Causes of Soil Salinization, Sodification, and Alkalinization
Abstract and Keywords
Driving forces for natural soil salinity and alkalinity are climate, rock weathering, ion exchange, and mineral equilibria reactions that ultimately control the chemical composition of soil and water. The major weathering reactions that produce soluble ions are tabled. Where evapotranspiration is greater than precipitation, downward water movement is insufficient to leach solutes out of the soil profile and salts can precipitate. Microbes involved in organic matter mineralization and thus the carbon, nitrogen, and sulfur biogeochemical cycles are also implicated. Seasonal contrast and evaporative concentration during dry periods accelerate short-term oxidation-reduction reactions and local and regional accumulation of carbonate and sulfur minerals. The presence of salts and alkaline conditions, together with the occurrence of drought and seasonal waterlogging, creates some of the most extreme soil environments where only specially adapted organisms are able to survive. Sodic soils are alkaline, rich in sodium carbonates, with an exchange complex dominated by sodium ions. Such sodic soils, when low in other salts, exhibit dispersive behavior, and they are difficult to manage for cropping. Maintaining the productivity of sodic soils requires control of the flocculation-dispersion behavior of the soil. Poor land management can also lead to anthropogenically induced secondary salinity. New developments in physical chemistry are providing insights into ion exchange and how it controls flocculation-dispersion in soil. New water and solute transport models are enabling better options of remediation of saline and/or sodic soils.
Soil salinization refers to the process of salt accumulation in terrestrial landscapes. It occurs naturally where evaporation is high relative to precipitation (there is a seasonal water deficit) and leaching is insufficient to move salts out of the soil profile (Duchaufour, 1982; Schofield & Kirkby, 2003), often in landscapes that do not drain into the ocean, called “endorheic” drainage basins. Soils with accumulations of gypsum are Gypsisols; those with accumulations of calcium carbonates are Calcisols or sometimes Chernozems and Kastanozems in the World Reference Base (IUSS, 2015). Salinity in the form of sodium (Na) salts is associated with “[s]oils with limitations to root growth,” especially Solonetz (with a high content of exchangeable Na) and Solonchak (with high concentration of soluble salts) Reference Soil Groups (IUSS, 2015; see the article “Classification and Mitigation of Soil Salinization”); these are the saline soils that have received the most attention from an agricultural perspective because of their extensive area in Europe (e.g., Szabolcs, 1974) and in Australia (e.g., Rengasamy, Chittleborough, & Helyar, 2003; see the article “Soil Salinization”). However, extensive areas of saline soils occur elsewhere in the world (Ghassemi, Jakeman, & Nix, 1995; Schofield & Kirkby, 2003). Soils with a “natric” (from Arabic natroon, salt) diagnostic horizon have a dense subsurface horizon with a distinctly higher clay content than in the overlying horizon(s). A natric horizon has a high content of exchangeable Na and, in some cases, a relatively high content of exchangeable Mg. Diagnostic soil characteristics reflect soil-forming processes, thus the presence of a natric horizon reflects “sodification,” the process of sodium accumulation.
Soluble salts dissolve in water into charged ionic species (electrolytes), and thus saline soil solutions can conduct electricity. Saline soils are defined those with an electrical conductivity in a saturated paste extract at 25°C, a , and exchangeable of total exchangeable cations (Richards, 1954; see the article “Soil Salinization”). They are distinct from sodic soils with exchangeable and, in older literature, alkali soils with a . Salt-affected soils can be saline-sodic, with a saturated paste extract , a , and an exchangeable of total exchangeable cations. Soil salinity implies the presence of any salt, including chlorides , sulfates , nitrates , borates , carbonates , and bicarbonates of Na, calcium (Ca), magnesium (Mg), potassium (K), and iron (Fe) (Rengasamy, 2006). Alkali soil implies the presence of and , therefore these are also sodic soils. Sodicity refers to the amount of sodium present in a soil. Alkalinization is the process of rise in pH during the accumulation of sodium carbonates (this process is explained in more detail in the section “Formation of Soil Carbonates and Alkalinity”).
This article summarizes the causes and processes that give rise to soil salinity and alkalinity in the environment more generally. It outlines the role of rock weathering, release of Na ions, and soil solution chemistry in the formation of sodic soils. It summarizes the role of arid and semi-arid climate and water movement through landscapes in regulating soil chemical equilibria and mineral precipitates. It briefly illustrates how understanding of ion exchange and mineral equilibria can help in land management. Finally, it includes a case study from an area in South Australia where salinity problems in a dynamic geochemical setting are linked with the formation of potentially acid sulfate soils inland (Fitzpatrick, Fritsch, & Self, 1996; Fitzpatrick et al., 1999).
