SCOPE 51 - Biogeochemistry of Small Catchments

5.1 SOIL CHEMICAL REACTIONS

5.1.1 INTRODUCTION

Biogeochemical processes in the terrestrial environment dominate the hydrochemical response of small catchments, because streamwater is largely made up of drainage water from soils. Biogeochemical processes can be categorized into three major groups (Table 5.1.; cf. van Breemen et al., 1983):

1. Biochemical processes, including interactions between biota and the atmosphere (e.g. photosynthesis, respiration, N₂ fixation), and interactions between biota and soil solution (e.g. assimilation and mineralization).
2. Geochemical and soil chemical processes, including interactions between solution and the soil solid phase (e.g. cation exchange, adsorption, chemical weathering).
3. Chemical reactions in solution (e.g. hydrolysis, complexation reactions) or between solution and atmosphere (e.g. degassing of CO₂).

Processes from all three categories modify the chemical composition of infiltration water. For all major solutes the quantitative importance of individual biogeochemical processes has been estimated from input-output budgets and net assimilation rates (e.g. Driscoll and Likens, 1982; van Breemen et al., 1984; Nilsson, 1985; Binkley and Richter, 1987; Lelong et al., 1988).

In this chapter we will focus on soil chemical reactions (i.e. categories 2 and 3), and how they may affect concentrations of macro-solutes in streamwater. Besides a brief presentation of the theory of the dominant soil chemical processes and the spatial and temporal patterns of soil chemical reactions and parameters, some examples of anthropogenic impacts on soil chemistry and subsequent recovery of the soils will be discussed. This chapter will conclude with a method section, dealing with sampling and analysis of soils and soil solutions. Aspects of the chemistry of trace metals in soils will be discussed in Chapter 13.

As indicated in Table 5.1, several biogeochemical processes involve the transfer of H⁺ ions, thus affecting the acid-base chemistry of soils and soil water. Net H⁺ (proton) transfer may be calculated from quantitative estimates of individual chemical processes. By accounting for all proton sources and sinks a proton budget may be constructed. Proton budgets, which express the relative importance of all major sources and sinks of acidity, have been used extensively in acidification research (Driscoll and Likens, 1982; van Breemen et al., 1983). Major sources of protons include CO₂ dissolution in water, cation assimilation, nitrification and atmospheric acid deposition, while major sinks are cation exchange and chemical weathering.

Table 5.1 Reaction equations of H⁺ transfer processes and related processes involving biota (after van Breemen et al., 1983)

5.1.2 REACTIONS INVOLVING INORGANIC CARBON

In soil solutions dissolved inorganic carbon (DIC) is abundant, and consists of H₂CO₃* (CO₂(aq) + H₂CO₃), HCO₃^- and CO₃^2-. The distribution of DIC species in water can be described by equilibrium relationships, where the H₂CO₃* activity is controlled by the partial pressure of CO₂ in the atmosphere (pCO₂):

where brackets indicate activity (mol l^-1), and pCO₂ is in atm.

Dissociation of carbonic acid depends on pH and can be described as (Bolt and Bruggenwert, 1976):                       

For systems open to the atmosphere, pCO₂ is c. 3x10^-4 atm. However, in soils pCO₂ (ranging from 10^-2 to 10^-1 atm; Bolt and Bruggenwert, 1976) is generally higher, due to respiration and oxidation of below-ground organic matter. Consequently, DIC concentrations tend to be higher in soil solutions than in surface water. Degassing of CO₂ is common when soil water emerges (Reuss and Johnson, 1986).

Carbon dioxide, dissolved in soil water, may react with minerals (including feldspars and calcite) according to:

generally resulting in soil solution pH values well above 6. Most natural freshwaters are in this carbonic acid buffer range (Stumm and Morgan, 1981).

The presence of dissociated carbonic acid in water gives rise to alkalinity (Alk), where Alk is the equivalent sum of bases that are titratable with strong acid:

with brackets indicating concentrations in mol l^-1. Alkalinity is also known as the acid neutralizing capacity (ANC). The equivalence point of the acidimetric titration (around pH 4.5) represents an approximate threshold below which most life  processes in natural waters are seriously impaired. Thus alkalinity is a convenient measure for estimating the maximum capacity of a natural water to neutralize acidity without permitting extreme disturbance of biological activities in the water (Stumm and Morgan, 1981).

In very dilute natural solutions (e.g. in acidic soils) DIC and therefore [Alk] in Equation (5.5) are low. In such systems additional protolytic systems (of which hydrolysed Al compounds and natural weak organic acids are the most prominent), may contribute to alkalinity. Assuming that at the equivalence point of the alkalinity titration Al is present as Al(OH)²⁺ (Sullivan et al., 1989) the definition of alkalinity becomes:

with [RCOO^-] representing the concentration of organic anions.

Alkalinity is a conservative parameter, i.e. it is pressure and temperature independent. For example, degassing of CO₂ results in the removal of equivalent amounts of H⁺ and HCO₃^- from solution, thus causing no change in [Alk] in Equations (5.5) and (5.6). By contrast, degassing of CO₂ may result in a significant increase in solution pH, particularly in solutions with positive alkalinity (Reuss and Johnson, 1986; Suarez, 1987). Degassing in solutions with a negative alkalinity (i.e. solutions where strong mineral acids dominate) will have little or no effect on pH, however.

