Sulphur Cycling in Upland Agricultural Systems
J. J. SCHOENAU and J. J. GERMIDA
|Department of Soil Science, University of Saskatchewan, Saskatoon,|
|10.2 THE NATURE OF SULPHUR IN UPLAND AGRICULTURAL SYSTEMS|
|10.3 SOIL-ATMOSPHERE INTERACTIONS|
|10.4 STABLE ISOTOPE INVESTIGATIONS|
|10.5 MODELLING OF MINERALIZATION- IMMOBILIZATION PROCESSES|
In upland agricultural systems (oxic soils), the major transformations of S are mineralization, immobilization, and oxidation. Such transformations often result in losses or gains of S in the soil-plant system through processes such as leaching and S gas evolution and absorption. As a result, the S load in adjacent hydrospheric and atmospheric systems is altered. S transformations can be greatly affected by small changes in the environment which, in turn, can cause large shifts in the size of the S pools. To predict accurately the effect of a perturbation such as global warming on S load, detailed knowledge of S forms and transformations in the system is required. In the past, progress in elucidating the soil-plant S cycle has been limited by the complex nature of soil S pools, in particular soil organic S, and the lack of suitable investigation techniques. Analytical difficulties have always been a major limitation in advancing knowledge of S cycling. For example, our progress in understanding soil organic S turnover has stagnated somewhat in recent years as we continue to rely largely on characterization methods developed 30 years ago which separate organic S into only two broad groups: HI-reducible S and C- bonded S.
A knowledge of process response to environmental factors is also critical to accurate prediction of S flux at all scales of interest. Only when the relationship between an environmental parameter and a process such as S mineralization is clearly defined can fluxes for all subunits in a body of interest be accurately predicted and combined to arrive at a flux estimate for the whole region. The predictive tools used to estimate flux are usually mathematical models of a process and must account for the complex interactive effect of environmental parameters such as temperature and moisture.
This chapter addresses the S cycle in upland agricultural systems by considering the forms and transformations of S and their relationship to other components of the global S cycle, in particular the atmosphere. Stable S isotope techniques and process modelling are introduced as current approaches to identifying and quantifying S forms and fluxes.
In well drained, non-saline surface soils, more than 90% of the S is organic, and this sulphur is the main source of plant available sulphur (sulphate) in soil (Bettany and Stewart, 1983; Freney, 1986). Organic S is separated into two major groups based on susceptibility to reduction by reducing agents: (1) HI- reducible S, and (2) C-bonded S. The HI-reducible fraction consists of organic S that is directly reducible to H2S by hydriodic acid. It includes S atoms which are not directly bonded to C but are linked to C via an oxygen, nitrogen, or sulphur atom (e.g. C-O-SO3- linkages) (Freney, 1961,1967, 1986; Williams, 1975; Biederbeck, 1978). The HI-reducible fraction, commonly referred to as organic sulphate, is thought to consist primarily of sulphate esters as well as thioglucosides and sulphamates. Organic sulphates are easily converted to inorganic sulphate by mild physical and chemical treatments such as wetting and drying and by treatments with acids and bases (Spencer and Freney, 1960; Freney, Melville and Williams, 1969). Organic sulphates are thought to be associated mainly with the low molecular weight fraction of soil organic matter (Bettany, Stewart and Halstead, 1973; Schoenau and Bettany, 1987), and are considered to be the most labile fraction. These conclusions were based on the ease of hydrolysis and predominance in the non-recalcitrant fractions of soil organic matter (Freney, 1967; Freney, Melville and Williams, 1975; Fitzgerald, 1976, 1978; Bettany, Stewart and Saggar, 1979). Organic sulphates comprise 30 to 70% of the total organic S in surface soils (Bettany, Stewart and Halstead, 1973).
Even though the organic sulphate fraction constitutes an important fraction of soil organic sulphur, its origins and mode of formation remain unclear. In plant residues, more than 90% of the S usually exists in the C-bonded form, whereas upwards of 70% of the humus S derived from such residues may be comprised of organic sulphates (Bettany, Stewart and Saggar, 1979; Roberts and Bettany, 1985). It is reasonable to suggest a microbial role in the production of such compounds during degradation. However, whether the organic sulphates are metabolic by-products, components of dead cells, or products formed by reaction of inorganic sulphate with humic constituents, remains a point of debate.