Sodium and Its Role in Sodification
Sodic soils and the natric diagnostic criteria in the World Reference Base for soil resources (IUSS, 2015) are defined on the basis of (1) their exchangeable sodium percentage (ESP), the proportion of on the cation exchange complex measured in similar units and expressed as a percentage; (2) their exchangeable sodium ratio (ESR), where ; or (3) their sodium adsorption ration (SAR), the ratio of the concentration of to and ions, where (Richards, 1954). In Australia an is used to indicate sodicity because at this ESP soils start to exhibit dispersion (Isbell, 1996), but traditionally, an is the threshold that defines sodic soil (Rengasamy, 2006). An ESP of 15 generally corresponds with an SAR of 13 (Richards, 1954).
The chemical composition of a soil solution is the result of acid-base reactions, precipitation and dissolution of solid phases, coordination reactions of metal ions and ligands, oxidation-reduction reactions, and adsorption-desorption process at soil–solution interfaces. The partial pressure of controls the dissolution and precipitation of soil carbonates and bicarbonates and the activities of ions, including , which control soil pH. These chemical processes are studied using principles of physical chemistry including chemical thermodynamics (Stumm & Morgan, 1970) and proper management of saline, sodic, and/or alkaline soils requires a good understanding of thermodynamics (e.g., Mau & Porporato, 2016; Šimůnek & Suarez, 1994).
Cation sorption and exchange reactions between ions in solution and clay mineral surfaces are responsible for the unique physicochemical properties and behavior of sodic soils. If cation exchange is represented by an equation similar to one used for chemical reactions in solutions, the reaction between a calcium-saturated soil exchange complex and a sodium chloride solution can be expressed as , where X designates the soil exchange complex (Richards, 1954). But generally the reaction does not go to completion as shown in the equation, because the soil solution contains a mixture of ions (i.e., it is not a pure sodium chloride solution), and as long as soluble calcium exists in the solution phase, there will be adsorbed calcium ions on the exchange complex. Near neutral pH, most soils have an exchange complex dominated by and . As the soil solution becomes concentrated through evaporation or transpiration by plants, the solubility limits of calcium sulfate, calcium carbonate, and magnesium carbonate are exceeded and these minerals precipitate. The result is an increase in the relative proportion of sodium ions in solution and consequently a replacement of the some exchangeable and by on the exchange complex.
Sodic soils with low EC can experience severe structural degradation and exhibit poor soil–water and soil–air relations (Rengasamy et al., 2003): Swelling and dispersion of soil aggregates dominated by on the exchange complex reduce the porosity and permeability of soils and increase the soil strength even at low suction (high water content). Soil strength is a measure of the capacity of soil to resist deformation and refers to the amount of energy (measured in megapascals, MPa) required to break apart aggregates or move implements through the soil; it affects the ability of plant roots to penetrate the soil. These adverse physicochemical conditions restrict water storage and movement, aeration, and nutrient uptake. The range of soil water content that does not limit plant growth and function (non-limiting water range) becomes narrow, and soils are either too wet immediately after rain or too dry within a few days for optimal plant growth (Rengasamy et al., 2003).
also has deleterious effects on soil structure under certain circumstances (Zhang & Norton, 2002). The saturated hydraulic conductivity (Ks) of an illitic soil column leached with a solution is much lower than if a solution is used (Quirk & Schofield, 1955), and Ks is consistently lower in Na–Mg systems than in Na–Ca ones (McNeal et al., 1968). Swelling and disaggregation, which reduces large pores, appear to be the dominant process causing the rapid initial decline of Ks, with clay dispersion and subsequent pore plugging becoming progressively important when electrolyte concentration is reduced to below the critical flocculation concentration (Zhang & Norton, 2002); the effects are observed for soils containing variable amounts of smectite, kaolinite, illite, and vermiculite, regardless of content (Keren, 1991). The general explanation given is that , compared with , has greater hydration energy and a larger hydration shell, which results in greater clay swelling; the greater swelling leads to weaker aggregate bonds and hence lower aggregate stability (Keren, 1991; Zhang & Norton, 2002).
