6.3.1: Soils and the carbon cycle

Soil C stocks:

Carbon (C) storage is an important ecosystem function of soils that has gained increasing attention in recent years. Changes in soil \(\mathrm{C}\) impacts on, and feedbacks to, the Earth's climate system through emissions of \(\mathrm{CO}_2\) and \(\mathrm{CH}_4\) as well as storage of carbon removed from the atmosphere during photosynthesis (climate regulation; Table \(\PageIndex{1}\)). Soil organic matter itself also confers multiple benefits for human society, e.g. enhancing water purification and water holding capacity, protecting against erosion risk, and enhancing food and fibre provision through improved soil fertility (Table \(\PageIndex{1}\); Pan et al., 2013, 2014).

Table \(\PageIndex{1}\): Management actions affecting the soil carbon cycle and their impact on ecosystem services

Management action or other driver of change	Provisioning service impact	Regulating service impact	Supporting service impact	Cultural service impact
Land use change (conversion of for- est/grassland/wetland to cropland)	Increased production of food, fibre, and energy crops; reduced availability of natural raw materials; potential change in hydrology/water availability	Decreased soil C sequestration and storage - increased GHG flux; increased erosion and sediment yield - reduced regulations of water flow and quality	Primary production may be changed; nutrient recycling reduced if no inputs, increased if there are inputs	Lower recreation value; may have impact on cultural value in recreating diverse landscapes
Land use change (establishment of forest or grassland on agricultural land)	Raw material provision may be increased; agricultural production likely decreased (but not always, e.g. agroforestry)	Increased C sequestration; increased regulation of water flow and quality	Primary production may be changed; increased water recycling	Increased recreation value; may have impact on cultural value in recreating diverse landscapes
Intensified nutrient management through fertilization and liming	Increased production of food and other raw materials	Effect on net soil C sequestration uncertain; increased GHG flux from fertilizer production and use; water and air pollution	Increased primary production; increased nutrient recycling
Soil amelioration using organic amendments such as compost and biochar	Increased food production; more raw materials; more water available for plant growth	Increased C sequestration; increased water purification value	Increased primary production; increased nutrient cycling; improved water infiltration and retention
Diversification of crop production systems (i.e. more perennials, reduced bare fallow)	Potential impact on agricultural production (±); more diverse products	Increased C sequestration; increased purification value	Changed primary production; increased nutrient retention; improved water infiltration and retention	Improved cultural value from more diverse landscapes
Replacement of hay forage production with pasture use on grasslands	No impact	Effect on C sequestration uncertain		Increased recreation value; may have impact on cultural value in recreating diverse landscapes
Improved grazing management	Increased food production; reduced runoff and improved water use	Increased C sequestration; increased purification value; water flow regulation	Increased primary production; improved water infiltration and retention

Soil is an important \(\mathrm{C}\) reservoir that contains more \(\mathrm{C}\) (at least \(1500-2400 \mathrm{Pg} \mathrm{C})\) than the atmosphere \((590 \mathrm{Pg} \mathrm{C})\) and terrestrial vegetation \((350-550 \mathrm{Pg} \mathrm{C})\) combined (Schlesinger and Bernhardt, 2013; Ciais et al., 2013), and an increase in soil \(\mathrm{C}\) storage can reduce atmospheric \(\mathrm{CO}_2\) concentrations (Table \(\PageIndex{1}\); Whitmore et al., 2014). All three reservoirs of \(C\) are in constant exchange but with various turnover times, with soil as the largest active terrestrial reservoir in the global C cycle (Lal, 2008). Carbon storage in soils occur in both organic and inorganic form. Organic C stocks in the world's soils have been estimated to comprise \(1500 \mathrm{Pg}\) of \(\mathrm{C}\) to \(1 \mathrm{~m}\) depth and \(2500 \mathrm{Pg}\) to \(2 \mathrm{~m}\) (Batjes, 1996). Recent studies have shown that the soil \(\mathrm{C}\) pool to \(1 \mathrm{~m}\) depth may be even greater and could account for as much as \(2000 \mathrm{Pg}\). These higher values are mainly based on increased estimates of the C stored in boreal soils under permafrost conditions (Tarnocai et al., 2009), in which decomposition is inhibited by low temperature, lack of oxygen, and low \(\mathrm{pH}\) in waterlogged soils, e.g. peats (Smith et al., 2010). Although the highest C concentrations are found in the top \(30 \mathrm{~cm}\) of soil, the major proportion of total \(\mathrm{C}\) stock is present below \(30 \mathrm{~cm}\) depth (Batjes, 1996). In the northern circumpolar permafrost region, at least \(61 \%\) of the total soil \(\mathrm{C}\) is stored below \(30 \mathrm{~cm}\) depth (Tarnocai et al., 2009). Peatlands are particularly important components of the global soil carbon store, covering only \(3 \%\) of the land area but containing around \(500 \mathrm{Pg} \mathrm{C}\) in organic-rich deposits ranging from 0.5 to \(8 \mathrm{~m}\) deep (Gorham, 1991; Yu, 2012), with storage in deeper layers as yet unquantified.

