A study was carried out on soils sampled at 0-10, 10-20, and 20-30 cm depths from both cultivated and uncultivated soils at four different locations (Awgu, Okigwe, Nsukka I, and Nsukka II), to evaluate the potentials of various aggregate size fractions of varying soil textures and depths to sequester carbon under different land uses. A 4 x 2 x 3 factorial experiment was conducted in a completely randomized design (CRD). Factor A was location at four levels, while factor B (land use) had two levels. Factor C (soil depth) comprised of three levels. Results showed that in both land uses, soil texture varied with depth in each location and included clay, loam, clay loam, sandy loam and sandy clay loam. Generally, all the soil properties varied with soil depth across the locations and land uses. Land use significantly  (P = 0.05) affected pH in KCl, Ca2+, Al3+, CEC, 0.50-1.00 mm water stable aggregates (WSA), total soil nitrogen (TSN) in 1.00-2.00 mm WSA, and soil organic carbon (SOC) in 1.00-2.00 mm and < 0.25 mm WSA. Cultivation at 0-30 cm depth significantly reduced SOC in 1.00-2.00 mm WSA by 19.30 %, and TSN in 1.00-2.00 mm WSA by 2.50 %. Land use effects on SOC in WSA at 0-30 cm depth of the various locations followed no consistent trend, except that SOC was higher in cultivated than in uncultivated soils of Nsukka II location. The SOC pool significantly decreased with soil depth. The SOC pool at 0-10 cm, 10-20 cm, and 20-30 cm depths averaged 17.62, 16.40 and 13.05 Mg C ha-1 respectively, in cultivated soils; and 19.59, 17.86 and 12.03 Mg C ha-1 respectively, in uncultivated soils. The SOC pool to the depth of 30 cm differed distinctly amongst the study sites in both land uses; however, cultivation had no significant effect on SOC pool. The effect due to soil texture on SOC pool indicated that C sequestration was significantly greater in clay loam > clay > sandy loam > loam > sandy clay loam. In all, SOC pool was most secluded at 10-20 cm depth, and least at 20-30 cm depth. Whereas SOC pool significantly correlated with dispersion ratio (DR), aggregated silt and clay (ASC), water dispersible clay (WDC), microporosity (Pmi), 0.50-1.00 mm WSA, mean weight diameter (MWD), soil pH, K+, and C/N ratio in cultivated soils; it correlated significantly with ASC, Na+, and CEC in uncultivated soils. Apart from Pmi, whose variability was largely due to the effect of SOC that significantly predicted up to 76 %, SOC significantly accounted between 34 % and 54 % of the variability in MWD, WDC, and WSA classes of > 2.00 mm, 1.00-2.00 mm and 1.00-0.50 mm of the cultivated soils.



            The fundamental basis of carbon (C) sequestration and its effect on global climate change and agriculture have become a major concern in recent years. Emissions of greenhouse gases (water vapour, carbon dioxide (CO2), methane, and nitrous oxide) as a result of human activities continue to alter the atmosphere in ways that are expected to affect change in climate. Anthropogenic activities produce CO2, which is the primary greenhouse gas that contributes to climate change to be released to the atmosphere at rates much faster than the earth’s natural processes can cycle. To help alleviate or possibly reverse the trend, a variety of means of enhancing natural sequestration processes are being explored. Increasing CO2 sink (C sequestration) has been acknowledged and accepted as a major possible mitigation to these effects. This is buttressed by the report of Rice and McVay (2002) indicating that through C sequestration, atmospheric CO2 levels are reduced as soil organic carbon (SOC) levels are increased”. 

            Among the three natural sinks for C (ocean, forest and soil), soils contain more C than is contained in vegetation and the atmosphere combined (Swift, 2001). The SOC pool which forms the largest sink after sedimentary rocks and fossil deposits however is the most vulnerable to disturbance (Schlamadinger and Marland, 2000) especially because of the competition between the various types of land use. Six et al. (2000) reported that tillage operations promote the loss of SOC through macroaggregate disruption and exposure of soil organic matter (SOM) to microbial decomposition. Also, Blum (1997) indicated that the decomposition and alteration (mineralization and metabolization) of organic compounds produces trace gases which can be harmful to the global atmospheric cycle.