Experiments show that, when soils are leached with salt solutions containing a mixture of a monovalent cation and a divalent cation until equilibrium between the soil and solution is established, the proportions of exchangeable monovalent and divalent cations present on the soil-exchange complex vary with the total cation concentration as well as with the monovalent:divalent cation ratio of the salt solutions. In the 1930s, Gapon found empirically that the influence of total cation concentration is taken into account and a linear relation with the exchangeable monovalent:divalent cation ratio is obtained when the molar concentration of the soluble monovalent cation is divided by the square root of the molar concentration of the soluble divalent cation (Richards, 1954). Thus for the reaction,
where is the Gapon coefficient.
The Gapon equation works well for describing exchange on smectite- and vermiculite-rich soils. It is still used extensively to predict exchange in arid environments (e.g., Mau & Porporato, 2016).
Cation Exchange and Solution Interface Chemistry
Understanding saline or sodic soil behavior requires an understanding of physicochemical phenomena of aqueous solutions of electrolytes at colloidal interfaces. Soil aggregrates are complex flocculated materials composed of small mineral particles and organic molecules, generally assumed to have a negatively charged surface, although this depends on the soil pH and point of zero charge (Frenkel, Levy, & Fey, 1992; Sparks, 1995). The electric field originating from the surface charge determines the force of adsorption of ions. Traditionally diffuse double layer theory or triple layer theory is invoked to explain the behavior of cations in solution on the negatively charged clay mineral-organic soil colloidal exchange complex (Sparks, 1995). In the first layer, the negative surface charge of the colloid is compensated by positively charged counterions tightly held due to electrostatic chemical interactions. The second layer is loosely associated with the colloidal interface and is thus called the “diffuse layer.” It consists of free ions that move in the fluid under the influence of weaker chemical attraction forces rather than being tightly held. The characteristic thickness of the double layer is the Debye length. In aqueous solutions it is typically on the scale of a few nanometers, and its thickness decreases with increasing concentration of the electrolyte.
If the electrolyte concentration is below the critical flocculation concentration, the Debye length increases and dispersion of the colloids occurs. The critical flocculation concentration for soil systems depends not only on the electrolyte concentration but also on the solution pH and point of zero charge (pzc) of soil colloids. If the , the colloid surface has a net positive charge, conversely, if the , the surface has a net negative charge. Flocculation is greatest at , and soils tend to weather toward a at which mineral dissolution is at a minimum.
The Debye-Hückel equation has been used to predict mean activity coefficients for ions in solutions that deviate from ideal behavior, when their thermodynamic properties would be analogous to those of a mixture of ideal gases; it works best for dilute solutions. The Debye-Hückel law predicts that the mean activity coefficient of an ion in an electrolyte solution is proportional to the square root of the solution’s ionic strength. Pitzer equations have been introduced to model ion interactions that cannot be neglected in concentrated solutions (Šimůnek & Suarez, 1994).
Some ions are held more tightly than others. Hydration shell and effective ionic radius are parameters introduced to fit experimental data for observed specific ion effects in soil chemistry and physics of water at interfaces, some of the Hofmeister effects (Lo Nostro & Ninham, 2012). With only monovalent alkali metal cations in solution, the general order of selectivity is and is related to the size of the hydrated ionic radius; , the ion with the smallest hydrated radius, can approach the negatively charged interface the closest and is held the most tightly. However, with divalent ions there is little consistency in the selectivity (Sparks, 1995). In soil solutions with ions of different valence, generally the higher charged ion will be preferred, for example, . Thus, when both and are present in a soil solution, the is held in the first layer and in the diffuse layer. Anions such as phosphate and can be held tightly on the surface of positively charged oxide particles (Sparks, 1995), but and are in the diffuse layer.
Furthermore, polarization must be considered. Polarization is the distortion of the electron cloud about an anion by a cation. The smaller the hydrated radius of the cation, the greater the polarization, and the greater its valence, the greater its polarizing power. With anions, the larger they are, the more easily they can be polarized (Sparks, 1995). Organic and inorganic anions in electrolyte solutions have been observed to increase the critical flocculation concentration of clay mineral suspensions by an order of magnitude (Frenkel et al., 1992). This means that the overall composition of the soil, its solid and solution components, needs to be considered to understand sodic soil behavior. Yet the role of organic matter (the composition of organics) in explaining sodic soil behavior has not been studied.