In arid and semi-arid soils, significant inorganic \(\mathrm{C}\) can be present as carbonate minerals (typically \(\mathrm{Ca} / \mathrm{MgCO}_3\), called "calcrete" or "caliche" in various parts of the world), formed from the reaction of bicarbonate (derived from \(\mathrm{CO}_2\) in the soil) with free base cations, which can then be precipitated in subsoil layers (Nordt et al., 2000). Soils derived from carbonate-containing parent material (e.g. limestone) can also have significant amounts of inorganic \(\mathrm{C}\). The inorganic C pool globally is large, estimated to be \(\sim 750 \mathrm{Pg}\) \(\mathrm{C}\) to a depth of \(1 \mathrm{~m}\) (Batjes, 1996). However, in most cases, changes in inorganic C stocks are slow, are not amenable to traditional soil management practices, and do not play a significant role in terms of most ecosystem services (though a major exception is the geoengineering proposal to add finely ground silicate minerals to soils, which will then weather to carbonates, taking up CO2 in the process; Köhler et al., 2010).

The net balance of soil \(\mathrm{C}\) depends on the inputs of \(\mathrm{C}\) to soils relative to \(\mathrm{C}\) losses. Losses can occur via mineralization (i.e. decomposition), leaching of dissolved \(\mathrm{C}\), and carbonate weathering (Smith, 2012; Schlesinger and Bernhardt, 2013). Thus, the soil organic C stock may either increase or decrease in response to changes in climate and land use practices (Smith et al., 2015). Furthermore, rates of SOC stock change in different parts of the profile can vary for different

soils and types of perturbation, because some portion of the C stored in soil, mainly in topsoil, turns over rapidly, while other soil \(\mathrm{C}\) fractions can have a long residence time (von Lützow et al., 2008; Rumpel and Kögel-Knabner, 2011). The accumulation of stabilized \(\mathrm{C}\) with long residence times in deep soil horizons may be due to continuous transport, temporary immobilization and microbial processing of dissolved organic matter within the soil profile (Kalbitz and Kaiser, 2012), and/or efficient stabilization of root-derived organic matter within the soil matrix (Rasse et al., 2005). The process of soil formation - i.e. the development of depth, horizons, and specific properties - is itself a supporting service (Table \(\PageIndex{1}\)).

High SOC content also improves other chemical and physical soil properties, such as nutrient storage (supporting service), water holding capacity (supporting and regulating service), aggregation, and sorption of organic or inorganic pollutants (regulating service). Carbon sequestration in soils may therefore be a cost-effective and environmentally friendly way to not only store \(\mathrm{C}\) for climate regulation but also enhance other ecosystem services derived from soil, such as agricultural production, clean water supply, and biodiversity (Table \(\PageIndex{1}\); Pan et al., 2013) by improving soil organic matter (SOM) content and thereby soil quality (Lal, 2004). Moreover, processes which improve SOM may themselves provide services, e.g. use of cover crops, which can provide provisioning or water regulation services while improving soil C (Table \(\PageIndex{1}\)). SOM or soil carbon are widely used proxy variables for soil health (e.g. Kibblewhite et al., 2008).
\(C\) cycling:

