Stability and storage of soil organic carbon in a heavy-textured Karst soil from south-eastern Australia.
Soil organic matter (SOM), which is made up of ~50% soil organic carbon (SOC), is a vital natural resource that promotes soil health and fertility by fulfilling numerous soil functions such as nutrient cycling (Paul 1984), enhancing water retention (Emerson 1995), enhancing cation exchange capacity (Chan et al. 1992) and promoting aggregation (Emerson 1959; Tisdall and Oades 1982). Additionally, SOC is recognised as an important pool of the global C cycle and is seen as an easy target for short-mid-term C sequestration (Smith et al. 2007), providing the possibility of mitigating atmospheric C[O.sub.2] levels. Soil C sequestration, i.e. an increase in the storage of SOC resulting from the transfer of atmospheric C[O.sub.2] to SOC, is facilitated by plant primary production and, thus, provides not only the possibility of reducing greenhouse gas concentrations but also securing global food security (Lai 2010).
A key concern regarding the efficacy of soil C sequestration is the longevity of C storage in soil. SOM is a complex mixture of substances, with retention times ranging from days to millennia (Trumbore 1997). These differing retention times result from different processes controlling SOC dynamics. Processes leading to an increase of SOC retention time in soil are known as stabilisation mechanisms. Storage and stability of SOC depend on environmental and site factors, which determine whether and how SOC is stabilised (and destabilised) in soil (Schmidt et al. 2011). These processes include sorption, physical separation from biodegraders, and freezing and thawing in soils. Mechanistically, they inhibit either microbial activity and/or SOC bioavailability, hence reducing SOC degradation and therefore increasing SOC retention times.
Highly effective stabilisation mechanisms can result in SOC retention in soils for centuries to millennia, and these mechanisms are viewed as the target for soil C-sequestration schemes. Two such long-term SOC stabilisation mechanisms are mineral association and aggregation.
Mineral association denotes the physico-chemical bonding of SOM onto the surfaces of soil fine minerals (the silt and clay fractions). The SOC can be physically (i.e. via sorption) or chemically (i.e. via predominantly ionic bonds) bound to the mineral matrix, resulting in a reduction of bioavailability. The SOM stabilised in this way is called mineral-associated OM (MAOM) (Oades 1988; Baldock and Skjemstad 2000; Sollins et al. 2006).
Aggregation leads to a physical encapsulation of SOM in the interior of soil aggregates (occlusion) (Golchin et al. 1994). This occlusion reduces the physical accessibility of SOC to microorganisms and leads to a change in the chemical environment within the aggregate, as diffusion limits the exchange of nutrients and degradation byproducts (e.g. [O.sub.2], C[O.sub.2]) between the aggregate interior and bulk soil matrix.
Despite the imperative to target long-term SOC stabilisation mechanisms for soil C sequestration, it has been reported that the largest storage pools of SOC may not be associated with the most stable SOC (Jastrow and Miller 1996; Hoblcy et al. 2013). This means that the success of SOC sequestration schemes is constrained by both the retention time of the sequestered C (SOC stability), and the capacity of the soil to store SOC (SOC storage), specifically with regard to stabilised C. As SOC stability and storage are environment- and site-specific, accurate identification of the key factors determining SOC storage and stability is required to ensure the success of such schemes.
We hypothesise that the factors influencing SOC storage are not necessarily the same as those that enhance SOC stability, resulting in a decoupling of SOC storage and stability. In other words, the mechanisms associated with the greatest SOC content are not necessarily those associated with the oldest SOC.
In this study, we test this hypothesis by investigating the storage and stability of SOC in a clayey soil from Wombeyan Karst Conservation Reserve, in the Southern Highlands of New South Wales, Australia. The site is under native vegetation and without history of agriculture, and so is a suitable baseline reference point for SOC storage and stability, e.g. for comparison with agricultural or anthropogenic soils of similar texture and mineralogy, or for comparison with other native soils of different texture, mineralogy or pedogenesis in the region (e.g. Hobley et al. 2013). Furthermore, as a soil formed on Karst and located above a cave system, the site provides future possibilities for investigating cave formations as paleoenvironmental archives of soil processes (Frisia and Borsato 2010). Specifically, wc use a combination of particle-size fractionation, elemental analysis and radiocarbon dating to investigate the amount and mean retention time of SOC in fine soil particles and aggregates in this heavy-textured soil.
Site, soil and sampling description
Wombeyan Karst Conservation Reserve (34[degrees]19'S, 149[degrees]59'E) is within the Southern Highlands of New South Wales, in southeastern Australia. The elevation of the area ranges from 540 to 920 m above sea level. Geologically, it includes predominantly Silurian limestone and marble outcrops intruded by Devonian acidic porphyry (dacites) and granites. The mean maximum annual temperature of the region is 18.2[degrees]C and the mean minimum annual temperature is 6.1[degrees]C. The mean annual precipitation is 801mm without notable seasonality (data from Taralga Post Office; BOM 2012). Mean annual precipitation is less than the mean annual evapotranspiration (Jennings et al. 1982; BOM 2012), so that the region experiences water-stress, which limits primary production and influences SOM dynamics. The region is vegetated with a mixture of dry sclerophyll eucalypt woodland and open forest, with patches of denser forest and rainforest. At the sampling site, the vegetation was sparse, with grasses and bushes predominating.
