72, 109-114.

Friedli H, Lotscher H, Oeschger H, Siegenthaler U, Stauffer B (1986) Ice core record of the [sup.13]C/[sup.12]C ratio of atmospheric C[O.sub.2] in the past two centuries. Nature 324, 237-238.

Fung I, Field CB, Berry JA, Thompson MV, Randerson JT, Malmstrom CM, Vitousek PM, Collatz GJ, Sellers PJ, Randall DA, Denning AS, Badeck F, John J (1997) Carbon 13 exchanges between the atmosphere and biosphere. Global Biogeochemical Cycles 11,507-533.

Garten CT, Cooper LW, Post WM, Hanson PJ (2000) Climate controls on forest soil C isotope ratios in the southern Appalachian Mountains. Ecology 81, 1108-1119.

Jackson RB, Canadell J, Ehleringer JR, Mooney HA, Sala OE, Schulze ED (1996) A global analysis of root distributions for terrestrial biomes. Oecologia 108, 389-411.

Jobbagy EG, Jackson RB (2000) The vertical distribution of soil organic carbon and its relation to climate and vegetation. Ecological Applications 10, 423-436.

Kershaw AP (1978) Record of last interglacial-glacial cycle from northeastern Queensland. Nature 272, 159-161.

Korner Ch, Farquhar GD, Wong SC (1991) Carbon isotope discrimination by plants follows latitudinal and altitudinal trends. Oecologia 88, 30-40.

Leavitt SW, Long A (1986) Stable-carbon isotope variability in tree foliage and wood. Ecology 67, 1002-1010.

Macko SA, Estep MLF (1984) Microbial alteration of stable nitrogen and carbon isotopic composition of organic matter. Organic Geochemistry 6, 787-790.

Martin A, Mariotti A, Balesdent J, Lavelle P, Vuattoux R (1990) Estimate of organic matter turnover rate in savanna soil by [sup.13]C natural abundance measurements. Soil Biology and Biochemistry 22, 517-523.

Mayer LM (1994) Relationships between mineral surfaces and organic carbon concentrations in soils and sediments. Chemical Geology 114, 347-363.

Natelhoffer KJ, Fry B (1988) Controls on natural nitrogen-15 and carbon-13 abundances in forest soil organic matter. Soil Science Society of America Journal 52, 1633-1640.

Neill C, Fry B, Melillo JM, Steudler PA, Moraes JFL, Cerri CC (1996) Forest- and pasture-derived contributions to carbon stocks and microbial respiration of tropical pasture soils. Oecologia 107, 113-119.

O'Brian BL, Stout JD (1978) Movements and turnover of soil organic matter as indicated by carbon isotope measurements. Soil Biology and Biochemistry 10, 309-317.

Popp M, Osmond CB, Summons RE (1982) Pathway of malic acid synthesis in response to ion uptake in wheat and lupin roots: evidence from fixation of [sup.13]C and [sup.14]C. Plant Physiology 69, 1289-1292.

Pressenda LCR, Aravena R, Melfi A J, Telles ECC, Boulet R, Valencia EPE, Tomazello M (1996) The use of carbon isotopes ([sup.13]C, [sup.14]C) in soil to evaluate vegetation changes during the Holocene in central Brazil. Radiocarbon 38, 191-201.

Roscoe R, Buurman P, Velthorst EJ (2000) Disruption of soil aggregates by varied amounts of ultrasonic energy in fractionation of organic matter of a clay latosol: carbon, nitrogen and [delta][sup.13]C distributions in particle-size fractions. European Journal of Soil Science 51,445--454.

Ross SM (1993) Organic matter in tropical soils: current conditions, concerns and prospects for conservation. Progress in Physical Geography 17, 265-305.

Santruckova H, Bird MI, Lloyd J (2000) Microbial processes and carbon isotope fractionation associated with heterotrophic metabolism in tropical and temperate grassland soils. Functional Ecology 14, 108-114.

Schimel DS, Braswell BH, Holland EA, McKeown R, Ojima DS, Painter TH, Parton WJ, Townsend AR (1994) Climatic, edaphic and biotic controls over storage and turnover of carbon in soils. Global Biogeochemical Cycles 8, 279-293.

Schleser GH, Jayasekera R (1985) [delta][sup.13]C variations of leaves in forests as an indication of reassimilated C[O.sub.2] from the soil. Oecologia 65,536-542.

Schmidt MWI, Rumpel C, Kogel-Knabner I (1999) Evaluation of an ultrasonic dispersion procedure to isolate primary organomineral complexes from soils. European Journal of Soil Science 50, 87-94

Schonwitz R, Stichler W, Ziegler H (1986) [delta][sup.13]C values of C[O.sub.2] from soil respiration on sites with crops of C3 and C4 type of photosynthesis. Oecologia 69, 305-308.