Carbon enters the soil as aboveground or belowground plant litter and exudates. \(\mathrm{C}\) input is not homogenous within the soil profile. Whereas topsoil receives higher amounts of aboveground litter, subsoil C originates from root \(\mathrm{C}\) as well as dissolved \(\mathrm{C}\), transported down the soil profile. Root \(\mathrm{C}\) has a greater likelihood of being preserved in soil compared to shoot \(\mathrm{C}\), and was therefore hypothesized to account for most of the SOC (Rasse et al., 2005). The majority of plant litter compounds pass through and are modified by the soil biota. Thus, SOM is composed of plant litter compounds as well as microbial and, to a smaller extent, faunal decomposition products (Paul, 2014). It is a complex biogeochemical mixture comprising molecules derived from organic material in all stages of decomposition. Some organic matter compounds, including microbial decomposition products, may be stabilized for centuries to millennia by binding to soil minerals or by physical occlusion into micro-aggregates (von Lützow et al., 2008), for example with iron oxyhydrates (Zhou et al., 2009), or through protection by occlusion within soil aggregates (Dungait et al., 2012). The inherent chemical recalcitrance of some plant litter compounds (e.g. lignin) has a minor influence on their longevity in soil (Thévenot et al., 2010), whereas the location of SOM within the soil matrix has a much stronger control on its turnover (Chabbi et al., 2009; Dungait et al., 2012). Mineralassociated SOM is predominantly composed of microbial

products (Miltner et al., 2012). Therefore, microbial use efficiency of plant inputs largely determines SOM stabilization through interaction with the mineral phase (Cotrufo et al., 2013), in addition to the environmental controls discussed elsewhere in this section. In peatlands, organic matter is stabilized by high water tables that slow down biological activity and decomposition. SOM is mineralized to carbon dioxide \(\left(\mathrm{CO}_2\right)\) in aerobic environments, or reduced to methane \(\left(\mathrm{CH}_4\right)\) in anaerobic environments. Soil \(\mathrm{CO}_2\) efflux, resulting from SOM mineralization, and from rhizosphere respiration and inorganic \(\mathrm{C}\) weathering, is the largest terrestrial flux of \(\mathrm{CO}_2\) to the atmosphere \((\sim 60 \mathrm{Pg} \mathrm{C}\); the sink of carbon on the other hand contributes to the climate regulation service; Smith, 2004). This flux is an order of magnitude larger than anthropogenic \(\mathrm{CO}_2\) emissions due to fossil fuel burning and land use change (1.1 \(\mathrm{PgCyr}^{-1}\); Ciais et al., 2013). Under anaerobic conditions, \(\mathrm{CH}_4\) is formed by methanogenic microorganisms. A proportion of this \(\mathrm{CH}_4\) is oxidized to \(\mathrm{CO}_2\) by methanotrophic microorganisms, but a proportion can be emitted from the soil surface (Reay et al., 2010). Since \(\mathrm{CH}_4\) is many times more potent as a greenhouse gas than \(\mathrm{CO}_2\) on a per-molecule or per-mass basis (Ciais et al., 2013), soil \(\mathrm{CH}_4\) emissions and their mitigation play an important role in climate regulation (Table \(\PageIndex{1}\)).

Fire may affect many ecosystem services, including \(\mathrm{C}\) sequestration. For fires in natural ecosystems, a decrease in soil C storage is often observed initially, but through positive effects on plant growth, as well as input of very stable pyrogenic C, C storage may increase at longer timescales (Knicker, 2007). An additional long-term C pool in many soils is pyrogenic carbon (PyC), formed from partially combusted (i.e. pyrolysed) biomass during wildfires or other combustion processes (Schmidt and Noack, 2000). Globally, soils are estimated to contain between 54 and \(109 \mathrm{Pg}\) PyC (Bird et al., 2015). Some of this PyC has a highly condensed aromatic structure that retards microbial decay, and can thus persist in soils for relatively long periods (Singh et al., 2012). Soil amended with industrially produced PyC (biochar) as a climate mitigation technique often shows no increase in soil respiration despite the additional carbon, the reduced ecosystem carbon turnover results in increased soil carbon storage (Stewart et al., 2013). PyC additions to soil affect regulating ecosystem services, such as \(\mathrm{C}\) sequestration, nutrient cycling, and adsorption of contaminants. However, PyC properties, and as result their effect on ecosystem services, may be strongly dependent on fire conditions.

Factors influencing soil C storage: Fundamentally, the amount of \(\mathrm{C}\) stored in a given soil is determined by the balance of \(\mathrm{C}\) entering the soil, mainly via plant production but also through manures or amendments such as organic sludge or biochar, and \(\mathrm{C}\) leaving the soil through mineralization (as \(\mathrm{CO}_2\) ), driven by microbial processes, and to a lesser extent leaching out of the soil of dissolved carbon and carbonate weathering. Locally, C can be lost or gained through soil erosion or deposition, leading to a redistribution of soil \(\mathrm{C}\), at landscape and regional scales (van Oost et al., 2007).