The investigated soil formed upon Palaeozoic marble bedrock in the Reserve. The sampling site was at an elevation of ~670m. The marble at the sampling site was highly weathered and the soil developed upon it had visible iron and manganese oxide deposits below 80 cm depth. The soil sampled is a Red Dermosol (Isbell 2002; traditionally known as Terra Rossa, Chromic Luvisol; IUSS Working Group WRB 2006) with a clay texture (Table 1). Two soil profiles, ~10m apart, were dug by hand to a depth of 100-110 cm, where bedrock hindered further excavation. Integrated depths were sampled at 0-30, 30-60 and 60-100 cm, which correspond with sampling depths of a site adjacent to this study and therefore enable future comparison of the sites (Hobley et al. 2013).
Textural analysis was performed on the fine fraction (<2mm) via wet-sieving after dispersion in sodium polymetaphosphatc (0.2%) (AS 12893.6.1 2009). The fraction <60 [micro]m was collected and analysed using a Sedigraph III Particle-size Analyzer (Micromeritics Instrument Corp., Norcross, GA, USA). Immediately before analysis, samples in 0.2% sodium polymetaphosphate were dispersed via ultrasound (26 W output). X-Ray diffraction analysis of the samples was performed with an X'Pert PRO MPD (Phillips, now PANalytical B.V., Almelo, Netherlands). Quartz ([alpha]-Si[O.sub.2]), clay minerals (montmorillonite, nacrite, kaolinite) and iron and manganese oxides were identifiable phases throughout the profile (Table 1). The soil is neutral to slightly acidic (pH in deionised water, soil: water ratio 1 : 5; Table 1).
Oven-dried (40[degrees]C) soil samples were disaggregated manually and the coarse fraction (>2 mm) was removed and weighed. The fine fraction (<2 mm, referred to hereafter as 'bulk sample') was dry-sieved using a set of nested sieves on a mechanical sieve (ELE International) for 15 min. Dry-sieving was chosen to avoid loss of dissolved OC from samples resulting from drying and rewetting (Kaiser et al. 2001; Fierer and Schimel 2003), and dry sieving is preferable to wet sieving for preservation of structure and soluble SOC components (Sarkhot et al. 2007). Three particle-size fractions (PSFs) were separated (AS 12893.6.1, Standards Australia 2009): 2000-200 [micro]m (including particulate OM, coarse sand and macroaggregates, referred to hereafter as the 'macro' fraction); 200-60 [micro]m (medium sand and microaggregates, referred to hereafter as the 'micro' fraction); and <60 [micro]m (fine sand, silt and clay, referred to hereafter as the MAOM fraction) (Hobley et al. 2013). The macro fraction was the dominant fraction by mass, and the MAOM fraction the smallest for all depths.
The relative degree of aggregation (RDA, Table 2) was estimated for each fraction according to the definition of (Hobley et al. 2013):
RDA = 1 + mass [proportion.sub.fraction] - mass [proportion.sub.primary particles] (1)
where RDA is the relative degree of aggregation of each PSF; mass [proportiort.sub.fraction] is the mass of a PSF divided by the mass of the bulk sample from which the sample was fractionated; and mass [proportion.sub.primary particles] is the mass of primary particles obtained in textural analysis using the same sieves as for the particle-size fractionation divided by the mass of the bulk sample. The RDA indicates the net proportion of the fraction that consists of aggregates. An RDA of 1 indicates no net aggregation in a fraction, i.e. the amount of aggregates (of smaller particles) in the fraction is equal to the amount of particles of the fraction that have formed aggregates and are found within in a coarser fraction. An RDA >1 indicates aggregates of smaller particles are accumulated in the fraction, and an RDA <1 indicates particles from this fraction have formed aggregates which are found in other fractions.
The mass concentration of SOC was determined via elemental analysis on oven-dried samples (40[degrees]C) in a CNS2000 analyser (LECO Corporation, St. Joseph, MI, USA) at Southern Cross University, Queensland, and on a Vario MICRO cube EA (Elementar Analysensysteme GmbH, Hanau, Germany) at the Australian Nuclear Science and Technology Organisation (ANSTO), Lucas Heights, according to the method of Rayment and Higginson (1992). Concentration of SOC is reported as mass percentage (%, g [g.sup.-1]) of 105[degrees]C dry sample.
Samples were tested for the presence of carbonates using a gas chromatograph isotope ratio mass spectrometer (GC1RMS, GV 2003; GV Instruments, Manchester, UK) at the University of Newcastle. An aliquot of ashed sample was weighed into an exetainer and flushed with He; 85% [H.sub.3]P[O.sub.4] was injected into the exetainer to cover the sample and the exetainer was heated for 2h at 70[degrees]C to allow the acid to react with carbonates. The exetainer was cooled to room temperature (20[degrees]C) before measuring C[O.sub.2] in the headspace. No C[O.sub.2] peaks were detected in the samples, indicating the absence of inorganic C in the soil.