5.1.3 REACTIONS INVOLVING ORGANIC CARBON

Soil organic matter (SOM) can be subdivided into non-humified and humified material. Non-humified substances are not or are only slightly altered after decay of tissue from living organisms and include, e.g. carbohydrates, amino acids, protein, lignin, hormones and low molecular weight organic acids (Tan, 1986). Humified substances are decomposition products of non-humified constituents and include complex compounds such as humin, fulvic acid (FA), hymatomelanic acid, humic acid (HA) and their hydroxybenzoic acid derivatives (Tan, 1986).

The concentrations of non-humified organic acids are generally low and many of these acids can only be detected by thin layer or gas chromatography. Nevertheless, with their rapid turnover, low molecular weight organic acids may play a significant role in mineral weathering. In most soils the contents of HA and FA are considerably higher than those of the non-humified organic acids. The major reason for the importance of FA and HA in soil chemistry is the presence and position of functional groups (particularly carboxyl and phenolic hydroxyl groups), which make FA and HA effective in cation exchange and complexation reactions (Tan, 1986). Charge characteristics of humic substances depend upon the extent of dissociation of the functional groups. At pH < 3 HA and FA behave as uncharged polymers, whereas at pH > 3 they behave as negatively charged polyelectrolytes, due to dissociation of carboxyl groups (3 < pH < 9) and phenolic hydroxyl groups (pH > 9) (Tate and Theng, 1980).

Fulvic and humic acids may form metal complexes of high stability through chelation, where metal ions are bound as bidentate e.g. by carboxyl and phenolic hydroxyl groups. Chelation promotes the dissolution of metals from soil minerals. For example, in podzols chelation of Al and Fe in the surface mineral soil and subsequent leaching eventually result in Al and Fe depleted eluvial (E) horizons. The solubility of chelates decreases with increasing sesquioxide content, however, thus resulting in precipitation in the illuvial (podzol B) horizon.

In soils, DOC concentrations decrease sharply from several tens of mg l^-1 in the O and E horizons to only a few mg l^-1 in the mineral soil (e.g. McDowell and Wood, 1984; Cronan and Aiken, 1985). Besides physicochemical sorption and precipitation, this decrease in DOC is likely due to mineralization.

As indicated in Section 5.1.2, DOC may contribute to the alkalinity of soil solutions and surface waters. On the average, organic acids in natural water contribute 5 to 10 µeq of anionic charge per mg DOC. To estimate the anionic charge associated with DOC in individual samples the empirical model by Oliver et al. (1983) has frequently been used (e.g. Driscoll et al., 1989a). Alternatively, Cronan and Aiken (1985) and Schecher and Driscoll (1987) used analogues of monoprotic, diprotic and triprotic acids with optimized dissociation constants. Recently, Tipping et al. (1988) proposed a deterministic model which treats organic acids as macroions, where interactions are modelled in terms of discrete binding sites. This model accounts for electrostatic (macroionic) effects which depend upon molecular charge (determined by the bound cations) and ionic strength.

Because soil organic matter is a major source of negative charge in many soils it acts as an important cation exchanger, which may represent a significant chemical buffer (e.g. the forest floor; James and Riha, 1986). In addition, soil organic matter may complex trace metals (e.g. Al), thus reducing phytotoxic effects (e.g. Bloom et al., 1979). Soil organic matter is also important as a store for N and S, which may be liberated slowly upon decay.

5.1.4 CATION EXCHANGE REACTIONS

In many soils the solid phase carries a net negative surface charge. For clay minerals this charge is a result of isomorphous substitution, where structural cations of higher valence are replaced by cations of lower valence (e.g. Si⁴⁺ is replaced by Al³⁺). As a result, clay minerals have a permanent negative surface charge.

Oxides and SOM also have charged surfaces: With increasing pH (i.e. increasing activity of OH^- ions) H⁺ is dissociated from oxide surfaces or from organic functional groups, thus resulting in a negative charge. Because dissociation increases with pH and ionic strength, such charge is termed variable charge.

At low pH values oxides may bind H⁺, which results in a positive surface charge. The pH values at which positively charged groups quantitatively equal negatively charged groups (i.e. the net surface charge is zero) are called zero point of charge (ZPC).

In soils the overall electroneutrality is maintained by an excess of electrostatically attracted counterions in proximity to the charged surface. In the case of negatively charged surfaces a diffuse double layer will result where cations are in excess of anions. The excess ions, termed exchangeable cations (anions) may, by definition, be exchanged with neutral salts. The quantity of exchangeable cations (in meq kg^-1 dry soil) is defined as the cation exchange capacity (CEC). Similarly the total amount of exchangeable anions is defined as the anion exchange capacity (AEC). In most soils the AEC is quantitatively far less important than the CEC.

Exchangeable cations are available to plants, for example through exchange with H⁺ liberated by the roots. Exchange reactions are also responsible for the retention of freshly introduced cations into the soil solution. In this way the CEC gives the soil a buffering capacity, which may slow down the leaching of nutrient cations and positively charged pollutants.

Since cation exchange is a relatively fast process it modifies the chemical composition of infiltration water. The exchange reaction of cations M and N, with a charge of m⁺ and n⁺, respectively, may be represented as:

Ads-Nm + nMm+ Ads-Mn + mNn+

(5.7)

Mathematically this reaction has been described by the Gapon equation:

  with the left-hand side being the ratio of adsorbed M over N (both in meq per mass unit exchanger). The right-hand side contains the reduced activity ratio, where the cation activities are raised to a power equal to the reciprocal of their valence. K_NMg in this equation represents the (Gapon) selectivity constant, which should be constant over a wide range of conditions. The selectivity constant differs from one, because small size (hydrated) ions are generally preferred over large ones (high K_NM), due to their smaller distance of approach of the charged surface. Also the surface structure of the exchanger may affect the selectivity constant (e.g. in case of a porous exchanger, where the binding may be determined by the "naked" cation; Bolt and Bruggenwert, 1976).