            The impact of organic carbon (OC) losses in soils may have a variety of serious environmental consequences. Lal (2004) reported that several depletion of SOC degrades soil quality, reduces biomass productivity, and adversely impacts water quality. Lal et al. (1998) observed that organic matter (OM) losses from soil worldwide contribute to increased atmospheric CO2 concentration. Lugo and Brown (1993) indicated that the net losses of SOC due to land use changes may occur as a result of decreased organic residue inputs and changes in litter composition, and increased rates of soil organic decomposition and soil erosion. The contribution of soil erosion to global C emission has also been recognized by Tans et al. (1990) as equally important to that of deforestation and fossil fuel burning. Lal (1995) estimated that the total SOC displaced by water erosion globally as 57 Pg yr-1 [Pg = Petagram. Where, 1 Pg = 1 Gt (Gigaton) = 1015 g = 1 billion tons]. Houghton et al. (1996) predicted that CO2 emission to the atmosphere would increase from 7.4 Gt C yr-1 in 1997 to approximately 26 Gt C yr-1 by 2010. Furthermore, the annual CO2 flux from the soil to the atmosphere (68 Pg yr-1) is 11.3 times the emissions from fossil fuel combustions (6 Pg yr-1) (Raich and Schlesinger, 1992). However, the Inter-Government Panel on Climate Change (IPCC) recognised three main options for the mitigation of atmospheric CO2 concentrations by the agricultural sector: (i) reduction of agriculture-related emissions, (ii) creation and strengthening of C sinks in the soil, and (iii) production of bio-fuels to replace fossil fuels (Batjes, 1998). Hence, the need to evaluate the role of soil as one of the natural C sinks that secludes organic C as stable humus for enhancing soil fertility and stability of soil microaggregates. Therefore, soil C pool and its dynamics play vital role and the knowledge of their spatial distribution is important for understanding the pedosphere in the global C cycle for the overall management of C. It is with this background that several attempts have been made to access the potential of cropland (Lal et al., 1999; Lal and Bruce, 1999), grazing systems (Follet et al., 2000), and forest ecosystem (Birdsey et al., 1993) to sequester C as possible strategies to curtail the rate of increase of atmospheric concentration of CO2

            Carbon sequestration refers to the removal of C, from the atmosphere through photosynthesis and dissolution, and storage in soil as OM or secondary carbonates (Lal, 2001). Through this process, C storage in soil is enhanced and its loss minimized, thereby reducing the chances of global warming by the reduction of atmospheric concentration of CO2. Recognizing the soil as one of the important potential sinks for C requires understanding of the processes that influence C sequestration. Soil aggregation has been observed as an important process of C sequestration and hence a useful strategy for mitigating increase in concentration of atmospheric CO2 (Shrestha et al., 2007). Igwe et al. (2006) stressed the importance of the study of the role of SOC in restoration of soil fertility and stability of soil microaggregates.

            The impact of C sequestration on greenhouse gases and agricultural sustainability has not been well elucidated at regional, national or global scales. Some available statistics are generally based on extrapolation. Lal (2004) reported that the rates of SOC sequestration in agricultural and restored ecosystems range from 0 to 150 kg C ha-1 yr-1 in dry and warm regions, and 100-1000 kg C ha-1 yr-1 in humid and cool climates. He also estimated the total potential of C sequestration in world soils as 0.4-1.2 Gt C yr-1, all of which were derived from national resource inventory. Improvement in the data base on the concentration of SOC needed to be validated with ground truth measurement/assessment, as the use of reliable data is essential for developing techniques of soil management and identifying policy options needed for promoting appropriate measures. Despite several studies carried out on the quantification of soil sequestered C in different geographical regions of the world (Cruz-Rodriguez, 2004; Denef et al., 2004; Lal et al., 1998; Lal, 2001; Shrestha et al., 2007), there are limited knowledge about SOC pool dynamics in the tropical humid agroecosystem of southeastern Nigeria. Quantification of SOC within aggregate size classes permits evaluation of aggregation under different soil management systems and its contribution to the accumulation and loss of OM (Sotomayor-Ram´ırez et al., 2006). The relevance of this study is to generate reliable information which is essential for developing techniques of soil/land management systems and for recommendation of agricultural practices that promote C sequestration for sustainable agriculture leading to advancement in food security and consequently, mitigate global warming. The hypothesis is that SOC sequestration is a function of soil texture and soil aggregation; and that SOC is similar between soil phases (cultivated and uncultivated) of the same soil series. Therefore, the main objective of the study was to assess the potentials of various aggregate size fractions of varying soil textures and depths to sequester C in cultivated and uncultivated soils. The specific objectives included to;