Recently quantum mechanical forces have been introduced to classical double layer theory (Lo Nostro & Ninham, 2012). Lo Nostro and Ninham (2012) have extended the Debye-Hückel theory and shown how Hofmeister effects (as ion-specific quantum forces) depend on an interplay between specific surface chemistry, surface charge density, pH, buffer, and counterion with polarizabilities and ion size. These new developments in physical chemistry have prompted Liu et al. (2013) to derive new selectivity coefficients for different ion exchange pairs as a function of surface potential in different soils. These developments herald a new era in soil chemistry that should help in better understanding the behavior of saline and sodic soils, where concentrated electrolyte solutions occur seasonally and structural stability changes as a result of leaching with more dilute rain or irrigation water.
Climate and the Role of Water Movement
Salt content changes down a soil profile with seasonal moisture fluctuations, evapotranspiration, and infiltration rate, and sometimes soils only experience transient salinization in the subsoil (Rengasamy et al., 2003). Secondary salinization can arise when salts accumulate near the soil surface as a result of rising water tables due to land management practices that change the soil hydrology, such as irrigation or tree clearing (Cisneros, Cantero, & Cantero, 1999; Rengasamy, 2006; Schofield, Thomas, & Kirkby, 2001; Williamson, 1986). The chemical composition of the irrigation water and underlying water table is critically important in determining the salts that precipitate.
A large fraction of natural and anthropogenically induced secondary salinity occurs in desert and grassland biomes, in savanna ecosystems and rangelands which are closely associated with semi-arid, seasonally contrasted climates in the tropics and subtropics. Carbonate precipitation also occurs in such climatic conditions. In Australia, soil carbonate is generally rare to absent in high-rainfall regions, including the Wet Tropics and much of the east coast of the continent, coastal southwest Western Australia, and Tasmania (Wilford, de Caritat, & Bui, 2015). Nevertheless, the occurrence of salt-affected soils is not a perfect match with climate, even when also taking lithology (a potential source of solutes released by weathering) into account (Schofield & Kirkby, 2003). Global patterns of salinization have changed with climate over geologic time and because climatic change will alter patterns of precipitation and evapotranspiration and landscape hydrology, spatial patterns of soil salinity and alkalinity will shift (Schofield & Kirkby, 2003). In fact, patterns of salinity respond to changes in rainfall in a matter of years (Schagerl, 2016).
Local-scale controls are associated with topography, parent material, and groundwater recharge and discharge areas (Gile et al., 2007). Salt lakes and/or deposits can form in endorheic topographic positions under semi-arid and arid climates (e.g., Nasri, Bouhlila, Saaltink, & Gamazo, 2015; Schagerl, 2016; Schofield et al., 2001). Hydrogeological connectivity between leachable source rocks and saline lakes controls the potential of salt lake systems to form economically viable mineral deposits (Mernagh et al., 2016). Groundwater recharge areas, generally in highlands, are well drained and do not accumulate salts.
Transient salinization, sodification, and alkalinization are seasonally dependent and fluctuate with depth (Fig. 1) as a function of the penetration of the wetting front associated with precipitation, evapotranspiration, and the time interval between rainfall events.
Tree clearing disrupts the hydrological balance of landscapes (George, Nulsen, Ferdowsian, & Raper, 1999; Runyan & D’Odorico, 2010; Williamson, 1986); it not only reduces the amount of evapotranspiration but also increases the amount of water that drains through soil and recharges groundwater systems (Fig. 2).
Rising groundwater tables due to increased drainage past the root zone after tree clearing can bring salts within < 2 m of the soil surface where evaporative pumping can then lead to the accumulation of salts at the soil surface. The salt can be remobilized from the subsoil where salt bulges form below the tree rooting zone or come from deeper confined saline aquifers that are leaking upward due to the increased pressure from the increased recharge—in this case the recharge area can be very distant from the saline discharge area. These hydrological processes lead not only to soil salinization but also to salinization of streams, rivers, and reservoirs (Ghassemi et al., 1995).
Formation of Soil Carbonates and Alkalinity
A soil indicates the presence of large amounts of carbonates. The carbonates can be precipitated from soil solution (Table 1, equation 1 in reverse) associated with alteration haloes or weathering zones around ultramafic rocks (e.g., Anand & Paine, 2002), be present in groundwater, be formed from biological processes, respiration and decay of organic matter (Quade, Chivas, & McCulloch, 1995; Table 1, equations 6, 7, and 9), or be inherited directly from calcareous parent material (Gile et al., 2007; Wilford et al., 2015).
Erosional and depositional processes along soil catenary sequences in South Australia control the lateral distribution and morphology of regolith carbonate (e.g., Hill, McQueen, & Foster, 1999). Topographic position influences the local soil microclimate, the moisture regime, and biological activity, which in turn can control salt and regolith carbonate distribution (e.g., Gile et al., 2007; Harrison, Hendrickx, Muldavin, McMahon, & Wardell, 2003). Valley calcretes are common where the axis of stream or paleochannel systems and groundwater discharge zones become saturated with respect to carbonate minerals (Anand & Paine, 2002; Gile et al., 2007) in the seepage zone in Figure 1.
