phenolics (C). Lower inhibition causes higher hydrolase activity (D) and organic matter decomposition (E) releasing CO2 and nutrients (F), which can feed back on microbial activity (G) through pH effects related to CO2 production and nutrition effects due to release of nitrogen, phosphorus, and other nutrients.
Source: Modified from Fenner & Freeman (2011).
Physical protection.
Organic matter can be physically protected from decomposition through chemical associations with mineral surfaces, by being physically inaccessible in soil pores, or as a result of encapsulation by humic materials. The importance of these mechanisms has been well‐illustrated by studies in terrestrial soils and marine sediments, where organic matter associated with mineral particles can be preserved for thousands of years yet is rapidly mineralized once desorbed (Keil et al., 1994; Nelson et al., 1994). For wetlands, more than half of the soil carbon pool could be protected by minerals at low soil carbon concentrations, but the mineral‐protected fraction necessarily drops as soil carbon concentrations increase (Needelman, Emmer, Emmett‐Mattox et al., 2018).
Evidence from soils and sedimentary systems indicates that interactions between organic carbon and mineral particles play a role in carbon preservation (Hedges & Keil, 1995; Hemingway et al., 2019; Mayer, 1994a; Torn et al., 1997). In this respect, aluminosilicate clays are especially important due to ionic surface charges and the high surface area to mass ratio of these smallest sediment particles. Early research suggested that organic carbon uniformly coated aluminosilicate minerals in a monolayer (Keil et al., 1994; Mayer, 1994b), but it now seems that distribution of organic matter is discontinuous across the mineral surface. Some areas of the aluminosilicate minerals are carbon‐free, whereas other regions contain organic carbon that is strongly adsorbed to the mineral plus an outer zone characterized by hydrophobic interactions between sorbed molecules and those in solution (Kleber et al., 2007). Physical sorption appears to be especially important in protecting organic molecules with low C:N ratios (e.g., amino acids, proteins) through electrostatic bonds between negatively charged portions of the clay and positively charged organic side chains (Aufdenkampe et al., 2001). Hydrophobic organic molecules in solution can interact with the hydrophobic end of molecules sorbed directly to the mineral surface, creating a membrane‐like layer that provides a degree of protection to the outermost layer of organic molecules (Kleber et al., 2007). Further, the surface of aluminosilicate minerals contains a multitude of pores of various sizes that (1) increase surface area versus a (theoretical) pore‐free mineral, and (2) exclude microbial and/or enzymatic access to sorbed organic molecules (Jastrow et al., 2007). In sandy subtropical marsh soils, the organic matter found in pores of 6 μm diameter (versus those of ~200 μm) had a greater thermodynamic potential (i.e., higher NOSC), a larger level of chemical complexity, and a higher degree of microbial reactivity (Bailey et al., 2017). Similarly, ~20% of the amino acid nitrogen in an Arctic tussock soil was physically isolated in pores (Darrouzet‐Nardi & Weintraub, 2014). At the nanoscale, even the smallest extracellular enzymes are largely excluded from pores with diameters ≤ 8 nm (Mayer, 1994a; Zimmerman et al., 2004). Typical pore sizes vary with mineralogy (Dalal & Bridge, 1996), but pore diameters are often < 8 nm (Mayer, 1994a), implying that organic carbon preservation in pores is a widespread mechanism.