  • Determine the soil physico-chemical properties of cultivated and uncultivated soils.
  • Quantify SOC and total soil nitrogen (TSN) stocks and assess their distribution across aggregate size fractions as stratified by location, land-use, soil texture and soil depth.
  • Determine the effect of SOC and TSN on soil aggregation and other soil properties.
  • Understand the SOC pool dynamics among different soil textures and depths, and between cultivated and uncultivated soils.



  •      An Overview of Global Carbon

       Understanding the concept of carbon (C) sequestration seems to be the most reasonable point to start a lesson, but C sequestration which is an aspect of C management cannot be absolutely understood and appreciated as one of the identified mitigation options against global warming and climatic change, without a critical survey of global C as documented by most researchers.

         Three main reservoirs regulate the C cycle on earth (IPCC, 1990): the oceans ≈ 39000 x 1015 g (or Pg) of C; the atmosphere (≈ 750 Pg C), and terrestrial systems (≈ 2200 Pg C). The fourth reservoir which is a permanent sink – the geological reservoir, is estimated at 65.5 x 106 Pg (Kempe, 1979). Table 1 shows the principal global C pools comprising of oceanic, geologic, pedological, atmospheric, and biotic. These pools are interconnected with sizeable fluxes among them. For example, the atmospheric pool is increasing at the rate of 3.3 Gt C yr-1. The oceanic pool is absorbing about 92 Gt C yr-1 and emitting 90 Gt yr-1, with a net gain of 2 Gt yr-1. The biotic pool photosynthesizes 120 Gt C plant respiration and the remaining 60 Gt C yr-1 by soil respiration (Lal et al., 2007). Although the soil-vegetation C pool is small compared with that of the oceans, potentially it is much more labile in the short term (Batjes, 1996). On average, the soil contains about 2.5 times more organic carbon (OC) than the vegetation (≈ 650 Pg C) and about twice as much C as is present in the atmosphere (≈ 750 Pg C) (Batjes, 1998). The soil is the largest terrestrial pool of OC, with global estimates ranging from 1115 to 2200 Pg C (Batjes, 1992), 1576 Pg C (Eswaran et al., 1995), 1400 Pg C (Falloon et al., 1998) and 1220 Pg C (Sombroek et al., 1993). Estimates of global soil C content have also been made by several researchers including, Kimble et al. (1990), Buringh (1984), Bohn (1982), Batjes (1998), Lal (2002), Post et al.(1990), Johnson and Kerns (1991).

         World soils or the pedological pool comprises two distinct components: soil organic carbon (SOC) and soil inorganic carbon (SIC) pools estimated at 1576 Gt and 938 Gt respectively, at 1 m depth (Post et al., 1982; Schlesinger, 1995; Eswaran et al., 1993). The SOC pool is concentrated in soils of arctic, boreal, and temperate regions which includes highly active humus and relatively inert charcoal C, while the SIC pool includes elemental C and carbonate minerals, such as calcite, dolomite, and gypsum and those of arid and semiarid climates (Schnitzer, 1991; Stevenson, 1994; Lal et al., 2000; Wagner, 1981; Paul and Clark, 1996). Lal (2004) recognized two types of carbonates in soils: primary or lithogenic carbonates and secondary or pedogenic carbonates. Of the global SOC pool (2500 Gt), which includes about 1550 Gt of SOC and 950 Gt of SIC therefore, the total soil C including both SOC and SIC pools in the active soil layer of 1 m depth constitutes about 23000 Pg (Lal, 2002b) which is 3.3 times the size of the atmospheric pool (760 Gt) and 4.5 times the size of the biotic pool (560 Gt) (Lal, 2004). Table 2 shows the global mass of SOC in the upper 30 cm, 1 m, 2 m, and 3 m of soil. The SOC pool in the top 1 m depth of world soils ranges between 1462 and 1600 Pg, which is nearly three times that in the aboveground biomass and approximately double that in the atmosphere; 32 % (or 506 Pg) of this is contributed by soils in the tropics (Eswaran et al., 1993; Lal et al., 1995; Batjes 1996). According to Batjes (1996) total soil C pools for the entire land area of the world, excluding carbon held in the litter layer and charcoal, amounts to 2157 to 2293 Pg of C in the upper 100 cm. Owning to the problem of making accurate global estimates of C, Eswaran et al. (1993) suggested that employing the coefficient of variation (CV) which is an expression of the variability will aid in understanding and accepting the reliability of most generalization.