Carbonate minerals are generally calcite , dolomite , and magnesite in calcareous soils, and and/or in alkali or sodic soils. Soil alkalinity develops because some , when in soil solution, dissociates into and and reacts with water:
Carbonic acid, , is poorly soluble, and the equilibrium reaction favors water, , and (carbon dioxide gas that escapes into the atmosphere). The remaining dissociated sodium hydroxide produces high pH values.
In layers where calcium carbonate has accumulated during pedogenesis, replacement of by on the exchange complex—for example, if a -rich irrigation water is applied—generates sodium bicarbonate and/or carbonate ( and/or ), increasing the soil pH above 9.
In addition to the toxicity of carbonate and bicarbonate species for plants (e.g., Yang, Guo, & Shi, 2010), high pH also leads to Fe, manganese (Mn), copper (Cu), zinc (Zn), and phosphorus (P) deficiency (Naidu & Rengasamy, 1993). Phytotoxicity of aluminum, present as in soils with , has also been shown (Brautigan, Rengasamy, & Chittleborough, 2012).
Sulfur cycling at the land–water–atmosphere interface is dynamic, especially in acid saline environments. Sulfides, such as pyrite, are present in igneous and metamorphic rocks, sediments, and peat deposits. Sulfate minerals include gypsum, anhydrite, jarosite, natrojarosite, alunite, and several other minerals that form in soils, grow diagenetically from shallow groundwaters, or precipitate directly from lake waters (Benison & Bowen, 2013; Long, Lyons, & Hines, 2009). In semi-arid environments, these sedimentary S-minerals are subjected to physical reworking by wind erosion during the dry season and chemical dissolution and reprecipitation in the wet season.
Microbial activity can influence saline soil chemistry and vice versa, particularly via the sulfur cycle (Benison & Bowen, 2013; Miletto et al., 2008; Mormile, Hong, & Benison, 2009; Whittig & Janitsky, 1963; Wolicka & Jarzynowska, 2012). Both S-oxidizing and S-reducing bacteria are involved, and therefore S-bearing minerals include both sulfides and sulfates. Unique microbial populations exist in hypersaline environments (Friedrich, Rother, Bardischewsky, Quentmeier, & Fischer, 2001; Oren, 2002), especially the acidic lakes in Australia (Mormile et al., 2009). Microbial oxidation of sulfide-bearing rocks and sediments is responsible for the occurrence of sulfuric acid in groundwater and the occurrence of inland acid sulfate soils (e.g., Appleyard, Wong, Willis-Jones, Angeloni, & Watkins, 2004; Fitzpatrick et al., 1996, 1999; Hall, Baldwin, Rees, & Richardson, 2006). Sulfate-reducing bacteria are a group of anaerobic heterotrophic microorganisms that mineralize organic matter in extreme environmental conditions, including high salinity and acidity, and participate in geological processes such as mineral precipitation and ore formation (Table 1, equations 9 and 10). Active S-reducing bacterial communities decreased the salinity of their environment by as much as 50% (Wolicka & Jarzynowska, 2012).
Weathering and Marine Influence
The sources of ions including carbonate and sulfate are weathering of rock and soil parent material, atmospheric deposition, and groundwater. Weathering occurs through the action of water and dissolved on minerals. Rain in equilibrium with the atmospheric contains a small amount of carbonic acid and has a pH value of about 5.7 (Bricker et al., 1994). Soil solutions commonly have higher concentrations of than rainwater as a consequence of production by respiration and microbial decomposition of organic material (Table 1, equation 7). The major anion produced by carbonic acid weathering is (Table 1, equations 1, 3, and 5). The rate of weathering is enhanced if the water contains organic acids (e.g., carboxylic acids; Table 1, equations 6 and 7) or mineral acids (e.g., , , ; Table 1, equation 8). Congruent weathering reactions, such as the dissolution of limestone or quartz arenite, dissolve the rock mineral completely, leaving no solid residual (Table 1, equations 1 and 2), whereas incongruent dissolution produces clay minerals such as kaolinite as a residual byproduct (Table 1, equations 3–5).