The C-bonded S fraction consists of S which is directly bonded to C, in the forms of S-containing amino acids in free and combined states, and sulphonates (Freney, 1967; Freney, Melville and Williams, 1970; Biederbeck, 1978). C-bonded S can be distinguished analytically from the organic sulphate fraction because it is not directly reducible to H2S by hydriodic acid. C- bonded S is calculated as the difference between total organic S and organic sulphate. A portion of C-bonded S, believed to be mainly amino acid S, can be directly measured by reduction to H2S with Raney Ni (Lowe and DeLong, 1963; Freney, Melville and Williams, 1970, 1975). The residual fraction is highly resistant to chemical degradation, and is probably not a significant source of available S (Biederbeck, 1978). The C-bonded S fraction is believed to be more associated with the highly condensed aromatic humus core and a more stable organic S form that organic sulphate (Bettany, Stewart and Halstead, 1973; Bettany Stewart and Saggar, 1979). However, a portion of the C-bonded S fraction, possibly consisting of simple amino acid S, contributed significantly to the soluble sulphate pool in short-term tracer experiments (Freney, Melville and Williams, 1975; Maynard, Stewart and Bettany, 1985) and field studies (McLachlan and DeMarco, 1975).
Soil microbial biomass is defined as the living part of soil organic matter . The microbial biomass sulphur (MB-S) in agricultural soils comprises about 2-3% of the total S. However, detailed information on the role of MB is limited as measurement methods have only recently been developed for this fraction. Saggar, Bettany and Stewart (1982) proposed a method involving fumigation with chloroform and extraction with 0.01 M CaCl2 to measure MB-S in soil. Using a modification of this method, Gupta (1989) determined that the quantities of MB-S in Saskatchewan agricultural soils ranged from 2.5 to 12.5 µg (S) g-l (0.08 to 0.39 µmol g-l), accounting for 1-4% of the total organic S.
Despite its small size, the MB is a highly active fraction that acts as the driving force behind mineralization-immobilization and oxidation-reduction processes. Microorganisms rapidly assimilate simple, free organics such as low molecular weight ester sulphates and amino acids, and also inorganic sulphate when the S in the substrate is insufficient to meet microbial requirements (immobilization). Organic S, assimilated in excess of microbial requirements, is mineralized by intracellular enzymes and released as sulphate (biological mineralization) (Biederbeck, 1978). Microorganisms can also secrete extracellular sulphatase enzymes such as arylsulphatases and other sulphohydrolases which are capable of hydrolysing organic sulphate functional groups in soil organic matter (biochemical mineralization) (Tabatabai and Bremner, 1970; Gupta, 1989). Mineralizable organic S found in soils appears to be closely associated with the microbial biomass present as binding agents in macroaggregates (Gupta and Germida, 1988a). Net mineralization of microbial biomass S is reported to be dependent on both microbial and soil C : S ratios as well as percentage organic sulphate (Gupta, 1989). For example, native pasture soils with a wide C : S ratio exhibited low net release of microbial sulphur due to re-assimilation of released microbial sulphur by the new flush of microbial growth.
A conceptual model for the various microbial components and processes of the sulphur cycle in soil is illustrated in Figure 10.1. The MB plays a central role as a sink and source for sulphate in the soil, with the interactions between various microbial populations being an important factor regulating sulphate levels. Gupta and Germida (1989) have recently observed that amoebal grazing of bacteria enhances the biochemical and biological mineralization of MB-S in soil. As well, addition of elemental S fertilizers has been reported to decrease populations of predatory protozoa, suggesting reduced nutrient turnover via microbial predation as a result of elemental S fertilization (Gupta and Germida, 1988b). Such findings point out the need to consider microbial ecology and population dynamics when attempting to model S turnover in soils.
Evidence for downward leaching of low molecular weight, S-rich organics during soil formation has led to the suggestion that organic S leaching is an important long-term translocation process in the soil-plant S cycle (Schoenau and Bettany, 1987; Roberts, Bettany and Stewart, 1989). The removal of labile, S-rich organics from surface horizons by leaching may influence S turnover rates and act as another S export mechanism if the organics are lost entirely from the profile.
