Mechanisms controlling soil carbon turnover and their potential ...

Climatic Change (2007) 80:5–23 DOI 10.1007/s10584-006-9178-3

Mechanisms controlling soil carbon turnover and their potential application for enhancing carbon sequestration Julie D. Jastrow & James E. Amonette & Vanessa L. Bailey

Received: 27 July 2005 / Accepted: 19 May 2006 / Published online: 21 December 2006 # Springer Science + Business Media B.V. 2006

Abstract In addition to increasing plant C inputs, strategies for enhancing soil C sequestration include reducing C turnover and increasing its residence time in soils. Two major mechanisms, (bio)chemical alteration and physicochemical protection, stabilize soil organic C (SOC) and thereby control its turnover. With (bio)chemical alteration, SOC is transformed by biotic and abiotic processes to chemical forms that are more resistant to decomposition and, in some cases, more easily retained by sorption to soil solids. With physicochemical protection, biochemical attack of SOC is inhibited by organomineral interactions at molecular to millimeter scales. Stabilization of otherwise decomposable SOC can occur via sorption to mineral and organic soil surfaces, occlusion within aggregates, and deposition in pores or other locations inaccessible to decomposers and extracellular enzymes. Soil structure is a master integrating variable that both controls and indicates the SOC stabilization status of a soil. One potential option for reducing SOC turnover and enhancing sequestration, is to modify the soil physicochemical environment to favor the activities of fungi. Specific practices that could accomplish this include manipulating the quality of plant C inputs, planting perennial species, minimizing tillage and other disturbances, maintaining a near-neutral soil pH and adequate amounts of exchangeable base cations (particularly calcium), ensuring adequate drainage, and minimizing erosion. In some soils, amendment with micro- and mesoporous sorbents that have a high specific surface – such as fly ash or charcoal – can be beneficial.

All authors contributed equally to this article. J. D. Jastrow (*) Biosciences Division, Argonne National Laboratory, Argonne, IL 60439, USA e-mail: [email protected] J. E. Amonette Chemical Sciences Division, Pacific Northwest National Laboratory, Richland, WA 99354, USA e-mail: [email protected] V. L. Bailey Biological Sciences Division, Pacific Northwest National Laboratory, Richland, WA 99354, USA e-mail: [email protected]

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Climatic Change (2007) 80:5–23