Strontium isotope studies in regolith carbonates throughout the world indicate a predominantly marine source for strontium and, by inference, Ca (Capo & Chadwick, 1999; Chiquet, Michard, Nahon, & Hamelin, 1999; Dart, Barovich, Chittleborough, & Hill, 2007; Naiman, Quade, & Patchett, 2000; Quade, Chivas, & McCulloch, 1995). Mg and Cl can also be of marine origin (Keywood et al., 1997). Playas and salt lakes can be inland dust sources of salts (Aryal, Kandel, Acharya, Chong, & Beecham, 2012; Naiman et al., 2000; Schofield, Thomas, & Kirkby, 2001). In Australia and other dry regions, eolian deposition is an important soil-forming process. Prevailing wind patterns and distance from the coast are thus important factors that influence soil salt chemistry.
Evaporative concentration processes eventually cause the precipitation of very soluble minerals. The mineral equilibria models that prevail in neutral to alkaline brine environments have been developed by Hardie and Eugster (1970) and Eugster and Hardie (1978), building on earlier work by Garrels and Mackenzie. These were applied to soils by van Beek and van Breemen (1973) and Al-Droubi, Fritz, Gac, and Tardy (1980). The Long-Lyons-Hines (2009) model (Fig. 3) introduced a low-pH pathway for moderately acid brines and their minerals using data from Lake Tyrrell in Victoria, Australia, and may need to be revised in light of data for even more acid saline lakes in Western Australia (Benison & Lowenstein, 2015).
Lake Tyrrell is a terminal lake in a closed surficial drainage (endorheic) basin, but it is connected to underlying groundwater systems (Long, Lyons, & Hines, 2009; Mernagh et al., 2016). It sits in a calcareous soil landscape where the high initial composition of water reflects a marine influence and concentration (Fig. 3, far left). Soil calcrete formation in recharge areas removes Ca as well as . Ferrolysis, a repeated process of alternating oxidation-reduction of dissolved Fe in wet-dry conditions that destroys clay minerals (Brinkman, 1970):
generates acidity along flow path until the spring zone area. Microbial oxidation of pyrite present in underlying sediments:
produces more acidity. During acidification, any remaining is removed. Gypsum precipitation removes more Ca, and the saline water becomes relatively enriched in Mg as well as in Al and Fe. Evaporation of the water results in precipitation of Fe-oxides, alunite:
and further enhances acidity. The resulting composition of the brine in the spring zone is a solution with low pH and high amounts of Al and Fe (Long et al., 2009).
In contrast to Western Australian saline lakes, Lake Magadi in southern Kenya (Fig. 3, right) and other East African soda lakes are young systems. Their genesis and chemistry are highly linked to active tectonic forces and volcanic bedrocks (Schagerl, 2016). The basaltic rocks are relatively rich in calcium, whereas the more felsic rocks are dominated by sodium-rich silicate minerals and abundant K-feldspar. Quartz is absent in many of the volcanic rocks. Carbonatite ashes composed of nahcolite , trona , sylvite , halite, kalicinite , and villiaumite (NaF) are released during volcanic eruptions (Deocampo & Renaut, 2016). These deposits weather rapidly to produce secondary nahcolite, trona, thermonatrite , pirssonite , gaylussite , and calcite; much of this material is completely soluble, thus releasing alkali elements and carbonate rapidly into surface water and soils. The impact on the alkalinity and pH of these landscapes is so rapid and strong that it mimics the effect of arid evaporative processes (Deocampo & Renaut, 2016). Weathering reactions in upslope soils and groundwater environments (and atmospheric input) dominate the geochemistry of dilute inflow to the soda lakes. These reactions produce initial solute ratios that determine the subsequent geochemistry of evolved brines; for example, soda lakes may develop only in basins where the dilute inflow has (Fig. 3, right). This condition is favored by silicate weathering, such as the idealized hydrolysis reaction describing the weathering of sodium-rich plagioclase feldspar (Table 1, equation 4), or reactions with water and as carbonic acid (Table 1, equations 3 and 5). The weathering of these feldspars releases hydroxyl or ions, thus raising the pH and increasing the alkalinity of soil solutions.
Coastal Soil Salinity
Salt marshes are coastal ecosystems affected by salinity. Tidal salt marshes typically consist of high, middle, and low marsh zones (Adam, 1990). There is a soil salinity maximum near the high marsh zone. One of the main reasons for this salinity increase is the decreased duration of the tidal inundation, which allows evapotranspiration to concentrate pore water salinity and salt to accumulate. Thus, elevation plays an important role in the structure and function of salt marsh ecosystems because it determines inundation frequency and duration of tides. Soil hydraulic conductivity is as important a factor as ET and temperature in determining soil salinity accumulation in a marsh (Wang, Hsieh, Harwell, & Huang, 2007). Higher hydraulic conductivity results in lower soil salinity because water rapidly moves out of the soil, flushing the soil and minimizing salt buildup.