Alternative mathematical descriptions of cation exchange reactions have been proposed (e.g. the Vanselow equation), but most frequently used in hydrochemical models for small catchments is the Gaines- Thomas equation:

Also here Ads(M) and Ads(N) are adsorbed amounts of M^m+ and Nⁿ⁺ in meq , per mass unit exchanger. The three different formulations have been compared by Reuss (1983).

Note that all three exchange equations were developed for a constant capacity (permanent charge) exchanger. This implies that if the change in cation exchange complex composition is small compared to the total store of exchangeable cations (which is generally the case, particularly in the short run) the adsorbed ion ratio Ads(N)^m/ Ads(M)ⁿ remains constant and hence the solution ratio (Nⁿ⁺)^m/(M^m+)ⁿ also remains constant. For Al³+-Na⁺ exchange this would suggest a cubic relationship in solution, whereas for Ca²⁺-Na⁺ exchange this relationship is expected to be squared (Reuss and Johnson, 1986; Neal et al., 1990).

However, in situ soil solution data from a great number of sites in Wales, Scotland and Norway show that cation concentrations are poorly predicted by the cation exchange equations given above (Neal et al., 1990). Possibly, application of a variable-charge instead of a permanent-charge exchanger model improves the quality of the predictions. A variable-charge cation exchange model (similar to the one already discussed for DOC) was recently proposed by Tipping and Hurley (1988).

5.1.5 SULPHATE TRANSPORT AND ADSORPTION

Oxidation of reduced S (e.g. pyrite), and atmospheric deposition of S compounds are the predominant sources of sulphate in most natural soils. Atmospheric S may be of anthropogenic (e.g. SO₂ due to burning of fossil fuels), and of natural origin (e.g. SO₂ due to volcanic activity, SO₄^2- in sea spray aerosol, gaseous biogenic S). Most S inputs are eventually transformed into H₂SO₄, and may acidify soil and water. Only S in seaspray is accompanied by base cations and is not acidifying. In coastal areas the acidifying potential of the atmospherically derived sulphate (often termed excess SO₄^2-) is estimated by subtracting seaspray S from the total S deposition.

The acidifying potential of H₂SO₄ in soils is largely determined by sulphate adsorption. Particularly, Al and Fe oxide surfaces have been reported to be major adsorption sites for soil solution SO₄ (Fuller et al., 1987; Mitchell et al., 1989). Adsorption of sulphate results in displacement of -OH^o ligands from oxides: 

Oxide-OHo + SO42- Oxide-SO4- + OH-

(5.10)

where OH^- is released and a cation exchange site is created (Rajan, 1978; Parfitt and Smart, 1978). Alternatively SO₄ adsorption may result in the displacement of an -OH₂⁺ ligand from the oxide:

Oxide-OH2+ + SO42- Oxide-SO4- + H2O

(5.11)

where water is released and the site changes from positively to negatively charged. Both mechanisms result in an increased cation exchange capacity, so that adsorption of sulphate and cations occurs simultaneously (Mitchell et al., 1992). At low pH a simple ligand exchange has been shown where two adjacent -OH 1igands are replaced by one SO₄^2- (Rajan, 1978). Most studies indicate that SO₄ binds more weakly to oxide surfaces than PO₄, but much more strongly than NO₃ and Cl (Marsh et al., 1987).

Sulphate adsorption increases with decreasing pH (Chao et al., 1962; Nodvin et al., 1986b; Puller et al., 1987). This pH-dependent adsorption probably results from the amphoteric character of the adsorption sites on oxide surfaces, which become negatively charged upon deprotonation, thus causing anion repulsion. The adsorption of sulphate is negatively affected by the presence of organic matter, even in soils where the contents of Al and Fe oxides are high (e.g. podzol B horizons; Johnson and Todd, 1983). This reduction in sulphate adsorption has been hypothesized to be the result of blocking of the sites by organic matter.

Sulphate adsorption is generally considered a reversible process and is frequently described by a Langmuir isotherm (e.g. Cosby et al., 1985):

where Es is adsorbed SO₄ (eq kg^-1), E_mx maximum SO₄ adsorption capacity (eq kg^-1), and C the half saturation constant of the adsorption process (eq m^-3). Brackets indicate concentration (mol m^-3). Note that SO₄ adsorption in Equation (5.12) is not pH dependent. To what extent the assumption of reversibility of SO₄ adsorption is realistic will be discussed in Section 5.3.2. 

Recently Nodvin et al. (1986a) developed the initial mass isotherm to describe adsorption reactions of anions in soils. This isotherm essentially is a simple model partitioning the anion between solid phase and solution.

In some acidic forest soils of northwestern Europe jurbanite (AlOHSO₄.5H₂O) has been reported to control the activity of SO₄ and Al (Nilsson and Bergkvist, 1983; Khanna et al., 1987), according to:

where K_SO is the solubility product of jurbanite (10^-17.8; Nordstrom, 1982), and brackets indicate activity (mol l^-1).