1 Introduction Accumulation of soil organic C (SOC) requires a positive imbalance between inputs to and outputs from soil organic matter stocks. Carbon accrual can be driven by an increase in photosynthetically derived C inputs, a decrease in C losses, or both. Decomposition, leaching, runoff, and erosion can all contribute to losses from any given location, but the latter three processes also have the potential to add to C inputs elsewhere. Carbon sequestration occurs when a positive disequilibrium is sustained over some period of time, with the system eventually achieving a new, higher steady-state. Because storage of SOC in soils is dynamic (i.e., C derived from the atmosphere is ultimately cycled back to the atmosphere), the residence time (τ) of C in the soil is a major determinant of the capacity of a soil to sequester C (Luo et al. 2003). An increase in τ can sequester SOC even without an increase in inputs. Conversely, if τ remains unchanged then increased inputs can only be sequestered for as long as the mechanisms controlling τ remain unsaturated (Six et al. 2002a). Under environmental extremes of moisture, temperature, pH, or nutrient availability, τ is largely controlled by how these factors limit decomposer activity. Inputs may be sequestered seemingly without constraint in a relatively undecomposed and uncomplexed state (e.g., boreal peat deposits) but the accumulated C is vulnerable to release from storage if environmental conditions become more moderate (Freeman et al. 2001). Fire is another significant factor affecting τ, particularly in grassland and savanna soils and peat deposits. Although the primary effect of fire is the rapid loss of aboveground C, the ultimate effects on SOC vary (Johnson and Curtis 2001). The generation of charcoal sequesters C directly and enhances soil fertility (Dai et al. 2005; Glaser et al. 2002). Furthermore, charcoal may play a beneficial role in the stabilization of SOC (Amonette et al. 2003a). Under more tempered environmental conditions, two major mechanisms – (bio)chemical alteration and physicochemical protection – are responsible for stabilizing SOC and thereby controlling τ. With (bio)chemical alteration, SOC is transformed by biotic and abiotic processes to chemical forms that are more resistant to decomposition and, in some cases, more easily retained by sorption to soil solids. The alteration process is generally referred to as humification, and the altered products as humic materials. With physicochemical protection, biochemical attack of SOC is inhibited by organomineral interactions at molecular to millimeter scales. Stabilization of otherwise decomposable SOC can occur via sorption to soil surfaces, complexation with soil minerals, occlusion within aggregates, and deposition in pores inaccessible to decomposers and extracellular enzymes. Both humified and non-humified SOC can be protected, although the greatest impact on τ is likely to arise from protection of humified SOC. Because of the specific nature of the protective sites, either transport of SOC to these sites or reorientation of mineral components around SOC is required for stabilization. The relative importance and potential saturation of these two stabilization mechanisms vary depending on soil type, vegetation, management practices, and environmental conditions. Although increasing C inputs to soil organic matter is an important component of strategies for enhancing SOC sequestration, we focus in this paper on increasing τ. Because turnover is primarily a biological process, we first consider the overall role of the soil biotic community and then briefly review how (bio)chemical alteration and physicochemical protection mechanisms control the stabilization of SOC. Lastly, we discuss how an understanding of the integrated functioning of these mechanisms within the context of soil structure suggest specific manipulations to increase the residence time of C in soils and, thereby, promote SOC sequestration.


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2 Soil biota and carbon stabilization The microbial community in soils is diverse, abundant, and only partly understood (Gans et al. 2005); it is often described as a “black box” (Tiedje et al. 1999). Some researchers estimate that as much as 90–99% of the soil microbial community is unculturable by current technologies (Hill et al. 2000). Therefore, we are limited in our ability to confidently and absolutely ascribe soil processes to specific microorganisms. For the purpose of discussing biogeochemical processes, it is easier to consider the community in terms of the greater biological domains, (e.g., fungi, bacteria, actinomycetes) or in terms of biochemical functions and specific processes (e.g., enzyme activities). In general, two of the dominant groups of soil microorganisms, fungi and bacteria, have substantially different biochemistries with consequent effects on SOC stabilization. A clear example, but by no means the exclusive cause of different sequestration potentials, is the difference in C utilization efficiency (CUE). Although some researchers have reported that CUE may not differ between fungi and bacteria (Payne 1970), others have found bacterial CUE to be lower than fungal CUE (Adu and Oades 1978; Claus et al. 1999). Organisms with a lower CUE respire a greater proportion of metabolized C as CO2. Thus, in metabolizing the same amount of C, bacteria with a lower CUE would contribute less to pools of newly stabilized SOC than fungi. However, as with most relationships detected in soils, these observations are highly dependent upon the soils studied, and generalizations must be interpreted broadly. Fungi and bacteria also differ in the nature of the extracellular enzymes they produce. Fungi produce large amounts of phenoloxidases, laccases, and peroxidases that help attack lignitic materials and promote condensation reactions. Bacteria, on the other hand, are more likely to produce lipases and cellulases needed to attack nonlignitic materials. The monomers derived from degradation of lignitic materials are perhaps the most common constituents of humic materials and, thus an abundance of fungal enzymes should tend to favor humification. A primary contribution of fungi to C stabilization derives from the very nature of their metabolism and biomass formation. Fungal cell walls are constructed of complex molecules such as melanin and chitin, in contrast to bacterial membranes (which are dominantly phospholipids). After cell death, phospholipids are rapidly metabolized by other bacteria, and overall these bacterial products are more vulnerable to grazers (Frey et al. 2001). Melanin and chitin residues are more recalcitrant and tend to persist in soils (Guggenberger et al. 1999; Holland and Coleman 1987). For example, when microbes were grown on leaf litter, the amount of C stored by fungi was 26 times greater than the C stored by bacteria (Suberkropp and Weyers 1996). The fungi discussed so far are largely saprotrophic in nature, i.e., they acquire C through degradation of complex organic compounds. A second subset of fungi, mycorrhizal fungi, also play a distinct and unique role in C sequestration. Although, mycorrhizal fungi are not directly involved in the decomposition of soil organic matter; they are obligate symbionts that coexist with most terrestrial plants. In this relationship, the plant gains nutrients gathered from the greater soil volume explored by mycorrhizal hyphae, and the fungus obtains photosynthate-C directly from the plant. Thus, the primary contribution of mycorrhizal fungi to C sequestration is as an additional source of C inputs to soil organic matter. This occurs, indirectly, through the generally positive effect of the mycorrhizal symbiosis on plant growth, particularly in low phosphorus soils. However, with C standing stocks of 54–900 kg C ha−1 and cell walls composed largely of chitin, the turnover of mycorrhizal fungi also can be a significant direct source of C inputs that are