While organic S is the main reservoir of S in the soil, inorganic S in the form of sulphate is the highly dynamic plant available pool. Inorganic sulphur may occur in agricultural soils as sulphate and as compounds of lower oxidation states such as sulphide, polysulphide, sulphite, thiosulphate, and elemental S. In well-drained, well-aerated soils, the most common form is sulphate and the amounts of reduced S are generally less than 1 µg (S) g-l (0.03 µmol g-l) (Freney, 1961; Williams, 1975). Sulphate may occur in soils as water-soluble salts, adsorbed sulphate, or insoluble forms. In surface soils of neutral to alkaline pH, sulphate exists mainly in the soil solution. In acid soils, considerable quantities of sulphate may be adsorbed to the mineral fraction of the soil. All of the water-soluble sulphate and most of the adsorbed sulphate is available for plant uptake (Freney and Spencer, 1960).
In highly leached soils, sulphate adsorption is an important mechanism for retaining sulphate in the profile. Sulphate adsorption is negligible above pH 6.5 and increases with decreasing pH (Kamprath, Nelson and Fitts, 1956). In soils containing large amounts of iron and aluminum oxides, sulphate adsorption is generally greater (Chao, Harward and Fang, 1962; Mitchell, David and Harrison, this volume). Sulphate adsorption occurs at many energetically different reaction sites so that sulphate adsorption may vary both within and among soils depending on the inorganic constituents present (Williams, 1975).
Insoluble sulphate can exist in the soil as barium and strontium sulphates, sulphates associated with calcium carbonate, and iron and aluminum sulphates (Williams, 1975). Sulphate as a co-crystallized impurity in calcium carbonate can account for up to 95% of the total S in calcareous soils (Williams and Steinbergs, 1964). This sulphate is suggested to make little, if any, contribution to the plant available sulphate pool (Williams and Steinbergs, 1964).
Figure 10.1. A conceptual model for the various microbial components and processes of the sulphur cycle in soil (Gupta. 1989). 1. Fertilizer dissolution and oxidation of elemental S fertilizers. 2. Oxidation of reduced S in minerals and release of sulphate. 3. Microbial assimilation of S from minerals. 4. Plant uptake of sulphate. 5. Microbial assimilation and immobilization of sulphate. 6. Mineralization. 7. Dead and decomposing root material and residual straw. 8. Microbial decomposition of plant material. 9. Humification and/or stabilization of resistant plant residues. 10. Humification and/ or stabilization of resistant microbial residues. 11. Microbial utilization of resistant organic S. 12. Release of labile S compounds during plant residue decomposition. 13. Microbial decomposition and organic matter turnover due to other environmental factors such as drying and wetting, freeze and thaw cycles. 14. Microbial assimilation of S from labile organic S compounds. 15. Release of labile microbial S due to grazing by predators. 16. Biological mineralization. 17. Biochemical or enzymatic mineralization. 18. End-product controlled activity of sulphatase enzymes. Microbial biomass: (A) Protozoa and nematodes feeding on bacteria; (B) Bacteria feeding on damaged fungal mycelium and on compounds released from dead hyphae; (C) Protozoa and nematodes feeding on fungi; (D) Protozoa grazing on algae; (E) Nematodes grazing on protozoa; and (F) Actinomycetes feeding on bacteria
In the subsurface horizons of soils with restricted downward leaching, large accumulations of sulphate may occur in the form of gypsum ( CaSO4. 2H2O ) or other water-soluble salts. Gypsum is a slightly soluble salt and produces a high sulphate activity in solution. Consequently, soils developed under arid or semi-arid environments with accumulations of gypsum in the rooting zone are seldom found to be S deficient (Bettany, Janzen and Stewart, 1983). Excessive quantities of sulphate salts within the rooting zone can increase osmotic potential to the point where plant growth is severely restricted. In contrast, in humid climates or under irrigation with good drainage, sulphate may be lost entirely from the soil profile by deep leaching, with reported losses as high as 30 g (S) m-2 a-l (1 mol m-2 a-1) (Harward and Reisenauer, 1966).
In poorly drained soils, large quantities of inorganic S may be present in reduced forms, mainly sulphides. Sulphides are usually found in subsoils below the water table. Re-oxidation of sulphide to sulphate under aerobic conditions can result in the formation of acid-sulphate soils (Fleming and Alexander, 1961).
