The terrestrial biosphere is dominated by sloping landscapes [Staub and Rosenzweig, 1986] where biogeochemical cycling of essential elements is controlled by interaction of geomorphic, pedogenic, and ecological processes that shape them. Recent studies have highlighted the important role of soil erosion in dynamics of soil organic matter (SOM) [Berhe et al., 2007, 2008; Boix-Fayos et al., 2009; Harden et al., 1999; Smith et al., 2001; Stallard, 1998a; Van Oost et al., 2007]. However, most process level studies of SOM cycling are dominantly located on nonsloping sites that experience minimal soil erosion and deposition and thus they fail to capture the influence of topography on SOM dynamics.
1.1. Erosion-Induced Terrestrial Carbon Sequestration
Annually, water erosion is estimated to move 30–100 Pg soil and 1–5 Pg carbon (C) globally—70 to 90% of which is deposited within the same or adjacent watersheds [McCarty and Ritchie, 2002; Stallard, 1998b; Starr et al., 2001]. It has been shown that erosion and terrestrial deposition can act as a C sink for the atmosphere with an estimated strength of up to 1.5 Pg C yr−1 [Berhe et al., 2007; Harden et al., 1999; Stallard, 1998b; Van Oost et al., 2007]. Soil erosion constitutes a C sink if posterosion watershed C stocks increase due to replacement of eroded C by production of new photosynthate at eroding positions and/or reduced decomposition rate of at least some of the eroded C in depositional landform positions [Berhe et al., 2007].
Studies on the role of soil erosion in terrestrial C sequestration typically focus on quantifying changes in the inventory of C in eroding watersheds over time [Quine and Van Oost, 2007; Smith et al., 2001; Stallard, 1998a; Van Oost et al., 2005, 2007]. Except in a few studies [Berhe et al., 2008; Billings et al., 2010; Harden et al., 1999; Nadeu et al., 2011], stability and stabilization mechanisms of SOM have not been addressed in the context of the erosion-induced terrestrial C sequestration. Hence we lack a good understanding of which mechanisms of SOM stabilization are most important during detachment of aggregates by rain splash, transport of eroded particles downslope, or decomposition postdeposition.
1.2. Mechanisms of Soil Organic Carbon Stabilization in the Context of Erosion and Sedimentation
In this study, we adopt the definition of stability of SOM as its persistence in the soil system [Schmidt et al., 2011]. In eroding watersheds that experience lateral redistribution of topsoil by soil erosion, stability of SOM (as inferred from MRT or 14C, for example) does not necessary indicate time spent in a specific profile, at a specific landform position, rather it is more accurate when used to indicate time spent in the whole watershed.
Three mechanisms had been previously recognized as governing stabilization of SOM: physical isolation of SOM inside aggregates, chemical interaction of OM with the soil matrix, and molecular composition of SOM [Christensen, 1992; Sollins et al., 1996; Stevenson, 1994]. Recently, the importance of molecular composition (also referred to as recalcitrance) as an important factor for SOM persistence has been challenged [Kleber, 2010b; Schmidt et al., 2011]. Here, we provide a brief discussion on how physical isolation, chemical interaction, and molecular structure of SOM are related to dynamics of SOM in eroding watersheds.
1.2.1. Physical Isolation of SOM
In both eroding and depositional positions of a given watershed, SOM can be protected from decomposition by physical isolation of OM inside aggregates of soil particles and/or burial in deep soil layers. Aggregation can render organic compounds physically inaccessible to soil microbes and fauna, and restrict the rate of diffusion of oxygen and enzymes [Kleber, 2010b; Sexstone et al., 1985; Six et al., 2002; Sollins et al., 1996]. In a similar manner, burial of topsoil material in subsoils of depositional positions could slow down decomposition by changing the environmental drivers of SOM decomposition, such as by reducing O2 availability and/or changing soil moisture content [Berhe, 2012].
Redistribution of SOM with soil erosion can both increase and decrease accessibility of SOM to decomposition. Early in the redistribution process, protective macroaggregates are broken down and finer mineral soil particles and light organic particles are suspended (in the case of water erosion), increasing the probability that previously physically protected SOM will be accessed by decomposer organisms and enzymes. Breakdown of aggregates also facilitates oxygenation and hydration of the previously enclosed OM, further enhancing the potential for its decomposition. However, after deposition, burial in depositional soil profiles with high bulk density, low total porosity, and small pore sizes impedes access by decomposers and their enzymes, and hinders the diffusion of oxygen and moisture. Moreover, redistribution can also put OM in a new matrix context, where an organic fragment that was previously surrounded by large sand grains may be deposited among silt and/or clay particles thus allowing the formation of tighter and more efficient physical barriers for decomposition. The net effect of redistribution on OM decomposition will depend on the duration of the exposure and transport phases; depth of burial, and the type of newly created matrix context.
