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Chapter 8 The redox–pH approach to the geochemistry of the Earth's land surface, with application to peatlands
Peatlands: Evolution and Records of Environmental and Climate Changes I.P. Martini, A. Martı´ nez Cortizas, W. Chesworth, Editors r 2006 Elsevier B.V. All rights reserved.
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Chapter 8
The redox–pH approach to the geochemistry of the Earth’s land surface, with application to peatlands W. Chesworth, A. Martı´ nez Cortizas and E. Garcı´ a-Rodeja
Introduction To a first approximation, any scheme that seeks to define the broad geochemical evolution of the land surface of the Earth needs to take the following features into account. The ubiquitous presence and importance of water. Geochemical change in materials in soils and other weathering systems depends primarily on water: its supply, abundance and residence time in the system. The water acts physically as an agent of transport, as a medium through which reactants diffuse to and from sites of reaction, and in freezing exerts a pressure capable of disintegrating rocks along discontinuities in their structure from the scale of a grain boundary to that of the jointing system. In addition it exerts a partial pressure that ranges from the saturation value to lower, but never completely zero, values in the arid regions. Chemically, water acts as a solvent, as a reactant in all important reactions in the weathering zone, a chemical constituent of many secondary minerals, especially clays and hydroxides, and as a chemical buffer fixing the thermodynamic activity of soil solution at one, or close to it, except in arid systems and episodes such as in the formation of sal/sodic soils. The ubiquitous presence and importance of organic matter. Organic materials may range from total dominance in the system as in ombrotrophic peatlands to a minor or insignificant role in very young or incipient soils such as lithosols. In the A horizon of a soil, organic matter undergoes decay, and organic acids and complexants are produced. These promote the weathering of inorganic constituents and the mobilization of inorganic ions in solution. The concentration of organic material attenuates with depth, and is generally taken to be negligible, or at least not obvious in the C horizon of a soil. The nature of the inorganic substrate. The dominant minerals in virtually all soils are silicate or aluminosilicate in composition. Over time they tend to be stripped of the so called basic cations, and to be converted into some combination of quartz, 1:1 sheet silicates, and hydroxides of Al and Fe, thereby producing acid soils of low ISSN: 0928-2025
DOI: 10.1016/S0928-2025(06)09008-0
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cation exchange capacity. If calcite is present, it reacts more rapidly than the (alumino)silicates and a pH between 7 and 8 is established. A rarer occurrence is the presence of sulfides, pyrite being the most likely. It too weathers more rapidly than the (alumino)silicates, and gives rise to waters with a pH of 3 or less. Of the elemental components in the inorganic substrate, Fe plays an extremely important role. It is the most abundant element with more than one valence, and takes part in reactions between the soil solution and ferric-hydroxy phases that relatively quickly set the overall redox state of the soil. The physical nature of the substrate is also of importance, particularly its hydrodynamic properties. The latter determine rates of drainage of the material and hence the degree of saturation with respect to water. The partial pressure of O2. Initial values (20 kPa) are set by atmospheric O2. In the vadose zone of the weathering system it will be essentially the same near the surface, but with depth and the increased biological oxygen demand of aerobic decay of organic matter the value will be less, and as microbial organisms proliferate and use up all the available resource of uncombined oxygen, will tend to reach the limit where PO2 is equal to PH 2 , and the system as a whole changes from aerobic to anaerobic. This limit will normally coincide with the water table, below which any pore space will be saturated with water to the exclusion of all oxygen except the small amounts actually dissolved in the water right at the water table. In some cases, lateral inflow of oxygenated water, may result in anomalously high levels of oxygen within the saturated zone. The partial pressure of CO2. Again the initial value is set by the atmospheric partial pressure: 101.5 kPa (103.5 atm). Respiration in the root zone, and decay of organic matter may raise this up to a maximum of about 100 times the atmospheric value. PCO2 determines the pH of the H2O–CO2 (carbonic acid) system, with 5.7 as the common value in unpolluted rainwater. In the absence of organic acids, carbonic acid is the most important one in acid–base reactions in the weathering mantle. This means that its chemical importance in soil tends to increase with depth as the importance of the organic component diminishes. Clearly there is ample opportunity for feedback among these factors, and taken together, they determine that the commonest types of chemical reaction within a weathering system are of acid–base (including cation-exchange) and redox type. Consequently, serviceable choices of master variables for the soil environment are those that are related to the thermodynamic activities of the proton and the electron. Regarding the activity of protons in a system, the universal variable of choice is pH. Regarding electrons, the equivalent parameter pe may be chosen, but the classic variable employed is Eh, which has the advantage of being measurable both in laboratory and field. As shown in Figure 8.1, there is a simple relationship between pe and pH. The application of this type of diagram to weathering systems generally and to peatlands in particular will now be described.