Microbiological transformations may play a crucial role in S cycling, as illustrated by the general dominance of organic S forms in soils (carbon-bonded S and ester-bonded S; e.g. Mitchell et al., 1992). Particularly in catchments with significant amounts of wetlands, reduction of SO₄ to H₂S and FeS may be an additional important sink for SO₄ (Urban et al., 1987). Biologically mediated transformation processes of S will be discussed in more detail in Chapters 6 and 11.

5.1.6 PHOSPHATE TRANSPORT AND ADSORPTION

Phosphorus, often one of the growth-limiting nutrients, is generally tightly conserved in natural ecosystems (e.g. Wood et al., 1984). Living organisms contribute significantly to P retention through close coupling of decomposition of organic matter and assimilation. These biological controls on phosphate are dominant in the O horizon and the upper mineral soil layers.

Inorganic P is adsorbed on clay minerals and on Fe and Al oxides as an inner- sphere complex through ligand exchange with -OH surface groups (Goldberg and Sposito, 1985). Adsorption processes largely control soil solution P in horizons high in sesquioxides (e.g. B horizons). The bioavailability of P in such horizons depends on the stability of the oxide-P surface complexes and the presence of ligands that can replace P. Fox et al. (1990) found that the release of both inorganic p and Al from a spodic horizon increased in the presence of simple organic acids that form stable complexes with Al. Overall the Al and P release was found to increase exponentially with increasing Al stability constant (log K_Al).

5.1.7 REACTIONS INVOLVING ALUMINIUM

Although Al is one of the most abundant elements in soils, it is not very mobile at the pH values encountered in most soils (i.e. between 6 and 8). The solubility of Al increases in acidic soils (pH <6) and in the presence of complexing ligands (e.g. F^-, SO₄^2-, DOC; Driscoll, 1989). In naturally acidic soils dissolved Al is predominantly in an organically complexed form (David and Driscoll, 1984; Driscoll et al., 1985). Introduction of high loads of mineral acidity results in a significant mobilization of inorganic Al, so that it may become the dominant cation in solution (Mulder et al., 1989b). Dissolved Al is potentially toxic to terrestrial biota and has been hypothesized to be one of the major causes of the occurrence of widespread forest decline in central Europe (e.g. Ulrich et al., 1980).

Major sources of dissolved Al include a number of operationally defined pools, such as exchangeable Al, organically bound Al, amorphous Al oxide and crystalline Al oxide (e.g. gibbsite). Unfortunately, fluxes between these (physicochemically poorly defined) pools and solution are unknown. Consequently, our understanding of Al solubility controls in soils is limited and largely based on equilibrium assumptions with respect to a few Al-containing minerals, e.g. gibbsite (Al(OH)₃). Gibbsite equilibria may be formulated as:

where brackets represent activities. At 25°C the equilibrium constant (K_SO) for gibbsite ranges from 10^8.11 (synthetic gibbsite) to 10^10.80 (amorphous Al(OH)₃). However, several studies suggest that dissolved Al does not follow gibbsite equilibrium, often not even where gibbsite is present (McMahon and Neal, 1990). Recently, Dahlgren et al. (1989) concluded from column experiments that inter- layer Al may control the solubility of this metal in mineral soils. Soil solutions in organic-rich horizons are often undersaturated with respect to gibbsite, and dissolved Al may be controlled by cation exchange (e.g. Bloom et al., 1979; Cronan et al., 1986), or complexation reactions (Tipping et al., 1988; Mulder et al., 1989a). Some studies in European forest soils suggest an Al solubility control by jurbanite (Equation 5.13).

5.1.8 CHEMICAL WEATHERING

Together with atmospheric input, chemical (and physical) weathering of minerals constitutes the major source of base cations in catchments (Figure 5.1). Liberation of base cations through chemical weathering increases the alkalinity of percolating water (see also Section 5.1.2). The base cation weathering rate and thus the rate of alkalinity production determines the long-term resistance of a soil to acidification (defined as a decrease in pH and base saturation) and is often taken as a measure for the critical load of strong mineral acid to a soil. The critical load is defined here as the maximum strong acid input that can be neutralized without depleting the stores of exchangeable base cations (Nilsson and Grennfelt, 1988). Rates of chemical weathering have been estimated from a combination of input-output budgets and laboratory experiments (see also Chapters 4 and 14).

Figure 5.1 Schematic diagram of major base cation fluxes in catchments: W = supply rate of base cations through weathering; I = input rate of base cations from precipitation, dry deposition and mineralization of organic matter; U = biological uptake rate of base cations; L = leaching rate of base cations with drainage water (from Christophersen et al., 1990b; reproduced by permission of Elsevier Science Publishers BV).

5.1.9 NITROGEN TRANSPORT

Biochemical transformations of N in terrestrial ecosystems generally result in a rapid decrease of dissolved NH₄⁺ with depth (due to cation exchange, assimilation and nitrification). Nitrogen leaching, if occurring at all, is mostly in the form of NO₃^-. Biochemical aspects of N transformation processes will be further discussed in Chapters 6 and 11.

Nitrogen transformations also affect the acid-base chemistry of soils and soil solutions (see Table 5.1). Enhanced nitrification and leaching of NO₃^- (associated with a cation) results in progressive soil acidification and consequently in a gradual decline of soil pH. As indicated earlier, soil pH is a master variable for many soil chemical reactions.