The C density of the soil was calculated based upon the SOC content and bulk density of the fine fraction (<2 mm) in which SOC was measured:
Carbon density = [SOC.sub.fine fraction] x [[rho].sub.fine fraction] (2)
where Carbon density is the amount of SOC per volume unit of soil at a given depth expressed in kg [m.sup.-3], [SOC.sub.fine fraction] is the SOC content (g[g.sup.-1]) of the fine fraction at a given depth; [[rho].sub.fine fraction] is the bulk density (kg [m.sup.-3]) of the fine fraction at each depth. Bulk density was estimated for the fine fraction according to:
[[rho].sub.fine fraction] = [mass.sub.fine fraction]/Vcorer--vrock (3)
where [V.sub.corer] is the volume of the corer used to calculate the bulk density of the sample; [V.sub.rock] is the volume of the rock in the bulk density sample (measured using a Helium Pycnometer; Micromeritics Instrument Corp., Norcross, GA, USA); [mass.sub.fine fraction] is the measured mass of the sample <2 mm in the bulk density corer.
Enrichment factors of SOC
To assess the relationship between particle size and SOC retention, the SOC enrichment factor, Ef was calculated for each fraction and depth according to the relationship reported in (Hobley et al. 2013):
Ef = %SOC in fraction / %SOC in bulk soil (4)
Radiocarbon dating of one soil profile was done via accelerator mass spectrometry (AMS) at ANSTO according to Fink et al. (2004). Aliquots of the PSFs were heated in 6n HC1 overnight to ensure absence of carbonate. In order to minimise loss of SOC, the samples were not rinsed before drying. Organic C was then oxidised from the samples at 900[degrees]C before graphitisation over an Fe catalyst at 600[degrees]C (Hua et al. 2001) to form a target for AMS measurement on the STAR accelerator (2MV Tandetron; High Voltage Engineering, Amersfoort, Netherlands). Radiocarbon concentration is reported as per cent modem carbon (pMC) and conventional radiocarbon age (years before present, years BP). The radiocarbon ages have not been calibrated. Samples with a radiocarbon content >10 pMC are considered modern (after 1950) as they contain radiocarbon derived from the radiocarbon produced by above-ground nuclear testing in the 1950s and early 1960s. The radiocarbon concentration of bulk samples was not determined independently but was instead calculated from the radiocarbon concentration of the PSFs and their content in the bulk sample for each depth.
Errors and statistical testing
For SOC content, errors were estimated as the standard deviation of SOC content of bulk samples at a given depth. This was calculated from SOC content of (i) different profiles, (ii) analysis in different laboratories, (iii) SOC content of bulk samples predicted from SOC content in PSFs and mass proportion of PSFs. The standard deviation of SOC content at a given depth was less than or equal to the analytical error. Errors are therefore depicted as the analytical error. For radiocarbon dating, only one profile was analysed, so that the error reported is analytical error. For the enrichment factors (Ef see Enrichment factors of SOC), errors shown are the standard deviation of estimates (n = 3) calculated with the different SOC measurements. Errors reported for the relative degree of aggregation (RDA, see Particle-size fractionation) have been propagated from the standard deviation of results from the fractionation procedure (n = 3) and the textural analysis (n = 2). Statistical differences between SOC concentrations were tested using the Student's t-test.
SOC concentration and storage
The SOC concentration decreased significantly (P < 0.01) from the topsoil to bedrock in the bulk samples and all PSFs. Figure 1 shows the mean SOC concentration of the samples at each depth. The SOC concentrations of PSFs and bulk samples ranged from 1.4% in the topsoil to 0.2% at a depth of 60-100 cm. The macro fraction consistently had lower SOC concentrations than the bulk samples; the micro fraction and MAOM fractions had higher SOC concentrations than the bulk samples.
The average C density was 13.3 [+ or -]2.1, 6.1 [+ or -]0.9 and 3.4 [+ or -] 1.1 kg [m.sup.-3], at 0-30, 30-60 and 60-100 cm, respectively, yielding SOC stocks of 7.2 [+ or -] 0.8 kg [m.sup.-2] for the soil profile, of which 4.0 [+ or -]0.6 kg [m.sup.-2] (56%) was in the top 30 cm. This is close to the mean for similar soils internationally, which lies at 3.9 kg [m.sup.-2] in the top 0-30 cm and 8.0 kg [m.sup.-2] in the top 100 cm (Batjes 1996), indicating that this soil is typical of its class.
Enrichment factors of SOC
The SOC enrichment factor, Ef was calculated for each fraction and depth (Fig. 2). The Ef was <1 in the macro fraction, and decreased slightly from the topsoil to bedrock. In the micro and MAOM fractions Ef was >1, similar at all depths and showed no trend.