         Lal et al. (2007) documented the estimates of global carbon pool up to 1 m depth of the various USDA soil orders. Another study that provided global mean of SOC estimates in soils of the tropics reported values of 8.3 for Ultisols, 9.7 for Oxisols, and 10.4 kg m-2 for Inceptisols among other soil orders (Lal, 2002a). These differences between soil orders in the tropics are mainly in relation to temperature, rainfall, soil texture and land use (Batjes, 2000). Similarly, Kimble et al. (1990) reported SOC density of principal soil orders of the world, where among others the mean SOC density is 9.7 kg m-2 with CV of 42 % for tropical Oxisols and 8.3 kg m-2 with CV of 70 % for tropical Ultisols. About 52 % of this C pool is held in the top 30 cm of the soil profile, the layer most susceptible to land use changes and responsive to management practices (Lal, 2002b).

2.2      Soil Organic Carbon Dynamics 

         Lal (2001) noted that the SOC pool which is a function of soil characteristics and climatic factors is a highly variable and dynamic entity. It is variable over space because its density differs widely among soils and ecoregions; and variable over time because it changes with change in land use and management. Lal (2002b) observed that SOC pool is in a dynamic equilibrium with its environment, with a balance of input and output at a steady state level. Thus, the SOC pool represents a dynamic equilibrium of gains and losses (Fig. 1). Follett (2001) pointed out that temporal variation in SOC level results from the balance between plant biomass input and decomposition and OC losses via leaching, oxidative, and erosional processes. Substrate quality is one of the main factors affecting decomposition and has been linked to the relative abundance of specific compounds such as nitrogen, lignin (Melillo et al., 1982; Tian et al., 1993), and phenolic acids (Martens, 2000). Lal (1999) hypothesized that lack of nutrients, especially N, could explain the low C conversion efficiency. Knops and Tilman (2000) observed

that the rate of carbon accumulation in agricultural abandoned fields was controlled by the rate of nitrogen accumulation, which in turn depended on atmospheric nitrogen deposition and symbiotic nitrogen fixation by legumes. More so, the turnover time of organic matter (OM) increases with depth in the soil, ranging from several years for litter to 15-40 years in the upper 10 cm and over 100 years below a depth of about 25 cm (Harrison et al., 1990; Lobo et al., 1990). In soils of the tropics, particle size fractionation techniques have been used to characterize relationships between SOC and aggregation at the macro and microaggregate scale (Feller et al., 1996). The concept is that soil organic fractions associated with different sized particles differ in structure and function, and therefore play different roles in SOC turnover (Christensen, 1992).       