Elemental S is commonly used as a fertilizer in S-deficient agricultural systems, but can also be a soil pollutant in the form of wind-blown dust from stockpiles near sour gas processing plants (Maynard, Germida and Addison, 1986). The elemental S is oxidized to sulphate by the soil microbial population. This conversion is necessary to render the S plant available, with the rate of oxidation being a major factor influencing the effectiveness of elemental S fertilizer. Oxidation of S occurs readily in some soils, but chemical, physical, and biological factors limit oxidation rates in other soils (Janzen and Bettany, 1987). Temperature and moisture affect microbial activity and consequently oxidation rate. Very slow rates are observed in cold, dry soils. Decreasing the particle size of the elemental S results in higher oxidation rates due to the increased surface area.
The composition of the microbial community plays an important role in regulating S oxidation. Autotrophic S oxidizers were once believed to entirely dominate S oxidation in soils, and are very active in certain environments (Lee, Watkinson and Lauren, 1988). However, enumeration of autotrophic and heterotrophic elemental S oxidizers in Saskatchewan soils has revealed heterotrophs to be the most abundant sulphur oxidizers (Germida, 1985; Lawrence and Germida, 1988). Between 3 and 37% of the total heterotrophic population of these soils were found to be capable of oxidizing elemental sulphur to thiosulphate. Heterotrophic sulphur oxidizers tended to increase with increasing pH, total S, and clay content, whereas populations of autotrophic thiosulphate oxidizers were negatively correlated with these factors (Lawrence and Germida, 1988). Nevertheless, it remains difficult to partition accurately the rates of sulphur oxidation between heterotrophic and autorophic oxidation. The plant rhizosphere is a potentially important site for S oxidation and reduction but S fluxes in crop rhizospheres are not well understood. Recent work by Grayston and Germida (1990) indicated that the rhizosphere supports greater sulphur oxidation with larger and more diverse populations of sulphur-oxidizing heterotrophs than non-rhizosphere soil.
In the last 20 years, the exchange of sulphur gases between soils, plants, and the atmosphere has received increased attention. A few studies (Banwart and Bremner, 1976; Minami and Fukushi, 1981) have shown that microbial breakdown of S-containing amino acids added to the soil can give rise to gaseous S compounds such as dimethyl sulphide, methyl mercaptan, dimethyl disulphide, carbonyl sulphide, and carbon disulphide. However, the contribution of indigenous organic S in soil to gas evolution during microbial decomposition is unknown. Recent work by Melillo and Stuedler (1989) has shown that addition of nitrogen to forest soils resulted in increased emission of carbonyl sulphide and carbon disulphide gases. Presumably this was due to increased production of S-containing amino acids, which were then transformed by the soil microbial population into volatile S compounds. It is highly likely that emissions of volatile S compounds are emitted from agricultural soils during residue breakdown in a fashion similar to that observed in forest soils. However, the lack of field measurements in cropping systems makes this difficult to corroborate.
The emission of volatile S compounds from plants has recently been proposed as an important constituent of the biogeochemical S cycle (Rennenberg, 1989). The importance of such emissions to the global atmospheric cycle is still uncertain but is probably not great (Andreae and Jaeschke, this volume). With regard to agroecosystems, Grundon and Asher (1988) have suggested losses of volatile S from fields of alfalfa and wheat to be in the order of tens of milligrams (S) m-2 a-l (in the order of 1 mmol m-2 a-l), whereas cotton may be losing upwards of hundreds of milligrams (S) m-2 a-l (in the order of 10 mmol m-2 a-l). These are typical of values reported for other types of vegetation, such as trees in tropical forests (Andreae and Jaeschke, this volume). The production of H2S and other volatile S compounds by plants is a light-dependent process which increases in response to root injury and S stress (Winner et al., 1981). Stress may be induced in the plants by high sulphate concentrations in the solution bathing the roots, or by exposure of the plants to high concentrations of SO2. This phenomenon may be important in soils containing high concentrations of sulphate, such as many saline soils, and where repeated injury to the roots occurs as a result of cultivation.