Widespread coastal salinization can be anticipated as a result of sea level rise under global warming. For example, in the Northern Territory of Australia, changes in the form and position of the coastline over the past 50 years have been paralleled by an extension of tidal creeks and the intrusion of salt water, resulting in the loss of 17,000 ha of freshwater wetlands in the Lower Mary River plains (Mulrennan & Woodroffe, 1998). Saline sulfidic soils are usually associated with tidal salt marshes and mangroves, and complex soil chemical changes are expected to result from salinization due to sea level rise.
Table 1. Weathering Reactions That Can Lead to Accumulation of Salts Under Arid Climatic Conditions
Incongruent reactions that produce kaolinite
Albite weathering by hydrolysis
Albite weathering by carbonic acid
Albite weathering by oxalic (dicarboxylic) acid
Oxalate converted to by microbial action
Albite weathering by sulfuric acid
Sulfate reduction by bacteria
Precipitation of metal sulfide ore mineral
Notes: Ebelmen-Urey reaction is from Berner (2012), equations 1–8 from Bricker et al. (1994), equations 9 and 10 from Webb et al. (1998). Chemical equilibrium reactions are those with =; theoretically they are reversible. The Ebelmen-Urey reaction is unidirectional and represents the long-term continental weathering of silicates and formation of marine carbonates and the short-term soil weathering and precipitation of pedogenic carbonates in arid and semi-arid conditions.
Land management can be a cause of soil salinization and/or of alkalinization. Tree clearing, whose consequent role in salinization was discussed above in the section “Climate and the Role of Water Movement,” is a prerequisite for agriculture and is recommended to stimulate pasture growth in Australian rangelands (McIvor & Monypenny, 1995). Irrigation is another land management practice that can result in soil salinization. The SAR and EC are also used to classify the quality of water for irrigation (Richards, 1954; see the article “Classification and Mitigation of Soil Salinization”.)
Maintaining a desirable stable soil structure with good permeability depends on controlling the flocculation-dispersion behavior of the soil clay fraction. To use saline-sodic soils for cropping, irrigation water with a high electrolyte concentration and large amounts of and is applied to remove from the exchange complex without initially changing the electrolyte concentration of the soil solution; then, once the divalent cations are the dominant ones on the exchange complex, the soil can be leached with water of lower electrolyte concentration to remove the excess salts (Richards, 1954). Gypsum or can be applied to remediate sodic soils with low EC by replacing the on the exchange complex with . is then leached out as a soluble salt, or NaCl. The addition of gypsum or also increases permeability by increasing the electrolyte concentration.
A plant-based solution for remediating calcareous saline-sodic soils has been proposed by Qadir, Steffens, Yan, and Schubert (2003) that is reportedly as successful in reducing levels as amendment with gypsum but cheaper. It relies on planting an alkaline/ tolerant crop, alfalfa (Medicago sativa L.), to enhance dissolution of calcite in the root zone and release to exchange with . In the phytoremediation treatment, the removed from soil through leaching, rather than taken up by plants, was the principal cause of reduction in salinity and sodicity (Qadir et al., 2003). Similar results have been reported using honeysuckle (Lonicera japonica Thunb.) in China (Yan, Xu, Zhao, Shan, & Chen, 2016).
Sulfuric acid can also be used for treating alkaline irrigation water and calcareous sodic soils (Miyamoto & Stroehlein, 1986): Applied directly to calcareous sodic soils, sulfuric acid dissolves and at solubilizes P, Fe, and Al, all of which help increase water infiltration. The amount of exchangeable Na removed by the sulfuric acid is approximately equal to that removed by the chemically equivalent rates of gypsum in moderately Na-affected soils. Sulfuric acid provides faster movement of leaching water than does the chemically equivalent rate of gypsum in severely Na-affected calcareous soils. When applied to moderately Na-affected soils, can provide faster leaching than applied at chemically equivalent rates.