5.2 SPATIAL AND TEMPORAL VARIABILITY OF THE SOIL

As discussed above, the chemical composition of soil solutions is affected by numerous chemical reactions. Thermodynamic and kinetic constraints involved may vary between soil types and even between soil horizons (layers), because of differences in chemical properties. Therefore, the distribution of soil types is an important factor in the hydrochemical response of catchments. In discussions of soil types we will use the FAO-UNESCO soil classification system (FAO-UNESCO, 1988). In addition to variations in space, soil chemical reactions may also vary temporarily, due to changes in temperature, moisture, biological activity, etc.

5.2.1 HORIZONS: VERTICAL VARIABILITY IN CHEMISTRY

Soil horizons (i.e. layers in the soil, matching a given set of criteria, e.g. FAO-UNESCO, 1988; Soil Survey Staff, 1975) develop as a result of long-term physical, chemical and biological processes. How a soil develops at a specific site depends on the relative importance of various soil-forming factors (i.e. climate, parent material, vegetation and fauna including man, topography, time).

Soil formation often involves transport of specific compounds resulting in depletion in some parts of the profile and enrichment in others. For example, due to the chelating power of percolating fulvic acids, Podzols have a strongly acidified, bleached Al- and Fe-depleted E horizon at the surface. At depth Fe- and Al- chelates precipitate, thus forming a spodic B, enriched in Fe, Al and organic C. At the surface, Podzols often develop a layer of acidic, slowly decomposing raw humus (O horizon). Clearly, the different chemical properties of horizons in Podzols (i.e. pH, CEC, and contents of Fe and Al oxides and soil organic matter) may lead to differences in chemical reactivity.

Other examples of soil formation, resulting in a vertical differentiation of chemical properties include the following:

1. Transport of clay minerals, giving rise to a clay-depleted horizon on top of a clay-enriched (argic B) horizon, as in Luvisols, Acrisols and Alisols.
2. Upward transport (capillary rise) of water and dissolved salts in areas with a high groundwater table and a precipitation deficit may lead to accumulation of salts at or near the surface (calcic, gypsic, natric-B horizon), as in Solonchaks and Solonetzes.
3. Intensive weathering and leaching of Si in the humid tropics may lead to a ferralic horizon, high in residual Fe- and Al-oxides and kaolinite (Ferralsols).
4. Peatlands (Histosols) may form as a result of accumulation ( of organic matter, due to incomplete decomposition of plant debris), rather than leaching.

Clearly, thermodynamic equilibria and kinetics of soil chemical reactions may differ per soil horizon. Particularly in catchments with shallow soils and a strongly fluctuating groundwater table, vertical gradients in soil solution composition may be reflected directly in the chemistry of streamwater, due to variable contributions of soil horizons to lateral transport (Mulder et al., 1990; Easthouse et al., 1992).

Some examples of differences in chemical constraints between various soil horizons are discussed below:

1. Adsorption of SO₄ and PO₄ in Podzols increases in the order Raw humus < E horizon < spodic B, because of the increase in oxide content.
2. Adsorption of SO₄ and PO₄ is generally higher in an argic B (of, e.g. Acrisols) than in a spodic B, due to the lower oxide and higher SOM content of the latter. In ferralic horizons (Ferralsols) the adsorption of SO₄ and PO₄ is even higher. In these variable charge horizons AEC is often even higher than CEC, so that NH₄⁺ and K⁺ are more mobile than, e.g. SO₄^2- and NO₃^- (Sollins et al., 1988).
3. In sandy Podzols the CEC, largely due to SOM, is generally highest in O horizons, where it represents a significant buffer capacity, and decreases sharply in the mineral soil (with a minor peak in the spodic B).
4. In the presence of strong acid the Al mobility in Podzols increases in the order O < E < B horizon.
5. In Podzols of pristine (i.e. unpolluted) areas the alkalinity contributed by the organic and carbonic acid systems changes with depth (Ugolini and Dahlgren, 1987). While organic acids dominate alkalinity in the surface horizons (B horizon and above), this gradually changes with depth into DIC dominated alkalinity (Ugolini and Dahlgren, 1987; Ugolini et al., 1988). This is due to the removal of DOC from solution (due to mineralization and sorption; see Section 5.1.3) and a simultaneous increase in the partial pressure of CO₂ (Bouten et al., 1984).

5.2.2 SOIL TYPES: LATERAL PATTERNS (CATENA)

In many catchments sequences of different soil types (catena) have developed in uniform parent material, due to variation in relief, drainage and vegetation. For example, soils with eluvial (leached) horizons generally develop on hill slopes and summits (e.g. Podzols, Luvisols, Acrisols, Ferralsols), while soils in depressions (e.g. adjacent to the stream) may exhibit gleyic properties or accumulate organic matter (Histosols). Consequently, water infiltrating in hill slopes may have intimate contact with several horizons of different soil types, before emerging into the stream.

In most current hydrochemical models the complexity of soil distribution patterns in catchments is simplified by representing soils by one or more vertically related "boxes" (e.g. surface soil and subsoil). Each of these boxes has "average" soil chemical properties. Catchments having Podzols, for example, on the hillslopes are generally represented by an upper O horizon and a lower B horizon box. Implicitly, chemical modification of laterally moving drainage water from these two horizons in the wetter soils adjacent to the stream is not considered.

However, field evidence suggests that chemical modification of solutions during transport in the terrestrial environment does occur and may be significant (Bishop et al., 1990; Billett and Cresser, 1992; Mulder et al., 1991). Simultaneously, soils in these wet zones may become chemically modified (e.g. increased stores of exchangeable Al, due to Al transport from up slope; Mulder et al., 1991). Ignoring the chemical potential of these wet zones at lower elevations seriously limits the predictive power of hydrochemical models.