Atmospheric inputs of S to agroecosystems are difficult to quantify due to large variations from location to location (Tabatabai, 1984). Tabatabai (1984) has concluded that the amounts of S added in precipitation in rural areas of North America (50 to 1000 mg (S) m-2 a-l; 1.6 to 31 mmol m-2 a-l) are important in crop production to maintain the soil S balance, particularly in those areas where use of S fertilizers is low.
Figure 10.2 shows rough estimates of some of the S fluxes for western Canadian agroecosystems. Considerable uncertainty surrounds many of the estimates. For example, we were unable to assign even approximate values to gaseous S emissions. This points out the difficulty in constructing S budgets for agroecosystems with present information and thus the direction that future work must take.
Figure 10.2. Conceptual S cycle in agroecosystems. Numbers represent flux estimates (g (S) m-2 a-1) for S transformations in western Canadian soils
Naturally occurring variations in stable S isotope ratios (34S / 32S) are measured and expressed relative to a standard using the 'delta' (d34S) notation (Nriagu and Krouse, this volume). Samples which are enriched in 34S relative to the meteoritic standard bear a positive d34S while samples depleted in 34S have negative d34S values.
Variations in 34S / 32S ratios are created as a result of isotopic fractionation in chemical and biological transformations of S. In many reactions, the lighter isotope (32S) reacts preferentially, forming a product depleted in the heavy isotope (34S) while the remaining reactant is enriched in 34S (Krouse and Tabatabai, 1986; Nriagu and Krouse, this volume). The degree to which isotope abundances are altered is often specific to a reaction and can be used as an indirect indicator of a process occurring in a system. For example, sulphate reduction is associated with large depletions of 34S in the product sulphide. The d34S technique has frequently been used in geological studies to assist in tracing the origin of petroleum and ore deposits and in environmental studies which have attempted to trace the fate of S emissions from industrial sources (Krouse and Case, 1981, 1983; Krouse, Legge and Brown, 1984; Nriagu and Krouse, this volume).
In one of the earliest d34S studies involving soils, differences in S isotope composition were noted between different soil organic S fractions (Lowe, Sasaki and Krouse, 1971). Fuller et al. (1986) observed that the B horizon organic S in a podzolic soil profile had a d34S at least 5 % less than the horizon above or below it, which they attributed to isotope fractionation in biotic S transformations. Enrichments in 34S of up to 3 % have been observed in the organic sulphate fraction relative to other S fractions in western Canadian soils (Table 10.1) (Schoenau and Bettany, 1989). This probably reflects the high lability and turnover of organic sulphate (Figure 10.3). 34S enrichment may thus offer a new means of identifying and tracing labile organic S compounds in soil.
Variations in 34S abundance may reveal soil-plant-atmosphere interactions. The total S in wheat straw grown on saline, sulphate-saturated cultivated grassland soils was found to be enriched in 34S by about 3 %o relative to its sulphate source (Schoenau and Bettany, 1989) (Table 10.1). Winner et al., (1981) and Krouse, Legge and Brown (1984) have noted that plants which are exposed to high levels of S, either in the atmosphere or the solution bathing the roots, become enriched in 34S compared to surrounding S sources due to the emission of 34S-depleted H2S from the plant as a stress relief mechanism. The enrichment of 34S in the wheat plants thus provides indirect field evidence for the emission of 34S-depleted gaseous S compounds in response to S stress invoked by the high concentration of sulphate.
Table 10.1. d34S values of S fractions in a western Canadian
grassland soil (Schoenau and Bettany, 1989)
|-1.7 (wheat straw)|
|-2.7 (blue grama grass)|
The d34S approach can provide clues about the origin and fate of sulphur in a soil. In sulphate co-precipitated with soil carbonates, Schoenau and Bettany (1989) observed d34S values that were much more positive than any of the possible S inputs during soil formation. This led to the suggestion that much of the sulphate in the carbonate minerals was of oceanic origin (highly positive d34S), being precipitated with the carbonates during the evaporation of ancient inland seas. d34S values have been used to evaluate the downward movement of manure S applied to a grassland soil (Chae and Krouse, 1986). The d34S values of the C-bonded organic S fraction indicated penetration of manure S in this form down to the water table.