Consequently, the main controls on soil \(\mathrm{C}\) storage are the amount and type of organic matter inputs, the efficiency by which this is used by microbes, and the capacity of the soil to retain it by physical or chemical stabilization (Cotrufo et al., 2013). In most natural and agricultural ecosystems, plant productivity and subsequent death and senescence of biomass provide the input of organic \(\mathrm{C}\) to the soil system (Table \(\PageIndex{1}\)). Thus, higher levels of plant residue inputs will tend to support higher soil organic carbon stocks, and vice versa (Paustian et al., 1997), though this does not continue indefinitely (Zvomuya et al., 2008). Plants also affect soil C cycling by their specific mycorrhizal associations (Brzostek et al., 2015). Shifts in specific mycorrhizal associations affect SOM storage by contributing to both SOM formation and decomposition. Ectomychorrizhal turnover is a dominant process of SOM formation (Godbold et al., 2006), possibly due to the more recalcitrant nature of the chitin in fungal tissues, compared to the cellulose and lignin in plant residues. In arbuscular mycorrhizal fungi, it has been suggested that glomalin, a highly resistant glycoprotein, has an active role in aggregate formation and SOM stocks (Rillig, 2004). Symbiotic mycorrhizal fungi can also directly impact the turnover of organic matter by the production of exo-enzymes (Averill et al., 2014; Finzi et al., 2015).

In many regions of the world, SOM accumulates because of inhibition of microbial SOM decomposition, due to cold, dry, or anoxic conditions (Trumbore, 2009). In general, when water is not limiting, higher soil temperatures increase the rate of microbial decomposition of organic matter. Thus soil temperature is a major control of SOM storage in soil C cycle models (Peltoniemi et al., 2007). The temperature sensitivity of SOM decomposition is not, however, as straightforward as represented in most models but varies between the many different forms of chemical and physical protection of organic matter in soil (Conant et al., 2011; Zheng et al., 2012). Water influences soil C storage through several processes. Moist, but well-aerated, soils are optimal for microbial activity and decomposition rates decrease as soils become drier. However, flooded (saturated) soils have lower rates of organic matter decay due to restricted aeration and thus often have very high amounts of soil C (e.g. peat soils). High precipitation may also lead to \(\mathrm{C}\) transport down the soil profile as dissolved and/or particulate organic matter, as well as lateral transport through soil erosion and deposition. During dry periods, SOM decomposition is decreased, but after rewetting there may be an accelerated pulse of \(\mathrm{CO}_2\) emission in aerobic soils (Borken and Matzner, 2008), whereas drought and lowering water tables may increase decomposition in naturally anaerobic peats (Freeman et al., 2001; Clark et al., 2012). However, the effect of drought is not only direct via soil microbial activity. There are feedback loops concerning drought and \(\mathrm{C}\) storage via plant activities, such as litter input and rhizodeposition. Drought was found to affect

plant litter composition (Sanaullah et al., 2014), plant C flow and root exudation (Sanaullah et al., 2012), as well as the resulting enzyme activities in the rhizosphere (Sanaullah et al., 2011).

C cycling in soils is strongly linked to the cycling of \(\mathrm{N}\) and \(\mathrm{P}\). Since the \(\mathrm{C}: \mathrm{N}: \mathrm{P}\) stoichiometry in SOM is generally lower than in plant material - i.e. there is more \(\mathrm{N}\) and \(\mathrm{P}\) per unit \(\mathrm{C}-\mathrm{C}\) generally accumulates in aerobic soil where nutrients are not limiting (Alberti et al., 2014). Nevertheless, an increase in organic \(\mathrm{C}\) is often accompanied by increased \(\mathrm{N}\) resource use efficiency in croplands (Pan et al., 2009), especially when SOC is increased with biochar (Huang et al., 2013). In nutrient-limited peatlands, inputs of nitrogen and/or phosphorus within the tolerance levels of sensitive plant species have increased rates of carbon accumulation (Aerts et al., 1992; Turunen et al., 2004; Olid et al., 2014). The relationship between nutrients and C cycling is not straightforward, since nutrients are also needed by soil microbes to degrade SOM. Thus, nutrient addition can either decrease or increase C storage, depending on the initial SOM stoichiometry, the ability of the soil minerals to stabilize microbial products of decomposition, and the simultaneous effects on plant productivity and organic matter inputs to soils.