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relatively resistant to decomposition (Zhu and Miller 2003). Another form of C inputs produced by mycorrhizal fungi is a glycoprotein called glomalin that appears to be somewhat resistant to degradation, perhaps due to interactions with soil minerals (Steinberg and Rillig 2003). In addition, mycorrhizal fungi contribute indirectly to SOC stabilization via their role in the formation and stabilization of soil aggregates (Jastrow et al. 1998; Miller and Jastrow 1990; Tisdall 1996; Tisdall and Oades 1982). Mycorrhizal hyphae together with fine roots create a “sticky-string bag” that enmeshes and entangles soil particles, helping to stabilize macroaggregates (Miller and Jastrow 2000; Oades and Waters 1991). The deposition of glomalin may also contribute to aggregate stabilization (Rillig 2004), although this hypothesis is still subject to scrutiny (Franzluebbers et al. 2000). Lastly, mycorrhizal fungi are hypothesized to further enhance C sequestration by translocating nutrients from the bulk soil to the host plant, thereby competing with freeliving decomposer microorganisms that would otherwise mineralize SOC to CO2 (Rillig and Allen 1999). Our consideration of the potential significance of fungi to C sequestration is not to minimize the role played by bacteria in these processes. On the contrary, bacteria have been extensively studied, and most of what is currently known about SOC dynamics and storage is derived from studies centered on bacteria (e.g., Hu et al. 1999; Marilley et al. 1999; Pelz et al. 1998; Schimel et al. 1999). We highlight the role of fungi in this paper as an emerging area of research and one that merits deeper study. Several recent reports have called attention to the distinct roles played by fungi in the C cycle. For example, fungal activity was correlated with soil C content (Bailey et al. 2002); shifts in mycorrhizal community structure (Treseder and Allen 2000) and increased grazing of fungi (Allen et al. 2005) were found in systems exposed to elevated atmospheric CO2 concentrations; and increases in fungal biomass were observed immediately following land abandonment and in association with undisturbed, high-C heathlands (van der Wal et al. 2006). In addition to fungi and bacteria, soil invertebrates play important roles in the fragmentation, comminution, and decomposition of plant litter and – through grazing and complex food web interactions – exert significant controls on microbial community structure, biomass, and activities (Beare et al. 1995; Mikola et al. 2002). Further, significant transformations of SOC are performed by microbes living within the guts or excrement of soil invertebrates (Hättenschwiler et al. 2005; Martin and Marinissen 1993; Seastedt 1984).