Figure 10.3 Predicted isotope composition of S fractions during mineralization of a plant residue with an initial d34S of 0%o (Schoenau and Bettany, 1989). Rapidly decomposed (labile) S fractions are characterized by 34S enrichment (more positive d34S) relative to recalcitrant fractions
As demonstrated above, 34S natural abundance variations are a useful tool in elucidating the soil-plant S cycle. As more information is obtained on the nature of stable S isotope fractionation, the utility of the d34S technique will undoubtedly increase. The recent availability of compounds artificially enriched or depleted in 34S offers a host of new research opportunities. For example, using 34S-labelled fertilizer materials and organic residues, it is now possible to trace the long-term fate of S additions in the field.
Mineralization-immobilization reactions regulate the supply of available S (solution sulphate) for biomass production. Sulphur mineralization- immobilization processes are strongly linked to the cycling of C, N, and
P and thus are often investigated simultaneously (Stewart, 1984).
In order to construct mathematical models which can be used to predict the net mineralization of S in different soils over a growing season, detailed information is required on how S mineralization rates change with time and environment. A first-order kinetic model derived from curve fitting to laboratory or field mineralization data has been used to predict N and S mineralization as a function of time (Stanford and Smith, 1972; Ellert and Bettany, 1988). This model is of the following form:
Ict = Mo (1 -e-kt)
where Ict is the cumulative inorganic S or N mineralized after time t (mg kg-1), t is the time from start of incubation (weeks), Mo is the potentially mineralizable organic S or N (may be determined by chemical extraction or extrapolation of mineralization curve), and k is the first-order rate constant (week-1).
Mathematical models of S cycling require temperature and moisture response functions to adjust mineralization rates for field conditions (Coughenour et al., 1980). If the rate constant as a function of temperature, moisture, and other environmental parameters is known, then the above expression may be used to predict S mineralization under a wide variety of conditions.
Figure 10.4. Cumulative sulphate-S mineralized from a wheat-fallow and perennial forage (alfalfa) field indicating S release and changes in mineralizable fractions predicted by best-fitting kinetic models. Included are schematic representations of the corresponding kinetic models
However, simple first-order expressions do not always adequately describe mineralization patterns observed in the field. For example, S mineralization in Saskatchewan soils cropped to perennial forage (alfalfa) has been observed to follow first-order kinetics, while soils cropped under a wheat/ fallow rotation showed an initial lag during the first two weeks of a 37-week incubation (Ellert and Bettany, 1988) (Figure 10.4). Similarly, S mineralization rates in native forest soils in Saskatchewan have shown much smaller responses to further temperature increases when temperature was warm (> 20° C) than cultivated soils, reflecting possible differences in substrate composition and/or the microbial community (B. Ellert, unpub. data).
Our understanding of mineralization-immobilization processes is far from complete. Temperature and moisture may influence the size and composition of the mineralizable fraction, which in itself is not yet well defined. If advances are to be made in understanding organic S turnover, improvements in characterization of organic S are vital. The separation of organic S into the two large, chemically 'fuzzy' pools of organic sulphate and carbon-bonded S is no longer sufficient for interpretation of the sophisticated incubation and tracer experiments now utilized. Attempts must be made to distinguish analytically the organic S pools of definite chemical and biological function. The recent development of the 'microbial biomass S pool' (Saggar, Bettany and Stewart, 1982; Gupta, 1989) is a good step in this direction.
At present, we have a basic understanding of the forms and amounts of S in agricultural soils and the processes controlling the supply of S to plants. The exchange of sulphur gases between the soil-plant system and the atmosphere is less well documented and understood. The current challenge is to predict accurately the impact of man-induced or natural changes on S loading in all components of the biosphere. Prerequisites to this are as follows:
To obtain a better definition of S forms and transformations in the soil plant S cycle accomplished through use of new approaches such as microbial biomass and stable isotope investigations.
To make more and better field measurements of soil-atmosphere S exchange from agroecosystems.
To understand how key processes in the S cycle respond to environmental factors, e.g. construction of models of mineralization and volatilization processes that include temperature, moisture, substrate, and microbial composition response functions.
To develop better approaches for extrapolating data from a few field sites or the laboratory to entire ecosystems, e.g. studies which emphasize how smaller components of a system combine to form the whole.
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