Schwartz D, de Foresta H, Mariotti A, Balesdent J, Massimba JP, Girardin C (1996) Present dynamics of the forest-savanna boundary in the Congolese Mayombe: a pedological, botanical and isotopic [sup.13]C and [sup.14]C study. Oecologia 106, 516-524.

Schweizer M, Fear J, Cadish G (1999) Isotopic (13C) fractionation during plant residue decomposition and its implications for soil organic matter studies. Rapid Communications in Mass Spectrometry 13, 1284-1290.

Sofer Z (1980) Preparation of carbon dioxide for stable carbon-isotope analysis of petroleum fractions. Analytical Chemistry 52, 1389-1391.

Stewart GR, Turnbull MH, Schmidt S, Erskine PD (1995) [sup.13]C Natural abundance in plant communities along a rainfall gradient: a biological indicator of water availability. Australian Journal of Plant Physiology 22, 51-55.

Syvertson JP, Smith ML, Lloyd J, Farquhar GD (1997) Net carbon dioxide assimilation, carbon isotope discrimination and water-use efficiency of citrus trees in response to nitrogen status. Journal of the American Society of Horticultural Science 122, 226-232.

Townsend AR, Vitousek PM, Trumbore SE (1995) Soil organic matter dynamics along gradients in temperature and land-use on the island of Hawaii. Ecology 76, 721-733.

Turchenek LW, Oades JM (1979) Fractionation of organo-mineral complexes by sedimentation and density techniques. Geoderma 21, 311-343.

Van der Merwe NJ, Medina E (1989) Photosynthesis and [sup.13]C/[sup.12]C ratios in Amazonian rain forests. Geochimica et Cosmochimica Acta 53, 1091-1094.

Wedin DA, Tieszen LL, Dewey B, Pastor J (1995) Carbon isotope dynamics during grass decomposition and soil organic matter decomposition. Ecology 76, 1383-1392.

Manuscript received 13 December 1999, accepted 5 August 2002

M. Bird (A,E), O. Kracht (B.F), D. Derrien (C), and Y. Zhou (D)

(A) Research School of Earth Sciences and Research School of Biological Sciences, Australian National University, Canberra, ACT 0200, Australia.

(B) Max Planck Institut fur Biogeochemie, Postfach 100164, 07701, Jena, Germany.

(C) Ecole Normale Superieure de Lyon, 69007, Lyon, France.

(D) Research School of Earth Sciences and Research School of Biological Sciences, Australian National University, Canberra, ACT, 0200, Australia.

(E) present address: National Institute of Education, Nanyang Technological University, 1 Nanyang Walk, Singapore 637616. Corresponding author; email: mibird@nie.edu.sg

(F) present address: Thermo Finnegan Mat GMBH, Boschring 12, Egelsbach 63329, Germany.
COPYRIGHT 2003 CSIRO Publishing
No portion of this article can be reproduced without the express written permission from the copyright holder.
Copyright 2003 Gale, Cengage Learning. All rights reserved.

Article Details
Printer friendly Cite/link Email Feedback
Author:Bird, M.; Kracht, O.; Derrien, D.; Zhou, Y.
Publication:Australian Journal of Soil Research
Geographic Code:8AUST
Date:Jan 1, 2003
Words:10211
Previous Article:Chemical characteristics of phosphorus in alkaline soils from southern Australia.
Next Article:Soil organic matter as influenced by straw management practices and inclusion of grass and clover seed crops in cereal rotations.
Topics:


Related Articles
Water retention characteristics of soils with contrasting clay mineral composition in semi-arid tropical regions.
Immobilisation and mineralisation of carbon and nitrogen from dairy farm effluent during laboratory soil incubations.
Organic carbon and soil porosity.
Soil carbon sequestration and density distribution in a Vertosol under different farming practices.
Calibration of the RothC model to turnover of soil carbon under eucalypts and pines.
The ability of Distichlis spicata to grow sustainably within a saline discharge zone while improving the soil chemical and physical properties.
Soil quality assessed by carbon management index in a subtropical Acrisol subjected to tillage systems and irrigation.
Traffic and tillage effects on wheat production on the Loess Plateau of China: 1. Crop yield and SOM.
Nitrogen isotope enrichment factor as an indicator of denitrification potential in top and subsoil in the Apace Valley, Slovenia.
Evidence for biocycling from Ba/Ca, Sr/Ca, and [sup.87]Sr/[sup.86]Sr in soils (Red Brown Earths) from South Australia.

Terms of use | Privacy policy | Copyright © 2020 Farlex, Inc. | Feedback | For webmasters