Besides physical evidence, additional support for the modification of laterally transported water by gleysoils and Histosols has been obtained from mixing models. For example, Christophersen et al. (1990a) found that streamwater at Birkenes (Norway) could largely be explained as a mixture of solutions from upslope O and B horizons and gleysoils in the valley bottom.

5.2.3 TEMPORAL PA1TERNS IN SOIL CHEMISTRY

5.2.3.1 Soil solution

The soil solution composition of a given soil horizon may vary considerably through the year, due to (i) a variable composition of infiltration water, (ii) variability in climatic conditions, and (iii) variability in biological activity.

In a recent study conducted over a six year period van Breemen et al. (1988) observed strong seasonal changes in the soil solution chemistry of an oak-birch woodland in The Netherlands. Particularly in the growing season nitrification of extreme loadings of atmospheric NH₃ were found to significantly increase the leaching of Al, Ca and Mg. Further increases in concentration were due to evapotranspiration throughout summer, as is evident from Cl (van Breemen et al., 1988). In ecosystems which receive relatively little atmospheric N, mineralization and subsequent nitrification of soil organic N through summer may also cause a significant increase in soil solution NO₃ and Ca (e.g. Foster et al., 1989).

Due to prolonged leaching, long-term natural changes in soil solution composition may occur. However, these changes are often too slow to be recorded by most monitoring programmes, which were nearly all established quite recently. Relatively fast changes in soil solution composition (e.g. decreases in base cations) have been reported in connection with anthropogenic impacts on soils (including forest production, acid rain, etc.) and will be discussed in more detail in Section 5.3.

5.2.3.2 Soil solid phase

The contents of chemical compounds in the soil solid phase are generally high relative to the rates of leaching and accumulation. Therefore, natural changes in the chemical composition of soils are generally slow. Anthropogenic impacts may speed up leaching and accumulation of chemical compounds considerably, thus causing measurable changes in soil chemical properties over relatively short time spans (Section 5.3).

5.3 ANTHROPOGENIC IMPACT ON SOILS AND SUBSEQUENT RECOVERY

5.3.1 ANTHROPOGENIC IMPACT AND RESULTING CHANGES

The impacts of humans on soils are manifold, and are generally associated with economic activities. They may be intentional, aimed at improving the economic value of a soil (e.g. for crop production through fertilization, liming, irrigation, artificial drainage), or they may be unintentional and often undesirable (e.g. salinization of irrigated land, pollution by pesticides used for crop protection, pollution by heavy metals from applied sewage sludge). Increased soil acidification is also an undesirable side effect, caused mostly by atmospheric deposition of SO_x and NO_x (from burning of fossil fuels), by oxidation of reduced S compounds in mine spoils or in artificially drained acid sulphate soils, by forest clearcutting and by plantation of coniferous forest.

Strong mineral acids (e.g. from acid deposition, pyrite oxidation, nitrification after deforestation) may cause increased leaching of base cations and depletion of exchangeable pools (Christophersen et al., 1990b; Johnson and Todd, 1990). Consequently pH gradually decreases and mobilization of inorganic Al increases (van Breemen, 1976; Driscoll et al., 1989c; Mulder et al., 1989b). At prolonged high rates of acidification (e.g. acid atmospheric inputs of several kmol ha^-1 year^-1) even the soil Al buffer in the surface horizons may be depleted, thus resulting in decreased pH buffering and consequently in a further decline of soil pH (Mulder et al., 1989a; de Vries et al., 1989).

Evidence for a decrease in soil pH attributed to acid atmospheric deposition has been found in a number of countries, including Germany (Ulrich et al., 1980; Butzke, 1981), Sweden (Hallbäcken and Tamm, 1986; Berden et al., 1987), Norway (Dahl, 1988), and the United Kingdom (Billett et al., 1990). In some studies forest growth has also been found to result in a detectable decrease in base saturation and acidification of forest soils within decades, due to net uptake of base cations (Hallbäcken and Tamm, 1986; Johnson and Todd, 1990).

A decrease in soil solution pH may also cause increased protonation of organic and bicarbonate anions (decreased alkalinity). Krug and Prink (1983) hypothesized that acidic deposition changed the source of streamwater acidity (from organic acid to mineral acid) but not the extent. However, a full neutralization of mineral acidity is unlikely as it would imply that all H⁺ from strong acids are consumed in the alkalinity titration. Some studies suggest a decreased solubility of organic acids, due to increased acidic deposition (Marmorek et al., 1987; Driscoll et al., 1989a). By contrast, Wright (1989) concludes from an acid exclusion experiment in Norway that the observed decrease in streamwater SO₄ is not associated with a change in DOC. Nevertheless the dissociation of organic acids was found to increase considerably.

Soil acidification (reduced pH) may also result in increased sorption of anions in solution (e.g. sulphate and organic anions). Since anion adsorption is an acid-consuming process, this may decrease the acidification potential for H₂SO₄. By contrast, desorption in periods of falling anion concentrations (decreased acid inputs) may significantly delay the de-acidification of soil drainage water.