Radiocarbon ages of SOC
Radiocarbon ages increased with depth for all PSFs (Table 3). The oldest radiocarbon ages were consistently in the macro fraction; the micro and MAOM fractions were younger and had similar ages. Radiocarbon ages calculated for the bulk samples were slightly younger than, but most similar to, the macro fraction at all depths.
SOC storage mechanisms
Both the absolute SOC concentrations and Ef indicate that SOC in the soil is preferentially associated with the finest particles (Figs 1 and 2). We attribute this to the larger surface area to volume ratio of fine particles, which results in an increase in physical sorption for SOC relative to coarse particles (Christensen 2001). The macro fraction consisted of nearly 50% aggregates, the remainder being primary particles (>200 [micro]m, i.e. sand; Table 2), so that SOC storage in the macro fraction will be less than in the finer fractions, as it has less clay particles. Furthermore, reactive mineral phases (layered clay minerals and metal oxides) promote SOC storage (Oades 1988; Kaiser and Guggenberger 2000; Bachmann et al. 2008), increasing the storage potential of the soil. This effect will be most pronounced in the fine fraction as it has the smallest particles and most clay minerals.
The relative degree of aggregation (Table 2) shows that the macro fraction is highly aggregated and that the particles in this fraction appear to be primarily derived from the MAOM fraction, as the increase in RDA of the macro fraction is nearly identical to the decrease in RDA of the MAOM fraction. Although the micro fraction may contribute to the aggregation of the macro fraction, the content of the micro fraction of the soil is not large enough to account for the increase in RDA observable in the macro fraction. The micro fraction does not show a large amount of aggregation. Macroaggregate formation in soils with kaolinitic clay minerals occurs independent of biological processes and does not lead to C sequestration in the newly formed aggregates (Denef and Six 2005), so that the lower C content of the macro fraction is likely due to the mineralogy of the soil driving aggregation, as opposed to the theory of aggregate formation promoted by OM (Tisdall and Oades 1982).
We therefore conclude that as a result of the aggregation in the soil being driven by mineralogy, not SOM composition, aggregation does not lead to an increase in SOC content of the aggregates and is therefore a less effective storage mechanism for OM than mineral association in this heavy-textured soil.
Despite the enhanced C storage capacity of the finest particles, the oldest radiocarbon ages are consistently found in the macro fraction (Table 3). This implies that aggregation, though less effective at enhancing C storage, is more effective at increasing C stability than mineral association. Occlusion has been found to increase SOC stability (Golchin et al. 1995; Skjemstad et al. 2008; Swanston et al. 2005), and has been found to be more effective at stabilising SOC than mineral association (Rasmussen et al. 2005), although microaggregates have been reported as more stable than macroaggregates (Jastrow and Miller 1996).
However, the high stability of microaggregates is believed to be due to differences in SOC composition between macro and microaggregates (Tisdall and Oades 1982). In this soil, aggregation does not appear to be linked with SOC retention (see above discussion SOC storage mechanisms), but is instead attributed to soil mineralogy, so that the aggregate hierarchy proposed by Tisdall and Oades (1982) does not apply. As a result, microaggregates are no more stable than macroaggregates. Indeed, the radiocarbon data indicate that the SOC associated with macroaggregates is more stable than that associated with microaggregates.
During aggregate formation in a soil where aggregation is driven by mineralogy, not OM, the OC that is already sorbed to the aggregating particles will be occluded within the formed aggregates, enhancing SOC stability without a concurrent increase in SOC content of the aggregates. This is reflected in the greater radiocarbon ages of the macro fraction compared with the other fractions, despite its lower SOC content.
An alternative hypothesis (to greater SOC stability within aggregates than outside the aggregate) is that the lower stability of the C in the finer particles may be the result of a preferential sorption of younger C to the finest fractions during translocation of dissolved OM (Eusterhues et al. 2007). This would lead to a shift of the radiocarbon age of the finer fractions in subsoil horizons to younger C. However, despite preferential sorption of OM to clay minerals (Golchin et al. 1995) and Fe and Al phases associated with the fine minerals in subsoils (Kaiser et al. 2002), sorption leads to stabilisation of the OM and a shift to older SOC in these fractions, so that this explanation seems less plausible. We therefore believe that occlusion within aggregates is responsible for the greater radiocarbon ages of the SOC in the macro fraction than in the other fractions.
Reported ages of SOC stabilised by occlusion within aggregates are in the order of decades (Marin-Spiotta et al. 2008) to centuries (Jastrow and Miller 1996; Rasmussen et al. 2005) at the soil surface, and centuries to millennia in subsurface horizons (Rasmussen et al. 2005; Rethemeyer et al. 2005). Our radiocarbon data agree well with these timescales, indicating stabilisation of SOC within aggregates in the order of centuries in the topsoil and millennia in the subsoil (Table 3). Compared with the finer fractions, this implies enhanced stabilisation due to aggregation in the order of centuries throughout the soil profile.