Organic matter can be chemically stabilized through sorption and coprecipitation with Fe(III) (oxyhydr)oxides (Kaiser & Guggenberger, 2000; Lalonde et al., 2012) or by forming a non‐crystalline floc with Fe2+ (Henneberry et al., 2012). For example, the aeration of porewater from a fen removed 27% of the dissolved organic carbon (DOC) due to coagulation with newly formed Fe(III) hydroxides (Riedel et al., 2013) and salt marsh soils had up to 50% of their soil organic carbon stabilized due to associations with Fe(III) (oxyhydr)oxides (Cui et al., 2014). Protection by iron helps explain why lignin is preferentially preserved in wetland soils. Iron minerals strongly sorb phenolic molecules (Riedel et al., 2013) and inhibit the mineralization of lignin but not that of bulk soil organic matter (Hall et al., 2016). At redox interfaces like the wetland plant rhizosphere, there is dynamic redox cycling of Fe (e.g., Weiss et al., 2004) where the microbial and chemical dissolution of Fe(III) can release sorbed carbon into solution (Chin et al., 1998; Knorr, 2013). However, many wetlands contain solid‐phase Fe(III) as a coating on vascular plant roots, in shallow soils where atmospheric O2 penetration occurs, and as Fe‐rich concretions (R. M. Chambers & Odum, 1990; Duan et al., 1996; Emerson et al., 1999; Mendelssohn et al., 1995). While there is an overall decline in Fe(III) with increasing soil depth (Cutter & Velinsky, 1988; Griffin et al., 1989), oxidized iron can persist under anaerobic conditions over geologically relevant timescales (Haese et al., 1997). We have focused here on the preservation of organic carbon, but wetlands can contain measurable amounts of inorganic carbon in the form of siderite (FeCO3) (Duan et al., 1996; Hansel et al., 2001; T. Wang & Peverly, 1999).
Lastly, proteins and amino acids can become encapsulated in humic acids and protected from hydrolysis. In soils and sediments, humic acid fractions can be hundreds or thousands of years old yet have high concentrations of amide and amino nitrogen, forms of organic matter which are often highly reactive (e.g., Hedges & Keil, 1995; Knicker et al., 1996; Mahieu et al., 2002; Zang et al., 2000). The humic acids may be forming a micelle‐like structure that traps reactive organic molecules within the hydrophobic interior of the structure (Zang et al., 2000), which is consistent with observations that hydrophobic organic contaminants also have a high affinity for humic acids (De Paolis & Kukkonen, 1997). This protective mechanism may be most important at low pH where humic acids form structures with a lower surface–volume ratio (versus a chainlike structure at higher pH), which enhances the ability of the humic materials to physically trap organic matter (Myneni et al., 1999). Given the low pH of many peatlands and their general paucity of mineral matter, the encapsulation of organic matter by humic acids in peat is likely to be more important than interactions with aluminosilicate clays or iron minerals.
pH.
Wetland soils span a wide range of pH values, from bogs and pocosins with pH values of ~4 or less to riparian floodplains and other wetlands where the pH can exceed 7.5 (e.g., Jacob et al., 2013; Richardson, 2003). We focus here on low pH wetlands since that is where pH has the largest inhibitory effect on carbon mineralization. Rates of CH4 production and emission are low in acidic wetlands and increase when pH is experimentally increased (Dunfield et al., 1993; Ye et al., 2012). The suppression of CH4 emissions by low pH occurs through direct inhibitory effects on the hydrogenotrophic and acetoclastic methanogenic pathways as well as interference with the fermentative processes that generate the substrates used by methanogens (Ye et al., 2012). Atmospheric acid deposition also depresses CH4 emission rates, although this effect is mediated by the competitive suppression of methanogenesis by NO3– and/or SO42– rather than a direct pH effect (Gauci et al., 2004; Watson & Nedwell, 1998). Rates of soil carbon mineralization to CO2 are also limited by low pH due to the inhibitory effects of pH on fermentation (Ye et al., 2012), the suppression of phenol oxidase activity (Williams et al., 2000; Xiang et al., 2013), a microbial community characterized by slow‐growing bacteria (Hartman et al., 2008), and/or the encapsulation of reactive organic matter by humic acids (see Physical protection in Section 3.3.2). Experimental increases of soil pH in the lab often lead to higher rates of CO2 production (e.g., Ye et al., 2012) although a multi‐year field experiment found a decrease in soil CO2 production rates in response to increased pH, perhaps because the native microbial community was well adapted to the original low pH environment (Keller et al., 2005).
Temperature.
Temperature affects the efficiency of carbon preservation through several related mechanisms. Firstly, biological processes such as decomposition generally slow down at cooler temperatures, as demonstrated for multiple indices of decomposition including litter decay, soil enzyme activities, biological oxygen demand, CO2 and CH4 production and emissions to the atmosphere, and the hydrologic export