         The SOC pool consists of “a mixture of plant and animal residues at various stages of decomposition, of substances synthesized microbiologically and/or chemically from the breakdown products, and of the bodies of live microorganisms and soil animals and their decomposing products” (Schnitzer, 1991). The different C pools existing in the soils have been described in terms of the different mean residence times, ranging from years (active fraction), to decade to hundreds of years (passive), to thousands of years (stable) (Carter et al., 2002). The C pools are relative concepts based on the rate of decomposition of particular constituents and are more related to biological function than to particular soil chemical C constituents. For example, the active fraction consists of live microorganisms (microbial biomass), microbial products, and unprotected chemical constituents such as proteins and polysaccharides with a turnover time of a few weeks or months. The slow fractions are more resistant to decomposition due to partial physical and chemical protection with a longer turnover time (Theng et al., 1989). The passive organic constituents include humic substances and other macromolecules that are intrinsically resistant against microbial attack due to chemical recalcitrance, physical protection by adsorption on mineral surfaces, or entrapment within soil aggregates (Gregorich et al., 1997). Lal (2001) pointed out that formation of stable microaggregates in the subsoil takes C out of circulation by encapsulating it (physical and chemical protection from microbial activity) and is thus sequestered. The long-term stabilization of C in temperate and tropical soils is mediated by soil biota (e.g. fungi, bacteria, roots and earthworms), soil structure (e.g. aggregation) and their interactions, and is influenced by agricultural management (Six et al., 2002).

            In most soils, C is organic and constituents approximately 57 % of the soil organic matter (SOM) that includes a wide spectrum of organic compounds, from labile components, such as relatively fresh plants material and microbial biomass, to refractory components such as charcoal, which accumulates slowly over thousands of years (Trumbore, 1993). Piccolo (1996) indicated that the process of turning agricultural soil into sink for OC sequestration would be complete if the stored OM were transformed into stable and recalcitrant humic substances. Accordingly, humified OC, humic acids and humin in particular, represent the most persistent pool for SOC

Fig. 1:  Processes affecting soil organic carbon dynamics.

Arrows pointed upward indicate emissions of Co2 into the atmosphere (Adapted from Lal, 2004).

accumulation with mean residence time of several hundreds of years. Spaccini et al., (2002) is of the view that hydrophobic protection provided by humified matter may substantially reduce decomposition of labile organic compounds in soils, thus they reported that the higher the hydrophobicity of a humic material, the larger the sequestration of OC in soil.  Bayer et al. (2000) reported that in southern Brazil, SOC associated with sand and silt fractions was less humified than that associated with finer-sized fractions. Nevertheless, Oades et al. (1987) and Six et al. (2000) demonstrated that the most humified or oldest fraction is associated with silt particles. Most of the C losses following soil disturbance such as tillage originate from the active and slow pools, which comprise the biologically defined SOM pools described as active (labile), slow (partially labile), and passive (stable) (Jenkinson and Rayner, 1977; Jenkinson, 1990; Duxbury and Nkambule, 1994). Biological separation of SOC empirically separates labile from recalcitrant forms by allowing microbes to mineralise C under controlled conditions with the most labile C mineralised first and recalcitrant C mineralised later.

         Physical fractionation of the soil according to aggregate size has been used to study the portioning of OC in the soil (Buyanovsky et al., 1994). Fractionation using physically based models is possible because OC is protected within and between aggregates (Cambardella and Elliot, 1992; Six et al., 2000; Snyder and V´azquez, 2004). In its simplest case, there is a free light fraction (LF) of labile C between the aggregates and intra-aggregate particulate OM (iPOM) within macroaggregates (Cambardella and Elliot, 1992; Six et al., 1998). The LF may be more related to residue input rates and soil environmental conditions and the iPOM more related to aggregate turnover, which is strongly affected by tillage management (Six et al., 1998). Aggregate hierarchy levels of formation occur in which the intra-macroaggregate POM (particulate organic matter) facilitates the binding of microaggregates into macroaggregates, which in turn affects the variation in the accessibility of soil microorganisms to SOC that leads to pools which differ in stability and dynamics. For example, in relatively undisrupted systems such as no-tilled agricultural and native systems, the greatest C concentration is usually found in the small macroaggregate size class (250–2000 μm), with C in this fraction being most affected by cultivation (Beare et al., 1994; Cambardella and Elliott 1994). Furthermore, a greater proportion of the SOC pool in large microaggregates implies greater C losses to the atmosphere if macroaggregates are broken by soil management practices (Cruz-Rodríguez, 2004). The losses of SOC are in turn associated with losses in the POM fraction and therefore the amount of aggregation and aggregate turnover (Six et al., 1999).