3 (Bio)chemical alteration In most soils, C stabilization necessarily involves alteration of organic matter to chemical forms that are more recalcitrant to microbial attack, more likely to adsorb to soil solids, or both. The alteration typically occurs in two stages. First, plant-derived residues are fragmented into particulate organic matter and then decomposed into smaller molecules. Similarly, the C assimilated by decomposers and used for growth is also decomposed when these organisms turn over. The decomposition stage is followed by condensation and polymerization reactions that create new larger molecules from the small molecules released during decomposition. Both stages of alteration are largely biologically driven, although there is also significant evidence for abiotic condensation and polymerization reactions at mineral surfaces. Consequently, the nature of the soil biotic community and the mineral phases present in the soil are key factors in the alteration process.


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3.1 Decomposition After physical fragmentation to particulate organic matter, the first stage of (bio)chemical alteration involves the decomposition of C inputs (chiefly cellulose and lignin biopolymers present in plant litter) into smaller molecules. This process is mediated principally by microorganisms and their extracellular enzymes. Cellulose is readily degraded by both fungi and bacteria. Lignin, however, is more recalcitrant, and its complete degradation is restricted to a select group of fungi that produce the extracellular lignin peroxidases (Conesa et al. 2002; Schubert 1965; ten Have and Teunissen 2001). As yet, no bacteria have been identified that are able to initiate degradation of lignin, and current knowledge suggests that they can only mineralize the partly transformed intermediates produced by the fungi. For a given cohort of C inputs, selective degradation by fungi and bacteria will decompose the labile portions first and lead to a progressive increase in the average recalcitrance of the remaining SOC with time. Decomposition of the more recalcitrant C inputs relies heavily on the activity of extracellular enzymes, which have a limited lifetime in soils (on the order of days; e.g., Amonette et al. 2003a, b; Dick and Tabatabai 1992; Shen et al. 2002). These enzymes typically sorb to soil solids, with varying effects on their ability to catalyze decomposition. Although the effect of sorption is enzyme- and mineral-specific, in many instances the enzyme activity decreases initially as a result of sorption, but longevity and total activity over the life of the enzyme increase (Boyd and Mortland 1990; Shen et al. 2002). Enzyme sorption to minerals, however, also limits access to potential substrates (particularly the large insoluble biopolymers) and restricts decomposition to regions near the source of the enzyme (George et al. 2005). Sorption of the enzyme to the substrate is the key process. Schimel and Weintraub (2003) suggested that a limited number of sorption sites exist on a substrate surface; hence, once these sites are saturated by enzymes, additional enzymes must diffuse farther from the source and return less product to the microorganism that synthesized them. This negative feedback loop (simulated by a reverse Michaelis–Menton relationship) regulates the production of extracellular enzymes by microorganisms. One model of extracellular enzyme production by bacteria suggests that the typical extracellular enzyme foraging distance is on the order of 10–50 μm (Vetter et al. 1998). At longer distances, the benefit to the organism does not exceed the cost of producing the enzyme. The limited zone of extracellular enzymatic activity is expanded, particularly in undisturbed soils, by fungi whose hyphal growth habit allows them to form “bridges” between the soil and surface litters (Beare et al. 1992; Holland and Coleman 1987). In relatively undisturbed systems, these hyphal networks can become a large component of the soil fabric (Haynes and Beare 1997). Because the fungal biomass is made up of complex polymers such as chitin and melanin (Guggenberger et al. 1999; Holland and Coleman 1987) that are relatively resistant to degradation, its senesced residues contribute to soil organic matter (West et al. 1987). In contrast, bacteria exhibit limited mobility in soils and thus seem to be restricted to using substrates located within their immediate proximity (Wolfaardt et al. 1994). Moreover, bacterial biomass is comprised of labile energy-rich molecules such as phospholipids and amino acids (Guggenberger et al. 1999). Thus senesced bacterial residues do not accumulate or contribute significantly to SOC in their unaltered state (West et al. 1987). In addition, wetting and drying cycles may have an effect on microbial community structure and, consequently, decomposition rates. Schimel and co-workers have shown significant long-term decreases in respiration for a grassland soil and for birch leaf litter

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subjected to repeated wetting and drying cycles (Clein and Schimel 1994; Fierer and Schimel 2002). The relationship is by no means robust, however. Increases in respiration were seen in a soil formed under perennial oak vegetation (Fierer and Schimel 2002), whereas long periods (e.g., six months) between cycles showed no effect (Haider and Martin 1981).