Physico-chemical background Figure 8.1 is variously referred to as a redox–pH diagram, an Eh–pH diagram, a pe–pH diagram, or simply a Pourbaix diagram, after the metallurgist who devised this
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Figure 8.1. Pourbaix diagram. The sloping, dashed line separating oxidizing from reduced conditions is for PH 2 ¼ PO2 in the equilibrium H2O ¼ H2+0.5 O2.
type of Cartesian graph for the purposes of examining corrosion (Pourbaix, 1974). In earth science their use was pioneered by Krumbein and Garrels (1952), and a good introduction is provided by Garrels and Christ (1965). A compilation of basic diagrams useful to the geochemist may be found in Brookins (1988). Truesdell (1968) suggested the adoption of pe rather than Eh as the redox parameter as a means of simplifying the calculation of redox–pH diagrams from fundamental thermodynamic data. Berner (1970) details a number of practical limitations related to the construction and interpretation of Eh–pH diagrams. At the most fundamental level, there is the problem of the quality of the thermodynamic data upon which the diagrams are based. In this regard, Woods and Garrels (1987, p. ix) refer to ‘‘the perils of indiscriminate selection of values from various sources’’. A major difficulty, and one which clearly has a determining influence on any ultimate conclusions to be drawn from the diagrams, lies in choosing what minerals to show on them, whether to consider metastable phases, whether to consider solid solution in condensed phases, and what to do about bacterial reductions. Furthermore, it is assumed that equilibrium (stable or metastable) holds for a particular diagram constructed. Finally, there is the major difficulty of obtaining good measurements of Eh in the field (Shotyk, 1988). In spite of all this, a consensus exists that the diagrams are useful in indicating general conditions and general tendencies in a system. Even for more parochial considerations, redox–pH diagrams may prove useful provided that matters of scale, and the possibilities of local reactions are kept in mind.
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Figure 8.2. Predominance fields in the system Fe–Ca–K–S–CO2–H2O. Conditions: 100 kPa total pressure, 251C, activities of all components except S, 106. S is undefined in terms of a specific activity in order to show the sulfide field at its maximum extent.
Predominance fields Figure 8.2 is a redox–pH diagram of the type used in the geochemical interpretation of low-pressure/temperature environments. Areas called predominance fields are labeled in terms of the minerals that would be expected to form under the Eh and pH combinations of those fields. The actual positions of the boundaries between fields depend on the choice of ionic activities made in calculating the appropriate mineralogical reactions for the system. A total pressure of 100 kPa and a standard temperature of 25 oC are usually chosen as representative ambient conditions. This diagram is of special interest in the context of the geochemistry of peatlands, as will be made clear later.
Geochemical fences A notable contribution of Krumbein and Garrels (1952) to the usefulness of the Eh–pH approach in earth science is the concept of the geochemical fence: a relatively narrow, linear zone in redox–pH space defined by a specific chemical reaction or by a related group of reactions. Their original diagram, developed to help in the
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Figure 8.3. The original system of geochemical fences devised by Krumbein and Garrels (1952).
interpretation of sedimentary environments, is shown in Figure 8.3. It requires modification if the full scope of soil conditions are to be covered, especially below pH 7. In practical terms this amounts to changing existing geochemical fences, and adding new ones. The modified fences, shown in Figure 8.4, are: Organic fence. Within the upper organic-rich part of the solum, microbially mediated breakdown of organic matter will give a range of possible Eh values for a given pH, up to the lower limit of oxidation by O2 in soil. The latter will likely coincide with the water table in the soil system as a whole, with the exclusion of oxygen to microsystems being controlled by details of micromorphology, for example by the presence of clay cutans protecting a mineral surface from oxidation. Above the water table the microbial biomass, using dead organic matter as its carbon source, will generally experience exponential growth up to the point where free O2 is fully utilized. Again, this will obviously coincide with the lower limit of oxidation by O2. It should therefore be placed at the Eh–pH contour along which PO2 is equal to PH 2 in the gaseous dissociation reaction of H2O, rather than where Eh is zero, the value chosen by Garrels and Christ (1965). In theory, the upper limit would be marked by the partial pressure of oxygen in the atmosphere. In practice, the biological oxygen demand in virtually all soils depletes oxygen to lower levels than this.