5.3.2 RECOVERY FROM SOIL DISTURBANCE

Recovery of acidified soils largely depends on the rate at which exchangeable acidity can be removed (i.e. the rate at which exchangeable base cations can be replenished). In a natural situation this depends on the rate of release of base cations through chemical weathering, the deposition rate of base cations from the atmosphere and redistribution of base cations in the profile by biological processes (Figure 5.1). Liming and fertilization may speed up the recovery considerably, but unintentional side effects may occur (e.g. changing rooting structure, changing soil microbiological community) having both additional chemical and hydrological consequences.

The rate of (chemical) recovery of soils also depends on the reversibility of SO₄ adsorption. This partly depends on whether SO₄ is held electrostatically or by a stronger molecular bond at oxide surfaces (Mitchell et al., 1992). Irreversibly adsorbed SO₄ will not redissolve when SO₄ concentrations decrease. By contrast, large pools of reversibly adsorbed SO₄ could considerably delay the response of surface waters to changes in S input. Field observations suggesting that SO₄ adsorption is largely irreversible include those from the Sudbury area in Canada (Keller and Pitblado, 1986), New Hampshire, USA (Driscoll et al., 1989b), New Jersey, USA (Morgan, 1990) and Norway (Christophersen et al., 1990b). At all these sites a rapid decline in stream SO₄ at decreasing SO₄ deposition rates is evident. Similar results were obtained for field manipulation studies (Wright et al., 1988) and soil column experiments in the laboratory (Dahlgren et al., 1990).

5.4 METHODS IN THE STUDY OF SOIL AND SOIL WATER

5.4.1 SOIL SURVEY, SAMPLING AND ANALYSIS

5.4.1.1 Soil survey

In a soil survey information about the spatial distribution of soil types within a catchment is obtained. First an appropriate soil legend has to be determined in the field. At the highest level this legend should have soil properties, which are believed to be most relevant for the hydrochemical response in catchments (e.g. acidic vs. neutral soils, and well-drained vs. poorly-drained soils).

 Soil sampling

After mapping, dominant soil types should be sampled with respect to both solid phase and solution. If possible, soils should be sampled according to horizon.

However, if boundaries between horizons are diffuse or if soils are only weakly developed, sampling may be done according to depth. In the latter case the thickness of the sampled layers may increase with depth (e.g. from a few cm at the surface to a few dm at depth). Since the variability in chemistry within one horizon may be significant, samples should be taken from several soil pits, randomly distributed within a soil type. The number of samples needed depends on the confidence limit for the standard error of the chemical parameter to be estimated (e.g. Johnson et al., 1990).

Soil samples are generally air-dried (sometimes freeze-dried) upon return in the laboratory, sieved (< 2 mm or fine earth fraction), homogenized (sample splitter), and stored in plastic containers until analysis. Although some soil chemical characteristics may change in the drying process (e.g. mineralization of organic N and S), drying is generally preferred over working with field-moist samples. The latter has severe drawbacks: (i) samples must be analysed directly (storage may cause significant mineralization, and nitrification); and (ii) samples are difficult to sieve and homogenize.

Physical parameters

Soil physical parameters that are routinely measured include texture and hydraulic conductivity (both saturated and non-saturated). For textural analysis air-dried and homogenized soil samples are pretreated to enhance dispersion. Pretreatment includes removal of cementing particles like CaCO₃ (by acidification with 1 M NaOAc), Fe-oxides (by reducing and complexing free Fe, using bicarbonate- buffered, Na-dithionite-citrate) and organic matter (by oxidation with H₂O₂). After dispersion (e.g. in Na-hexametaphosphate) various sand fractions (>50mm) are determined by sieving, while clay and silt fractions are obtained from sedimentation analysis (Gee and Bauder, 1986). The hydraulic conductivity is determined in undisturbed soil columns.

Chemical parameters

Here we will briefly discuss the most common soil chemical parameters determined in catchment studies. For details concerning methods and computational procedures reference is made to the extensive reviews in Klute (1986) and Page et al. ( 1986).

Routine chemical analyses include measurement of pH, CEC and exchangeable cations. Generally pH is determined in 1 : 5 (weight to volume) ratios in both water and neutral salt solutions (often 1N KCl). In acidic soils, high in organic matter (and thus CEC), pH(KCl) is lowest, due to the contribution of some exchangeable acidity. Note that the pH of the soil extract is generally not directly comparable to the in situ pH of the soil solution. In the most recent years measurement of the CEC with neutral salts (often with Ba, NH4 or both; e.g. Hendershot and Duquette, 1986; Amacher et al., 1990) has become common practice. With the same technique exchangeable cations may be determined. Further routine soil chemical analyses include the contents of CaCO₃, organic C, total N, total S, and free Fe and Al.

For specific purposes additional analyses may be needed, including contents of inorganic N (i.e. NH₄ and NO₃) as well as fractionation of free Al and Fe into exchangeable, organically complexed, amorphous and crystalline forms (e.g. Driscoll et al., 1985; Parfitt and Childs, 1988). Total S and P are often fractionated in various forms (Page et al., 1986). In particular cases adsorption characte ristics of soils are determined with respect to PO₄ (van der Zee, 1988) and SO₄ (Nodvin et al., 1986b; Singh et al., 1980).

5.4.2 Soil water and groundwater collection

Sampling soil solutions and groundwater remains a problem area. Different techniques may each sample a specific "type" of soil water (with a unique composition). In comparative studies, solutions collected at zero tension showed fewer signs of chemical interaction with the soil or were less concentrated by evapotranspiration than solutions collected with tension lysimeters (Cozzarelli et al., 1987; Joslin et al., 1987; Swistock et al., 1990). The difference between the zero tension and the tension Iysimeter solutions decreased with increasing depth in the soil. These results suggest that preferential flow is most important in the surface layers, where the macro-pore density (related to biological activity) is highest.