3.2 Condensation/polymerization In the second stage of (bio)chemical alteration, new compounds form from some of the molecules released during the decomposition stage. These compounds are condensates of a variety of different “monomers” and, as a consequence, do not have well-defined composition or structure. Indeed, there is increasing evidence to suggest that the condensates may consist of dynamic clusters of chemically altered and unaltered compounds held together loosely by hydrophobic interactions and hydrogen bonding (Sutton and Sposito 2005). In general, the condensates have higher aromatic character and oxygen content than the average biomass present in soils (Hayes and Malcolm 2001). When compared to plant litter, the condensates have higher levels of carboxylic and fatty-acid C and lower levels of polysaccharide C (Chefez et al. 2002; Zech et al. 1997). Their most important attribute, however, is that they are more recalcitrant to decomposition, either as a result of their intrinsic biochemical properties or enhanced sorption affinities. As shown in Fig. 1, consensus has emerged among researchers that the chemically altered fraction of the condensates are formed by the reactions of amino compounds (acids and sugars) with quinones or reducing sugars to form melanin-type compounds (Flaig 1975; Haider et al. 1975; Hedges 1988; Kononova 1961; Maillard 1916; Martin and Haider 1971; Martin et al. 1975; Stevenson 1994; Tan 2003; Waksman 1932). The amino compounds and reducing sugars are readily available from the lysing of microorganisms, whereas the quinones are the result of oxidation of polyphenols derived from lignin and other plant materials as well as from microorganisms. The condensation of reducing sugars with amino compounds proceeds spontaneously but slowly at typical soil temperatures, and is greatly enhanced by repeated wetting and drying cycles. Many investigators, however, consider the quinone/amino-compound condensation reaction to be the dominant humification pathway in soils (Stevenson 1994). The ratedetermining step for this reaction is believed to be the oxidation of the polyphenol to form the quinone. The reaction is faster at high pH, and at the near-neutral pH of most soils assistance from a catalyst is needed for the reaction to proceed at a measurable rate. Biological catalysts include polyphenol oxidase, peroxidase, and laccase enzymes produced by fungi, which mediate the electron transfer from the polyphenol to molecular oxygen (Sjoblad and Bollag 1981, Tate 1992). Other substances in soils either catalyze the reaction – e.g., amorphous silica, charcoal, iron-sorbed smectite (Amonette et al. 2003a, b, 2004; Booth et al. 2004) – or serve as oxidants – e.g., oxides and hydroxides of manganese and iron, smectites (Amonette et al. 2000; Naidja et al. 1998; Shindo and Huang 1984; Wang and Huang 2005). Recent evidence suggests the overall reaction rate increases synergistically when both biological and inorganic catalysts are present (Amonette et al. 2000, 2003a, b). Ultimately, the availability of molecular oxygen determines whether the polyphenol oxidation reaction occurs, as this species also oxidizes the metal oxides and phyllosilicates. Indeed, insufficient oxygen is likely to stop the quinone/amino-compound condensation reaction from occurring in soils, just as it prevents the enzymatic decomposition of peat bogs (Freeman et al. 2001). Too much oxygen, however, and the monomers and newly formed condensates will continue to be altered until their C is fully oxidized to CO2. In


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Fig. 1 Major pathways for humification showing C inputs and the types of condensation reactions that are likely to occur. Pathways are (a) sugar-amine, (b) polyphenol-quinone, (c) lignin-quinone, and (d) modifiedlignin (after Stevenson 1994, p. 189).