The amount and type of clay particles (and to a lesser extent silt particles) are the major factors controlling the quantity and composition of soil C (Sollins et al., 1996; von Lützow et al., 2006). Clays are mainly sheet-like crystals of silicon and aluminium, known as phyllosilicates, often located as skins coating soil aggregates. In clay-rich soils, higher organic matter content and a greater concentration of O-alkyl C derived from polysaccharides may be expected compared to sandy soil, which are characterized by lower C contents and high concentrations of alkyl C (Rumpel and Kögel-Knabner, 2011). Aliphatic material may be responsible for the hydrophobicity of soils, which can lead to reduced microbial accessibility and therefore increased C storage (Lorenz et al., 2007). Many of the OM-matrix interactions are driven by expandable and non-expandable phyllosilicates, which interact with organic compounds through their large surface areas, micropores, and micro-aggregation, particularly in acid soils. In neutral and calcareous soils, polyvalent cations (especially \(\mathrm{Ca}^{2+}\) ) predominate in the interaction mechanism, forming bridges between the largely negatively charged SOM and negatively charged phyllosilicates (Cotrufo et al., 2013). Short-order silicates, like allophane, provide some of the strongest organo-mineral interactions and stabilize both proteins and carbohydrate monomers, though their occurrence is very geographically restricted (Buurman et al., 2007; Dümig et al., 2012: Mikutta and Kaiser, 2011). Pedogenic oxides (for example iron oxyhydrates in rice paddies) usually act as a coating of soil mineral particles and stabilize carbon, contributing to a higher \(\mathrm{C}\) storage and stability compared to other soils (Song et al., 2012).

Bioturbation (the mixing of soil by organisms) may further influence the amount as well as the chemical nature of soil C. It greatly influences the heterogeneity of soils by creating hotspots of carbon and biological activity. On biologically active sites, incorporation and transformation of organic compounds into soil is usually enhanced, leading to more organo-mineral interactions and increased \(\mathrm{C}\) storage (Wilkinson et al., 2009).

Microbial decomposition of SOM may be stimulated by the input of labile (easily decomposed) organic matter through the priming effect (Jenkinson et al., 1971). Positive priming refers to greater mineralization of otherwise stable \(\mathrm{C}\) through shifts in microbial community composition and activity (Fontaine et al., 2003). However, in some cases, the addition of organic matter to soil may also impede mineralization of native SOM (negative priming effect), thereby protecting SOM from its decomposition. Plant communities (Table \(\PageIndex{1}\)) are the main controlling factors of these processes because they influence organic matter input and microbial activity by their effects on soil water, labile \(\mathrm{C}\) input, \(\mathrm{pH}\), and nutrient cycling (Kuzyakov et al., 2000).

By storing and cycling C, nutrients, and water, soils provide supporting services like soil formation and nutrient and water retention, which underpin both primary production and landscape hydrology (the processes which deliver provisioning services such as food, fibre, and water; Table \(\PageIndex{1}\)), in addition to the regulating services such as climate regulation already discussed (Fig. 1). To ensure that soils continue to provide these key services, soil will require to be managed for \(\mathrm{C}\) preservation - thus mitigating climate change - while simultaneously permitting continued SOM recycling (Table \(\PageIndex{1}\)). Janzen (2006) pointed to this dilemma, that there is a tradeoff between improved soil fertility to support the provisioning services of food/timber production and the regulating service of soil carbon sequestration aiding climate regulation. Despite knowledge on which practices are likely to lead to improved SOC status, a better understanding of the controls on SOM distribution, stabilization, and turnover will help to better target these practices. This will be an important contribution to the mitigation of greenhouse gases, while assuring decomposition and, with it, the cycling of nutrients necessary to support food production. Table \(\PageIndex{1}\) summarizes management actions affecting the soil carbon cycle and their impacts on ecosystem services.

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Excerpted from:

Smith, P., Cotrufo, M. F., Rumpel, C., Paustian, K., Kuikman, P. J., Elliott, J. A., McDowell, R., Griffiths, R. I., Asakawa, S., Bustamante, M., House, J. I., Sobocká, J., Harper, R., Pan, G., West, P. C., Gerber, J. S., Clark, J. M., Adhya, T., Scholes, R. J., and Scholes, M. C.: Biogeochemical cycles and biodiversity as key drivers of ecosystem services provided by soils, SOIL, 1, 665–685, https://doi.org/10.5194/soil-1-665-2015, 2015. https://soil.copernicus.org/articles/1/665/2015/