Despite the lower stability of mineral-associated SOC than aggregate-associated SOC, the MAOM fraction had radiocarbon ages of centuries to millennia in the subsoil, indicating a much greater stability of mineral-associated SOC below the soil surface than at the surface. Radiocarbon content is commonly depleted with increasing soil depth (e.g. Trumbore et al. 1989; Tom et al. 1997; Rumpel et al. 2004; Eusterhues et al. 2007; Hobley et al. 2013) and the depletion in radiocarbon with depth is often attributed to enhanced stability of subsoil C.
In particular, increased stability of subsoil SOC is frequently attributed to sorption to the mineral matrix (Kaiser et al. 2002; Rumpel et al. 2002; Mikutta et al. 2006), and reported retention times of mineral-associated OM are often in the order of centuries to millennia (e.g. Mikutta et al. 2006; Eusterhues et al. 2007; Chabbi et al. 2009). Our data agree with these timescales of centurial to millennial SOC retention. As such, sorption to the mineral matrix in the subsoil can be viewed as an effective stabilisation mechanism, though not as effective as aggregation in our soil.
We investigated the efficacy of both mineral association and aggregation at storing and stabilising SOC in the topsoil and subsoil of a native Terra Rossa soil in south-eastern Australia using the techniques of soil particle-size fractionation, elemental analysis and radiocarbon analysis. Our results indicate that throughout the soil profile the greatest storage of SOC occurs in the finest particles, which is attributed to their greater surface-area to volume ratio, as well as the mineralogy of the soil. However, the oldest radiocarbon ages were not measured in the finest particle-size fractions, but in the coarsest fraction, which was highly aggregated. We therefore believe that occlusion within the aggregates leads to an enhanced stability of SOC in this soil. This result is valid for both the topsoil (0-30 cm) and subsoil (30-100 cm). Despite being younger than the aggregates, the radiocarbon ages of the mineral-associated OM in the subsoil were in the order of centuries to millennia, so that mineral association can also be viewed as a long-term C stabilisation mechanism in the subsoil. Our results show that enhanced SOC storage is not necessary linked with enhanced SOC stability. This has implications for soil C sequestration, namely whether it is possible to increase SOC storage significantly while at the same time targeting highly stable SOC storage mechanisms.
Received 11 October 2013, accepted 3 February 2014, published online 9 May 2014
We thank L. Henderson for provision of the site map, Eckhard Ferber and Andrea Borsato for help in the field, and the friendly staff at ANSTO for laboratory assistance. We are grateful to the Australian Institute of Nuclear Science and Engineering, the NSW Department of Environment, Climate Change and Water, and the University of Newcastle for funding this research. Eleanor Hobley was supported by an Australian Postgraduate Award, and Garry R. Willgoose by an ARC Australian Professorial Fellowship.
Bachmann J, Guggenberger G, Baumgartl T, Ellerbrock RH, Urbanek E, Goebel M, Kaiser K, Horn R, Fischer WR (2008) Physical carbon-sequestration mechanisms under special consideration of soil wettability. Journal of Plant Nutrition and Soil Science 171, 14-26. doi: 10.1002/jpln.200700054
Baldock JA, Skjemstad JO (2000) Role of the soil matrix and minerals in protecting natural organic materials against biological attack. Organic Geochemistry 31, 697-710. doi: 10.1016/SO146-6380(00)00049-8 Batjes NH (1996) Total carbon and nitrogen in the soils of the world. European Journal of Soil Science 47, 151 163. doi: 10.1111/j.13652389.1996.tb01386.x
BOM (2012) "Climate statistics for Taralga Post Office and Goulbum Airport.' (Australian Government Bureau of Meteorology)
Chabbi A, Kogel-Knabner 1, Rumpel C (2009) Stabilised carbon in subsoil horizons is located in spatially distinct parts of the soil profile. Soil Biology & Biochemistry 41.256-261. doi: 10.1016/j.soilbio.2008.10.033
Chan K, Roberts W, Heenan D (1992) Organic carbon and associated soil properties of a red earth after 10 years of rotation under different stubble and tillage practices. Australian Journal of Soil Research 30, 71-83. doi: 10.1071/SR9920071