contrast to the oxidation step, the subsequent condensation reaction with the amine group occurs spontaneously, although many factors (primarily pH) can affect its rate. The mineral phases in soils can have an important impact on (bio)chemical alteration, principally as catalysts of condensation and as oxidants. The dominant minerals in soils on a mass basis are typically quartz and the feldspars, but these minerals have very low specific surfaces (ca. 0.1 m2 g−1) and correspondingly low impact on chemical alterations of soil organic matter. Minerals with high specific surfaces (>10 m2 g−1) and high chemical impact on alterations include the phyllosilicate clays, allophanes, and the oxides and hydroxides of manganese and iron. The phyllosilicate clays can be grouped according to whether they are swelling (e.g., smectites, vermiculites) or non-swelling (e.g., kaolinite, illite), with swelling clays offering internal surfaces as high as several hundred square meters per gram. Although some studies have shown correlations between clay content (or clay plus silt) and the amount of soil organic matter (e.g., Six et al. 2002a, b), the surface reactivity and specific surface of soil minerals appear to be better predictors of SOC (Baldock and Skjemstad 2000).

4 Physicochemical protection Protection of (bio)chemically altered soil organic matter from further microbial decomposition or oxidation by molecular oxygen and extracellular enzymes is essential to significantly lengthen the residence time of C in soils. For protection of new C to occur, some change in the arrangement between this C and other soil particles must occur. Rearrangement can occur in a number of ways but the end state generally involves chemical or physical sorption of the new C to an existing surface coupled with some sort of physical

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barrier to prevent further access by agents that would decompose or oxidize the C. If not already at a surface, the new C may diffuse or advect to a surface where it can be adsorbed. Alternatively, the soil particles may change position to become reoriented in association with SOC as a result of advection, the mechanical actions of plant and fungal growth, bioturbation by earthworms and other soil fauna, or changes in hydration status.

4.1 Chemical protection Relatively labile organic compounds (e.g., microbial byproducts) as well as humified C can be protected by strong chemical associations with soil minerals. Sorption occurs via a variety of organomineral associations, such as polyvalent cation bridges, hydrogenbonding, van der Waals forces, and interactions with hydrous oxides and aluminosilicates. Drying also appears to play an important role in increasing the stability of organomineral associations by enhancing hydrogen bonds in addition to polymerizing and denaturing the organics. Thus, many sorption and complexation interactions are possible depending on various factors, including the chemical characteristics of the soil organic matter, the reactivity and specific surface of soil minerals, base-cation status, presence of iron and aluminum oxides, pH, and redox conditions (for overviews see Baldock and Skjemstad 2000; Blanco-Canqui and Lal 2004; Oades 1984; Sollins et al. 1996). Clay mineralogy plays an important role in the chemical protection of SOC (Baldock and Skjemstad 2000; Dalal and Bridge 1996; Sollins et al. 1996). In soils dominated by 2:1 clay minerals, such as smectite and illite, complexation is believed to occur via the formation of multivalent cation bridges between negatively charged organic groups and negatively charged clay platelets. In contrast, for soils dominated by 1:1 clays (kaolinite) and iron- and aluminum-oxides, surface-charge densities are generally lower and offer fewer sites for complexation with organics. Thus, the availability of multivalent cations – particularly calcium, iron, aluminum, and manganese – is an important factor in the chemical protection of SOC. In addition to their role as bridging agents, high saturation of clays with multivalent cations helps to keep organomineral complexes more flocculated and condensed, thereby reducing the efficiency of attack by microbes and enzymes (Baldock and Skjemstad 2000). This mechanism may be particularly important in soils with low-charge clays. Variations in specific surface are important to chemical protection as they determine potential sites for interactions (Baldock and Skjemstad 2000; Kaiser and Guggenberger 2003; Sollins et al. 1996). Specific surface, which can range from ca. 0.1 to several hundred square meters per gram, varies with clay mineralogy and also decreases with increasing particle size. Thus, sandy or silty soils with small amounts of high-specific-surface clays can likely sorb more C than a similarly textured soil with low-specific-surface clays. Sorption does not appear to occur uniformly across mineral surfaces, but rather has been shown to exist in patches (Baldock and Skjemstad 2000; Kaiser and Guggenberger 2003). The most reactive sites may be located at the mouths of micropores between two domains of a mineral, where multiple organomineral attachments can occur. Sorption at such sites is hypothesized to contribute to observed variations in turnover rates for mineral-associated SOC (Kaiser and Guggenberger 2003). The first molecules sorbed at these sites are the best stabilized, but after these pore openings are filled, additional organic matter loading is less strongly sorbed.