The effectiveness of sulfuric acid as an amendment to improve water infiltration depends upon the actual chemical properties of the soils and water and the application methods (Miyamoto & Stroehlein, 1986): when injected into sprinkler lines (a closed water system), the sulfuric acid reacts with and to form carbonic acid and reduces the pH of the water below 7, as long as application rates are less than equivalent concentrations of . Upon sprinkling, the carbonic acid decomposes to , thus leading to Ca precipitation and limited effects on sodicity and water infiltration. When applied to open-ditch flow, sulfuric acid removes and , prevents Ca precipitation at , and may solubilize Ca from soil carbonates. Resulting effects are to maintain or reduce sodicity and to increase electrolyte concentrations, all of which contribute to increasing infiltration of sodic irrigation waters. Water-run application of , a common post-planting fertilization method in row crops, causes precipitation of Ca and induces sodium hazards. Sulfuric acid applied to such water neutralizes alkalinity and minimizes or prevents Ca precipitation and associated infiltration reduction.
Numerical models that account for variations in water content, chemical composition of soil and irrigation water, temperature, and concentrations in the soil are now used for irrigation scheduling in arid and semi-arid areas. The major chemical variables that are considered are Ca, Mg, Na, K, , Cl, , alkalinity, and . For example, the UNSATCHEM-2d model accounts for equilibrium chemical reactions between these components, such as complexation, cation exchange, and precipitation-dissolution of calcite, dolomite, gypsum, hydromagnesite, and nesquehonite (Šimůnek & Suarez, 1994). Because the ionic strength of soil solutions can reach high values, both modified Debye-Hückel and Pitzer expressions have been incorporated into the UNSATCHEM-2d model to calculate single ion activities (Šimůnek & Suarez, 1994). Most recently, Mau and Porporato (2016) have developed a control theory approach to reclaim sodic and saline-sodic soils by irrigation more rapidly.
The fate of the sulfate applied to remediate sodic soils has not been examined much beyond the point of application, but the S cycle is potentially impacted by the introduction of large quantities of gypsum or sulfuric acid into agricultural landscapes. If this sulfate eventually accumulates in wetlands and is reduced to sulfide, it may increase the area of inland potential acid-sulfate soils (Hall et al., 2006).
Case Study: Mt Lofty Ranges, South Australia—A Mediterranean Climatic Region (>600 mm Rain per Annum, Winter Rainfall Maximum)/Koppen Climate Class: Grassland, Warm, Summer Drought
Fitzpatrick et al. (1996, 1999) used soil-landscape process understanding with remotely sensed and point data, within a GIS framework, to produce a regional scale assessment of drainage conditions, salinity, and soil acidity/alkalinity for ~80 km2 in the Mt Lofty Ranges near the city of Adelaide in South Australia. They found saline sulfidic soils forming in waterlogged low-lying areas as a consequence of dryland salinity (Figs. 4–6). These unusual soils are due to contemporary weathering of pyrite lenses in the underlying rock, intercepted by two water-flow systems: a rising saline, sulfate-rich groundwater table responding to land clearing since European settlement, and the seasonal discharge of fresh water via a perched water table (Fig. 4). Three main processes control the formation of these saline sulfidic soils: (1) poor drainage, waterlogging, and the development of saline conditions throughout the solum, (2) the accumulation of organic matter from which Fe- and S-reducing/oxidizing bacteria derive their energy, and (3) a continuous supply of Fe and S in groundwater aquifers. A decade of detailed pedological, mineralogical, hydrological, and physicochemical investigations culminated in a conceptual framework to explain the formation of these saline sulfidic soils in the context of the overall processes shaping pedogenesis in the region (Fritsch & Fitzpatrick, 1994). Similar processes may be prevalent in wetlands throughout the Murray basin in Australia (Hall et al., 2006). Currently these soils have an alkaline pH, typical of anoxic conditions when sulfate-reducing bacteria thrive (Table 1, equation 9), but if they were drained, the sulfide would oxidize upon exposure to air and these soils would become acidic (inland potential acid sulfate soils).
This article has described the major chemical processes in salt-affected landscapes and causes of salinity, sodicity, and alkalinity in soils. Weathering of soil parent material under strongly seasonally contrasted climate, microbial activity, mineral equilibria, and ion exchange reactions controls the chemical composition of soil and water. Salinity is naturally present in the environment and presents major challenges to agriculture. Patterns of soil salinity can respond to changes in rainfall in a matter of years. But natural terrestrial ecosystems have evolved that are adapted to the wide range of saline chemistries observed; they are dynamic systems. Future challenges will be to understand how organisms have adapted to extreme saline, sodic, and alkaline conditions that exert multiple stresses on them concurrently—this may be the key to breeding salt-tolerant crops for farming on saline land—and to preserve some of the ecosystems in which they thrive.
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