Christensen BT (2001) Physical fractionation of soil and structural and functional complexity in organic matter turnover. European Journal of Soil Science 52, 345-353. doi: 10.1046/j.1365-2389.2001.00417.x
Denef K, Six J (2005) Clay mineralogy determines the importance of biological versus abiotic processes for macroaggregate formation and stabilization. European Journal of Soil Science 56, 469 479. doi: 10.1111/j. 1365-2389.2004.00682.x
Emerson WW (1959) Stability of soil crumbs. Nature 183,538. doi: 10.1038/183538a0
Emerson WW (1995) Water-retention, organic-C and soil texture. Australian Journal of Soil Research 33, 241--251. doi:10.1071/SR9950241
Eusterhues K, Rumpel C, Kogel-Knabner I (2007) Composition and radiocarbon age of HF-resistant soil organic matter in a Podzol and a Cambisol. Organic Geochemistry 38, 1356-1372. doi: 10.1016/j.orggeochem.2007.04.001
Fierer N, Schimel JP (2003) A proposed mechanism for the pulse in carbon dioxide production commonly observed following the rapid rewetting of a dry soil. Soil Science Society of America Journal 67, 798-805. doi: 10.2136/sssaj2003.0798
Fink D, Hotchkis M, Hua Q, Jacobsen G, Smith AM, Zoppi U, Child D, Mifsud C, van der Gaast H, Williams A, Williams M (2004) The ANTARES AMS facility at ANSTO. Nuclear Instruments & Methods in Physics Research. Section B, Beam Interactions with Materials and Atoms 223-224, 109-115. doi:10.1016/j.nimb.2004.04.025
Frisia S, Borsato A (2010) Karst. In "Carbonates in continental settings. Vol. 61.' (Eds AM Alonso-Zarza, LH Tanner) pp. 269-318. (Elsevier: The Netherlands)
Golchin A, Oades JM, Skjemstad JO, Clarke P (1994) Soil structure and carbon cycling. Australian Journal of Soil Research 32, 1043-1068. doi: 10.1071/SR9941043
Golchin A, Oades JM, Skjemstad JO, Clarke P (1995) Structural and dynamic properties of soil organic-matter as reflected by [sup.13]C natural-abundance, pyrolysis mass-spectrometry and solid-state [sup.13]C NMR-spectroscopy in density fractions of an oxisol under forest and pasture. Australian Journal of Soil Research 33, 59-76. doi: 10.1071/SR9950059
Hobley E, Willgoose GR, Frisia S, Jacobsen G (2013) Environmental and site factors controlling the vertical distribution and radiocarbon ages of organic carbon in a sandy soil. Biology and Fertility of Soils 49, 1015-1[0.sub.2]6. doi: 10.1007/s00374-013-0800-z
Hua Q, Jacobsen G, Zoppi U, Lawson E, Williams A, McGann M (2001) Progress in radiocarbon target preparation at the ANTARES AMS Centre. Radiocarbon 43, 275-282.
Isbell RF (2002) 'The Australian Soil Classification.' Revised edn (CS1RO Publishing: Melbourne)
IUSS Working Group WRB (2006) 'World reference base for soil resources 2006.' 2nd edn. World Soil Resources Reports No. 103. (FAO: Rome)
Jastrow JD, Miller RM (1996) Carbon dynamics of aggregate-associated organic matter estimated by carbon-13 natural abundance. Soil Science Society of America Journal 60, 801-807. doi: 10.2136/sssaj1996.03615995006000030017x
Jennings JN, James JM, Montgomery NR (1982) The development of the landscape. In 'Wombeyan Caves'. (Eds HJ Dyson, R Ellis, JM James) (The Sydney Speleological Society: Sydney)
Kaiser K, Guggenberger G (2000) The role of DOM sorption to mineral surfaces in the preservation of organic matter in soils. Organic Geochemistry 31, 711-725. doi:10.1016/S0146-6380(00)00046-2
Kaiser K, Kaupenjohann M, Zech W (2001) Sorption of dissolved organic carbon in soils: effects of soil sample storage, soil-to-solution ratio, and temperature. Geoderma 99, 317-328. doi: 10.1016/S0016-7061(00)00077-X
Kaiser K, Eusterhues K, Rumpel C, Guggenberger G, Kogel-Knabner I (2002) Stabilization of organic matter by soil minerals--investigations of density and particle-size fractions from two acid forest soils. Journal of Plant Nutrition and Soil Science 165, 451-459. doi: 10.1002/15222624(200208) 165:4<451:: AID-JPLN451>3.0.CO;2-B
Lai R (2010) Managing soils and ecosystems for mitigating anthropogenic carbon emissions and advancing global food security. Bioscience 60, 708-721. doi: 10.1525/bio.2010.60.9.8
Marin-Spiotta E, Swanston CW, Tom MS, Silver WL, Burton SD (2008) Chemical and mineral control of soil carbon turnover in abandoned tropical pastures. Geoderma 143, 49-62. doi:10.1016/j.geoderma.2007. 10.001
Mikutta R, Kleber M, Tom M, Jahn R (2006) Stabilization of soil organic matter: Association with minerals or chemical recalcitrance? Biogeochemistry 77, 25-56. doi: 10.1007/s 10533-005-0712-6