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Figure 8.4. The pedogenic grid (after Chesworth 1992), an extension of the Krumbein and Garrells diagram (Fig. 8.3) to include new fences and fences modified from the diagram. The extended diagram is applicable to pH values lower than Krumbein and Garrells consider, including values found in bogs, many mires, and acid soils (see Figs 8.9, 8.11, 8.13).
The width of the organic fence shown in Figure 8.7 is consistent with measurements in such materials. Neutral fence. Garrels and Christ (1965) give no mineralogical justification for a neutral fence, and in the ubiquitous presence of atmospheric and respiratory CO2 in a weathering system, it is unlikely that the ionic dissociation of water will buffer a natural system to a pH of 7. However, an average smectite (Weaver and Pollard, 1973), with so-called basic cations (Mg, Ca, Na, K) in exchange positions, marks a near-neutral fence according to the reaction: Mg0:2 ðSi3:81 Al1:71 Mg0:29 O10 ðOHÞ2 ðmontmorilloniteÞ þ 0:98 Hþ þ 0:22e þ 0:33H2 O ¼ 0:86 Al2 Si2 O5 ðOHÞ4 ðkaoliniteÞ þ 0:22 FeOOH ðgoethiteÞ þ 2:1SiO2 ðquartzÞ þ 0:49Mg2þ Because of minor substitution of Fe in the octahedral layer it will not be a vertical fence as in the Garrels and Christ diagram (Garrels and Christ, 1965). It should also be realized that this particular fence does not mark the lower pH limit of a smectite predominance field, since there exists a spectrum of smectites down to beidellite that is stable at low pH. Fe hydroxide fence. Garrels and Christ (1965) based their Fe-oxide fence on the solubility of Fe2O3. In the weathering regime Fe-hydroxy phases are overwhelmingly
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more important. Goethite is generally accepted as the stable hydroxide, though ferrihydrite is probably its precursor. The new fences in Figure 8.4 are : An aluminosilicate fence. This is defined by reactions in the system SiO2– Al2O3–H2O. For example by the congruent dissolution of kaolinite: Al2 Si2 O5 ðOHÞ4 ðkaoliniteÞ þ 6Hþ ¼ 2Al3þ þ 2SiO2 ðquartzÞ þ 5H2 O ðpH4SiO4 o4:7Þ by the incongruent dissolution of kaolinite: Al2 Si2 O5 ðOHÞ4 ðkaoliniteÞ þ 5H2 O ¼ 2AlðOHÞ3 ðgibbsiteÞ þ 2H4 SiO4 ðquartzÞ ðpH4SiO4 o4:7Þ and by dissolution of gibbsite: AlðOHÞ3 ðgibbsiteÞ þ 3Hþ ¼ Al3þ þ 3H2 O A sodium carbonate fence. This defines the high-pH extreme of the soil environment, and lies in the pH range of feldspar dissolution, if the latter were to reach a saturated equilibrium, as well as the range in which zeolites form authigenically under earth-surface conditions. In addition, two fences on the original Krumbein and Garrels diagram are left essentially in their original form. These are: The limestone fence. Renamed the calcite fence, on the high pH side of which solid calcite persists in a soil system. The sulfate– sulfide fence. Name simplified to sulfide fence, marking the upper limit of Eh of the field wherein sulfides, and in particular pyrite, may form in a soil. Other geochemical fences may be useful in specific or local situations. For example, in ombrotrophic bogs with low-atmospheric input of particulate matter, the lowpH limit could be marked by a fence defined by the dissociation of H+ from carboxylic functional groups in the organic component of the soil system. An even lower pH fence can be defined for the interpretation of acid-sulfate soils by the breakdown products of pyrite. Where neither organic matter nor pyrite is common, the system H2O–CO2 will be useful in defining a fence from about pH 5.7 to 5.5 depending upon the balance between atmospheric and respiratory inputs. Under extreme reduction a CH4 fence comes into play.
The pedogenic grid The resulting adaptation (Fig. 8.4) of the Krumbein and Garrels (1952) diagram is referred to as the pedogenic grid (Chesworth, 1992) and covers the broad range of conditions in common mineral soils. Certain organic soils and peatlands, and less common soils such as acid-sulfate soils, lie outside this range. Parallels exist between the pedogenic grid and the pressure–temperature (PT) diagram of the petrologist. The invariant point of a PT diagram, where pressure and temperature are held constant, is similar to those regions on the grid where two
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fences cross. Chesworth (1992) called such crossing points nodes. They act as quasiinvariant points where the pH and Eh of a system is buffered and poised, respectively. A further analogy arises from the fact that the essentially univariant (or pseudounivariant) fences of the grid delimit separate ranges of conditions with each area being characterized by its own individual assemblage of minerals. This is implicit in the diagram of Krumbein and Garrels (Fig. 8.3).