More recently Zabowski and Ugolini (1990) found significant (seasonal) differences between the chemical composition of lysimeter samples (extracted at 10 kPa) and centrifugates from the same horizon (extracted at 30 kPa and 300 kPa, respectively). These authors concluded that the lysimeter solutions were more representative for moving water, whereas centrifugation would be more appropriate for forest nutrition studies.

In addition to the poorly defined "type" of soil water that is being sampled, sampling procedure and equipment may induce a chemical change in the collected solution. For example, Suarez (1987) found that vacuum extractors could result in a significant increase in soil solution pH, due to degassing of CO₂. This effect may be corrected for (Suarez, 1987). Alternatively, Takkar et al. (1987) proposed the use of a closed tension lysimeter system, without exposure to the atmosphere, to sample soil solutions and preserve the dissolved inorganic carbon.

Although numerous types of suction lysimeters have been proposed, the material used for the construction of the collection device (alundum, ceramic, teflon, acrylic copolymer with internal nylon support) may have a pronounced influence on the composition of the sample (Litaor, 1988). From his review Litaor concluded that there is no single and simple solution to the collection of soil water. The design and properties of a soil solution sampler should be evaluated critically for each specific problem under study.

5.4.3 CHEMICAL ANALYSIS OF SOLUTIONS

A common analysis of soil solutions includes pH, and total concentrations of DOC, DIC, Si, K, Na, Ca, Mg, Al, NH₄, F, CI, NO₃, SO₄, and PO4. In addition, alkalinity may be estimated from titration experiments. Often solutions are also analysed for a whole suite of trace elements, including Fe, Mn, Pb, Zn and Cu. To minimize changes in sample composition it is common practice to analyse right after sampling. Particularly sensitive parameters in this respect are pH, DOC, DIC, Al, F, NH₄, NO₃ and PO₄.

Analysed solutes generally consist of a number of species, which often differ in their bioavailability. Therefore speciation of several solutes has become common practice in catchment research. Often speciation also allows for the estimation of solute activities, necessary for the study of element solubility controls in soils.

Speciation of Al is probably most common, and involves a separation in inorganic and organic monomeric Al (generally using an acidic cation exchanger) plus polymeric Al (Barnes, 1975; Driscoll, 1984). The Al³⁺ activity may be estimated from inorganic monomeric Al (including Al³⁺, as well as monomeric hydrolysed Al and complexes with Si, F and SO₄) using thermodynamics (e.g. Schecher and Driscoll, 1987). Dissolved F is often separated into total F and free F (after addition of TISAB buffer, which decomplexes F in solution). The difference between total and free F may be taken as an estimate of Al-bound F, and may be used as an alternative procedure to estimate the Al activity (from thermodynamics).

Since DOC consists of a wide range of organic substances with highly variable physicochemical characteristics (Pohlman and McColl, 1988), an analytical fractionation scheme was proposed by Leenheer and Huffman (1979) and Leenheer (1981). By using macro reticular resins DOC is fractionated into six classes: hydrophobic and hydrophilic acids, neutrals and bases.

5.5 SUMMARY

Streamwater runoff is dominated by drainage water from the terrestrial environment. Therefore biogeochemical processes in the terrestrial environment are of f paramount importance for the chemistry of streamwater.

This chapter discusses abiotic chemical reactions in solution, gas and solid c phases of the soil, with special reference to reactions involving inorganic C, organic C, SO₄, PO₄, Al, N and cation exchange. Reaction equilibria and kinetics are largely determined by soil chemical properties, which result from geological processes and subsequent soil formation.

Soil formation is the long-term vertical differentiation of physical, biological and chemical properties of rocks and sediments under the influence of soil formation factors, including parent material, climate, biota, topography and time. As soil-forming factors generally differ within one catchment, various soil types may result.

The importance of specific soil types for stream water quality depends on the water pathways in the terrestrial system. These pathways may vary as a function of precipitation intensity. Therefore the contribution of various soils (or soil horizons) to the stream varies dynamically with runoff.

Routines for sampling and pretreatment of samples from soils and soil solutions are numerous and arbitrary. Researchers need to critically evaluate their procedures in the light of their specific aim.

5.6 FURTHER READING5.5 SUMMARY

There are several excellent textbooks on soils in general, such as these of Brady (1974) and Scheffer and Schachtschabel (1982). Soil classification which is widely used is based on Soil Survey Staff (1975) and FAO- UNESCO (1988). Sampling and analysis of soils are covered in authoritative volumes of Klute (1986) and Page et al. (1986). Sampling of soil solutions is discussed by Litaor (1988).

General discussion of abiotic chemical reactions in soils appears in Bolt and Bruggenwert (1976), Stumm and Morgan (1981) and Greenland and Hayes (1981). More detailed topics of chemistry of humic substances in soils are given in Aiken et al. (1985) and Huang and Schnitzer (1986); aluminium chemistry of soils is discussed in depth by Sposito (1989).

Modelling soil solution chemistry and the underlying chemical concepts are explained in Reuss and Johnson (1986). Anthropogenically induced changes in forest soils are summarized and discussed by Johnson et al. (1991).

Wright and Hauhs ( 1991) discuss the question of the reversibility of acidification.

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