Oades JM (1988) The retention of organic matter in soils. Biogeochemistry 5, 35-70. doi: 10.1007/BF02180317
Paul E (1984) Dynamics of organic matter in soils. Plant and Soil 76, 275-285. doi: 10.1007/BF02205586
Rasmussen C, Tom MS, Southard RJ (2005) Mineral assemblage and aggregates control carbon dynamics in a California conifer forest. Soil Science Society of America Journal 69, 1711-1721. doi:10.2136/sssaj2005.0040
Rayment GE, Higginson FR (1992) 'Australian laboratory handbook of soil and water chemical methods.' (Inkata Press: Melbourne)
Rethemeyer J, Kramer C, Gleixner G, John B, Yamashita T, Flessa H, Andersen N, Nadeau M-J, Grootes PM (2005) Transformation of organic matter in agricultural soils: radiocarbon concentration versus soil depth. Geoderma 128, 94-105. doi:10.l016/j.geoderma.2004.12.017
Rumpel C, Kogel-Knabner I, Bruhn F (2002) Vertical distribution, age, and chemical composition of organic carbon in two forest soils of different pedogenesis. Organic Geochemistry 33, 1131-1142. doi: 10.1016/S0146-6380(2)00088-8
Rumpel C, Seraphin A, Goebel M-O, Wiesenberg G, Gonzales-Vila F, Bachmann J, Schwark L, Michaelis W, Mariotti S, Kogel-Knabner I (2004) Alkyl C and hydrophobicity in B and C horizons of an acid forest soil. Journal of Plant Nutrition and Soil Science 167, 685-692. doi: 10.1002/jpln.200421484
Sarkhot DV, Comcrford NB, Jokela EJ, Reeves 1JB (2007) Effects of forest management intensity on carbon and nitrogen content in different soil size fractions of a North Florida Spodosol. Plant and Soil 294, 291-303. doi: 10.1007/s 11104-007-9255-z
Schmidt M, Tom M, Abiven S, Dittmar T, Guggenberger G, Janssens I, Kleber M, Kogel-Knabner I, Lehmann J, Manning D, Nannipieri P, Rasse D, Weiner S, Trumbore S (2011) Persistence of soil organic matter as an ecosystem property. Nature 478, 49-56. doi: 10.1038/nature10386
Skjemstad JO, Kmll ES, Swift RS, Szarvas S (2008) Mechanisms of protection of soil organic matter under pasture following clearing of rainforest on an [O.sub.x]isol. Geoderma 143, 231-242. doi:10.1016/j.geoderma.2007.11.006
Smith P, Martino D, Cai Z, Gwary D, Janzen H, Kumar P, McCarl B, Ogle S, O'Mara F, Rice C, Scholes B, Sirotenko O (Eds) (2007) Agriculture. In 'Climate Change 2007: Mitigation. Contribution of Working Group III to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change'. (Cambridge University Press: Cambridge, UK/New York)
Sollins P, Swanston C, Kleber M, Filley T, Kramer M, Crow S, Caldwell BA, Lajtha K, Bowden R (2006) Organic C and N stabilization in a forest soil: Evidence from sequential density fractionation. Soil Biology & Biochemistry 38, 3313-3324. doi: 10.1016/j.soilbio.2006.04.014
Standards Australia (2009) AS12220.127.116.11: Soil classification tests--Determination of the particle size distribution of a soil--Standard method of analysis by sieving. In 'Methods of testing soils for engineering purposes'. (Standards Australia: Sydney)
Swanston CW, Tom MS, Hanson PJ, Southon JR, Garten CT, Hanlon EM, Ganio L (2005) Initial characterization of processes of soil carbon stabilization using forest stand-level radiocarbon enrichment. Geoderma 128, 52-62. doi:10.1016/j.geodcrma.2004.12.015
Tisdall JM, Oades JM (1982) Organic matter and water-stable aggregates in soils. Journal of Soil Science 33, 141-163. doi: 10.1111/j.13652389.1982.tb01755.x
Tom MS, Trumbore SE, Chadwick OA, Vitousek PM, Hendricks DM (1997) Mineral control of soil organic carbon storage and turnover. Nature 389, 170-173. doi: 10.1038/38260
Trumbore S (1997) Potential responses of soil organic carbon to global environmental change. Proceedings of the National Academy of Sciences of the United States of America 94, 8284-8291. doi: 10.1073/pnas.94.16.8284
Trumbore SE, Vogel JS, Southon JR (1989) AMS [sup.14]C measurements of fractionated soil organic matter; an approach to deciphering the soil carbon cycle. Radiocarbon 31, 644-654.
Eleanor Hobley (A,B,D), Garry R. Willgoose (A), Silvia Frisia (B), and Geraldine Jacobsen (c)
(A) School of Engineering, The University of Newcastle, Callaghan, NSW 2308, Australia.
(B) School of Environmental and Life Sciences, The University of Newcastle, Callaghan, NSW 2308, Australia.
(c) Institute for Environmental Research, Australian Nuclear Science and Technology Organisation, Lucas Heights, NSW 2234, Australia.