1.2.2. Interaction With Soil Minerals
In aerobic soils, the oldest SOC tends to be associated with soils that have high clay content, high concentration of aluminum (Al) and iron (Fe) oxy(hydr)oxides, large specific surface area (SSA, externally available sorption sites), and high sum of exchangeable cations (as indicated by cation exchange capacity, CEC) [Kahle et al., 2002; Kaiser and Guggenberger, 2001; Masiello et al., 2004; Mayer and Xing, 2001; Mikutta et al., 2006; Oades, 1988]. Thus considerable SOC storage and stabilization is associated with organo-mineral complexation phenomena controlled by soil mineralogy and pedogenesis [Eusterhues et al., 2003]. Decomposition may be retarded by chelation of organic acids with Fe3+ and Al3+ to form metastable intermediate organo-metal complexes [Boudot, 1992; Masiello et al., 2004; Oades, 1995; Rasmussen et al., 2005; Tate, 1992; Torn et al., 1997] or by interactions with poorly crystalline or amorphous minerals such as ferrihydrite. These metastable oxyhydroxides are characterized by high specific surface area, variable charge, high degree of hydration, and are capable of forming OM-mineral bonds through innersphere ligand exchange reactions [Masiello et al., 2004] as well as adsorb anions by a combination of electrostatic attraction and surface complexation [Masiello et al., 2004; Oades, 1988; Torn et al., 1997].
Erosion can bring dissolved OM and/or previously unprotected particulate OM into contact with mineral surfaces and enable sorptive protection. Upon deposition, eroded OM enters a new matrix context where the finer soils particles, including reactive colloids that are preferentially transported by water erosion, are deposited. This chemical mechanism of SOM stabilization should be most effective if erosion removes OM from an area that has relatively higher concentration of unreactive (such as low-charge siloxane) mineral surfaces and deposits it in environments with relatively higher abundance of more reactive minerals (such as hydroxylated Fe oxides). Sorption of OM on surfaces of Fe oxides can lead to formation of innersphere complex bonds that provide increased protection against desorption of OM from mineral surfaces.
1.2.3. Molecular Structure
SOM comprises a collection of simple and macromolecular organic functional groups [Krull et al., 2003; Sollins et al., 1996]. The mean residence times (MRT) of different compounds can vary widely depending on their association with the soil matrix and environmental conditions. By changing the physical, biological, and microclimatic conditions in environments where a given SOM pool or compound class resides, transport and deposition can change its decomposition rate, and some types of SOM might be expected to be more strongly influenced by the change in environment [Berhe et al., 2007; Schmidt et al., 2011; Trumbore, 2009]. For example, let's consider a case where SOM from slopes is laterally transported by soil erosion to an environment that is waterlogged, as is the case in valley floors or depositional positions such as our plain position that is waterlogged for at least for part of the year. Lateral redistribution with soil erosion can lead to transfer of OM from well aerated soil profiles at slopes to depositional environments that are waterlogged, less oxygenated, have lower density of active microbes—where the conditions in the depositional valley floor or plain are less optimal for decomposition. In theory, in the short term this could lead to reduced decomposition of organic compounds that are less soluble, larger, and have high stoichiometric oxygen demand (for decomposition) than others. But, these short-term dynamics should not be confused with long-term resistance of organic functional groups to decomposition. Molecular structure alone does not control long-term decomposition dynamics of SOM [Kleber, 2010a, 2010b; Schmidt et al., 2011].
1.3. Objectives and Hypotheses
The objective of this study was to characterize the relative importance of different mechanisms that may extend the residence time of SOM in eroding versus depositional landform positions. We hypothesized that (1) physical stabilization (aggregate protection) of SOM is more important in depositional landform positions, represented by hollow and plain in this study, compared to the eroding positions; (2) chemical stabilization of SOM (sorptive protection against decomposition by soil minerals) is more important in depositional, compared to eroding landform positions; and (3) there is less change in SOM composition with soil depth in the depositional landform positions, compared to eroding positions. We studied an annual grassland watershed in California with four types of landform positions: summit, slope, hollow, and plain. As we investigate whether SOM is stored effectively in depositional versus eroding positions, we expected the residence time of soil mineral particles to be longer in depositional compared to eroding landform positions. We evaluate the effectiveness of different landform positions for SOM storage by (1) computing the weighted average 14C content of each profile, (2) determining if more C is stored inside aggregates in the depositional versus eroding landform positions, (3) correlating the stock of bulk SOM and fraction of OM in the dense fraction (DF) with Fe and Al oxides, and (4) determining the distribution of SOM functional groups and ratio of alkyl to O-alkyl groups in SOM at the different landform positions.