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4.2 Physical protection Substantially greater protection is obtained when decomposer access to chemically protected substrates is physically impeded or when soil structural controls on gas exchange and moisture conditions inhibit decomposer activity (Elliot and Coleman 1988; Elliot et al. 1980; Six et al. 2002a; Sollins et al. 1996; Young and Ritz 2000). Physical protection mechanisms are intimately tied to the processes of aggregate turnover and stabilization at multiple scales. Aggregates are formed and degraded as a result of physical processes including wetting and drying, freezing and thawing, the growth of plant roots and associated mycorrhizal fungi (including physical entanglement and localized drying due to water uptake), bioturbation (i.e., the activities of earthworms, termites, ants, and other soil fauna), and tillage or other physical disturbance (Allison 1968; Dexter 1988; Kay 1990; Oades 1993; Six et al. 2004). In many soils, the deposition and transformation of organic matter plays a major role in aggregate stabilization, and strong feedbacks exist between aggregate turnover and SOC dynamics (Dalal and Bridge 1996; Feller and Beare 1997; Jastrow and Miller 1998; Oades 1993; Six et al. 2004, 2002b; Tisdall 1996). In soils where organic matter is a major aggregate-binding agent, plant growth and decomposition of organic inputs (especially roots) lead to the development of a hierarchical aggregate structure (Fig. 2), where the mechanisms of aggregate formation and stabilization and their relative importance change with spatial scale (Jastrow et al. 1998; Oades and Waters 1991; Tisdall and Oades 1982). Primary particles and extremely stable fine-silt-sized aggregates (250 μm diameter) by labile organic materials and by fine roots, fungal hyphae, bacteria, and algae. By its very nature, aggregate hierarchy creates a parallel hierarchy of pores between and within aggregates of varying sizes that control gas exchange, water movement, and soil food web structure and dynamics (Elliot and Coleman 1988; Young and Ritz 2000). Thus, the nature and effectiveness of various aggregate-binding agents depend on their physical dimensions relative to the dimensions of the pores (i.e., planes of weakness) being bridged, and smaller aggregated units are generally more stable and less susceptible to disturbance (Dexter 1988; Jastrow et al. 1998; Kay 1990; Tisdall and Oades 1982). Current evidence indicates that microaggregates are formed inside macroaggregates (Fig. 3), particularly as root-derived particulate organic matter is decomposed (Angers et al. 1997; Gale et al. 2000; Golchin et al. 1994; Six et al. 2000). Similarly, silt-sized aggregates may be stabilized (Fig. 3) by the deposition of microbial byproducts and condensation of humates to form complexes with mineral particles and decomposing organic debris inside microaggregates (Golchin et al. 1994). Disturbance and other factors that increase the rate of macroaggregate turnover can, thereby, decrease the formation and stabilization of microaggregates (Six et al. 2000). Several lines of evidence indicate that protection of SOC by microaggregates is greater than the protection afforded by macroaggregates. The amount of C mineralized from crushed or dispersed aggregates in laboratory incubation studies is often greater for microaggregates than for macroaggregates (Bossuyt et al. 2002; Gregorich et al. 1989). In tracer studies, labeled C preferentially accumulated in microaggregates compared to other locations within the soil (Angers et al. 1997; Besnard et al. 1996; Denef et al. 2001; Gale et al.