Corresponding author. Email: email@example.com
Table 1. General soil parameters Rock content is reported as mass percentage of whole sample. Sand, silt and clay content are given as mass percentage of the fine soil (<2 mm). Phases identified with XRD: Q, quartz [alpha]-Si[O.sub.2]; N, nacrite [Al.sub.2][Si.sub.2][O.sub.5][(0H).sub.4]; [Fe.sup.III], ferric oxide [alpha]-[Fe.sub.2][O.sub.3]; P, periclase MgO; Ne, neotocite (Mn,Mg,Fe)Si[O.sub.3] x [H.sub.2]0; K, kaolinite [Al.sub.2] [Si.sub.2][O.sub.5][(OH).sub.4]; D, dickite [Al.sub.2][Si.sub.2] [O.sub.5][(OH).sub.4](HCON[H.sub.2]); A, anatase Ti[0.sub.2]; M, montmorillonite [Na.sub.0.3][(Al,Mg).sub.2][Si.sub.4][O.sub.10] [(OH).sub.2] x 8[H.sub.2]O; [Mn.sup.III], manganese(III) oxide [Mn.sub.2][O.sub.3]; I, illite (K,[H.sub.3]O)[Al.sub.2][Si.sub.3]Al [O.sub.10][(OH).sub.2]; CVM, chlorite-vermiculite-montmorillonite [Na.sub.0.5][Al.sub.6][(Si,Al).sub.8][O.sub.20][(OH).sub.10] x [H.sub.2]O Depth Rock content Sand Silt (cm) (>2 mm) (2000-20 [micro]m) (20-2 [micro]m) 0-30 14 [+ or -] 7 45 [+ or -] 1 14 [+ or -] 1 30-60 14 [+ or -] 4 28 [+ or -] 3 10 [+ or -] 1 60-100 25 [+ or -] 19 32 [+ or -] 1 13 [+ or -] 1 Depth Clay pH EC (cm) (<2 [micro]m) ([micro]S [cm.sup.-1]) 0-30 41 [+ or -] 1 6.5 [+ or -] 0.7 28 [+ or -] 7 30-60 62 [+ or -] 1 5.9 [+ or -] 0.1 28 [+ or -] 10 60-100 55 [+ or -] 1 5.7 [+ or -] 0.3 49 [+ or -] 38 Depth XRD phases (cm) 0-30 Q, N, [Fe.sup.III], P, Ne, K, D, A 30-60 Q, M, N, K, [Fe.sup.III], [Mn.sup.III], I 60-100 Q, N, [Fe.sup.III], CVM, K, [Mn.sup.III] Table 2. Mass content and relative degree of aggregation (RDA) of particle-size fractions at varying depths Macro, 2000-200 [micro]m; micro, 200-60 [micro]m; MOAM, <60 [micro]m Depth Macro (cm) Content (%) RDA 0-30 77.7 [+ or -] 2.1 1.47 [+ or -] 0.04 30-60 82.4 [+ or -] 6.8 1.63 [+ or -] 0.13 60-100 81.7 [+ or -] 3.9 1.61 [+ or -] 0.08 Depth Micro (cm) Content (%) RDA 0-30 15.8 [+ or -] 1.6 1.06 [+ or -] 0.13 30-60 12.0 [+ or -] 4.3 1.06 [+ or -] 0.40 60-100 13.1 [+ or -] 2.6 1.05 [+ or -] 0.21 Depth MAOM (cm) Content (%) RDA 0-30 6.2 [+ or -] 0.6 0.47 [+ or -] 0.05 30-60 5.5 [+ or -] 2.6 0.31 [+ or -] 0.15 60-100 5.1 [+ or -] 1.2 0.34 [+ or -] 0.08 Table 3. Radiocarbon concentrations and ages of particle-size fractions and bulk samples at different depths pMC, per cent modern carbon; years BP, years before present. Macro, 2000-200 [micro]m; micro, 200-60 [micro]m; MOAM, <60 [micro]m. Bulk sample radiocarbon content was calculated from PSF data Depth Macro Micro (cm) pMC (years BP) pMC (years BP) 0-30 95.6 [+ or -] 0.4 100.6 [+ or -] 0.3 (365 [+ or -] 35) (modern) 30-60 85.8 [+ or -] 0.4 92.1 [+ or -] 0.3 (1235 [+ or -] 40) (665 [+ or -] 30) 60-100 77.3 [+ or -] 0.4 80.5 [+ or -] 0.4 (2070 [+ or -] 40) (1740 [+ or -] 40) Depth MAOM Bulk (cm) pMC (years BP) pMC (years BP) 0-30 101.2 [+ or -] 0.3 96.8 [+ or -] 0.6 (modern) (265 [+ or -] 50) 30-60 91.9 [+ or -] 0.3 87.4 [+ or -] 0.6 (675 [+ or -] 30) (1085 [+ or -] 60) 60-100 83.6 [+ or -] 0.3 77.9 [+ or -] 0.6 (1435 [+ or -] 35) (2010 [+ or -] 60)
|Printer friendly Cite/link Email Feedback|
|Author:||Hobley, Eleanor; Willgoose, Garry R.; Frisia, Silvia; Jacobsen, Geraldine|
|Date:||Aug 1, 2014|
|Previous Article:||Potential soil organic carbon stock and its uncertainty under various cropping systems in Australian cropland.|
|Next Article:||Manganese oxidation and reduction in soils: effects of temperature, water potential, pH and their interactions.|