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Fig. 2 Conceptual diagram of aggregate hierarchy illustrating microaggregates inside a macroaggregate (modified from Jastrow and Miller 1998, p. 209).

2000). Furthermore, for soils exhibiting aggregate hierarchy, estimates of the mean residence time of C in macroaggregates are shorter than for C in microaggregates (Six and Jastrow 2002). Although the amount of long-term protection provided by macroaggregates appears to be minimal, the rate of macroaggregate turnover plays an important role in the stabilization of SOC. Some level of macroaggregate turnover may be necessary to incorporate and protect fresh C inputs from rapid mineralization (Plante and McGill 2002). Yet, if macroaggregate turnover is too rapid, microaggregate formation and stabilization can be inhibited, leading to a reduction in the amount of microaggregate-protected SOC – especially microaggregate-occluded particulate organic matter (Six et al. 2004). In addition to the physical barriers created by the encrustation and occlusion of soil organic matter with soil mineral components, the size distribution and tortuosity of pores (which is related to aggregate size distribution and hierarchy) also plays an important role in the physical protection of SOC. Pore size, and in particular pore neck size, controls the distribution, movement and activity of decomposers and soil food web dynamics (Elliot and Coleman 1988; Elliot et al. 1980; Young and Ritz 2000). Labile C substrates filling pores smaller than 1 μm or located inside larger pores with necks less than 1 μm are not accessible to most soil microorganisms. Similarly, the diffusion of extracellular enzymes into micron-sized pores can be limited and is completely excluded from nano-scale pores (Zimmerman et al. 2004). Pore size distributions and, hence, enzyme accessibility at this scale differ with clay mineralogy. For example, pore size in clay tactoids of Ca-saturated illite is ∼10 nm compared to ∼100 nm in kaolinite (Dalal and Bridge 1996). Decomposer activity is also limited by localized oxygen and water availability. Steep declines in oxygen concentrations have been measured within small distances from aggregate surfaces, and interactions between water films and small pore necks can lead to


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Fig. 3 Conceptual model of microaggregate formation and turnover (based on Gale et al. 2000, Golchin et al. 1994, Six et al. 2000). Illustrated stages are (a) fragmented and decomposing roots and litter become encrusted with mineral particles forming microaggregates; (b) decomposition continues but at a slower rate due to physical protection; microaggregate is stabilized; (c) slowed decomposition and intimate contact with soil minerals in stable microaggregate enables organic matter to be humified or chemically protected by association with mineral fraction; (d) organic binding agents decompose sufficiently for aggregate to be destabilized; mineral fraction enriched with new organomineral associations becomes available for incorporation into new microaggregates.

anaerobic patches within largely aerated aggregates (Sexstone et al. 1985; Young and Ritz 2000). Recent work by Strong et al. (2004) demonstrated initial rates of decomposition were most rapid in soils with large volumes of intermediate-sized pores of about 15–60 μm. In particular, decomposition was enhanced near the gas–water interface, most likely because these locations optimize (1) organism motility; (2) the diffusion of nutrients, toxins, and enzymes; and (3) oxygen supply. In larger pores (60–300 μm) oxygen was abundant, but decomposition was slower – probably because motility, diffusion, and direct contact between residues and mineral surfaces were all reduced. However, the greatest protection and slowest rates of decomposition were found in soils that had large volumes of pores with neck diameters < 4 μm. Thus, the interactions of the microbial community with the complexity of pore space and surfaces that compose their physical habitat leads to protection of SOC via what has been termed “partial refuge” (Ekschmitt et al. 2005). In essence, this means that not all degradable SOC is within range of decomposing organisms all the time because of (1) the limited zones of microbial habitats – estimated by Nannipieri et al. (2003) to be