Geochemical trends in the weathering zone Three general trends are followed in the geochemical evolution of surface materials in the weathering zone (Fig. 8.5). Essentially they are the result of the pumping of protons and electrons between sources and sinks in the land environment. Proton and electron pumps A consequence of considering the activities of the proton (H+) and electron (e) as the major driving forces in effecting chemical change in soils (H+ in hydrolysis, and acid-base reactions generally, e in oxidation–reduction reactions) is that conceptually both can be considered as being pumped between a source (or sources) and sink (or sinks) in the weathering regime. First consider H+. On the largest scale, weathering may be described as the titration of acids from atmospheric and organic sources, against bases represented by the aluminosilicate, carbonate, and other mineralogical constituents of the earth’s surface. Figure 8.6 illustrates this within Eh–pH space as a proton pump between source and sink. In a humid climate, with adequate atmospheric precipitation, the land surface is ‘‘inevitably over-titrated, acting as a sink for protons’’ (Edmond et al., 1979, p. 21). Given enough time, all so-called base cations (Na, K, Ca, Mg) will be stripped from solid phases and removed by leaching. Starting from calcareous materials, e.g., solution of calcite, would buffer the system between pH 7 and 8 until it was completely dissolved. Then progressive overtitration would lead to hydrolysis and the formation of a residuum made up of such minerals as kaolinite, gibbsite, and goethite, together with resistant minerals, of which quartz is the commonest. Eventually (and with the proviso that no metal sulfides were present) a steady (or quasi-steady) state would be set up, buffered by the aluminosilicate fence of the pedogenic grid. How close a given weathering system will approach this steady state will depend largely on the availability of water, and the hydrodynamics of the landscape. A close approach is favored by high rainfall in warm climates, and with a good drainage to carry away dissolved cations. In addition, the persistence of phases such as calcite and smectites saturated with the so-called basic cations Ca, Mg, Na, and K will buffer pH and retard the progress of acidification. In humid conditions, and taking the long-term geological perspective, this can only be considered a transient state, the lifespan of which will depend on such factors as landscape position (shorter lifespan on slopes and topographic highs, a longer one in less well-drained topographic lows)
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Figure 8.5. The three major geochemical trends in the weathering zone on the land surface of the Earth. The arrows indicate an acid trend characteristic of regions of humid climate, an alkaline trend characteristic of dry climates, and a reduced (or hydromorphic) trend found, for example, under waterlogged conditions in peatlands and gleysolic soils.
and texture of the weathering materials, coarse, well-drained deposits leading to a shorter persistence of the basic minerals, than finer, less well-drained ones. Semi-arid and arid conditions will lead to an opposite trend; that is, a trend of undertitration wherein it is the basis in the system that predominates over the additions of H+, and the system is buffered at an alkaline pH. The development of sodicity, and in the limit the precipitation of salts of the alkali metals, is the likely result. Figure 8.7 is a similar interpretation regarding e. The most reduced material in the biosphere is the living biomass. When death and decay intervene, organic materials become a prolific source of electrons in soil and in effect pump electrons to any available sink, the largest of which is atmospheric oxygen. For Malthusian reasons (the tendency of organisms to use resources to the maximum), any resource they depend upon will normally mean that the overall Eh of the soil will be poised along the contour in Figure 8.1 that represents this limit (PO2 ¼ PH 2 ). For example, if the ready diffusion of oxygen into the soil is impeded, for reasons of water logging, heavy texture, or compaction , microbial activity will be fuelled by
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Figure 8.6. The proton pump, showing the movement of hydrogen ions from sources such as dissociating carboxylic groups in decaying organic matter, and from carbonic acid from atmospheric precipitations, to sinks represented by the inorganic materials of the Earth surface, ranging from carbonate to silicate materials. In effect, this is a natural titration, which in humid climates leads to overtitration (and low pH values at the Earth surface) and in dry climates to undertitration (and higher values at the Earth surface).
the movement of electrons from the organic matter to some other available sink. The range of operation of these alternative electron acceptors is well known, and is shown in Figure 8.8. At finer scales, particularly in the contact zones between phases, and on the reactive interface between solid and soil solution, an immense number of transfers of H+ and e are possible. On the whole, however, and on the macroscopic scale of the solum presented here, the general tendencies outlined are reasonable ones, at least semi-quantitatively. The peatland environment From the geochemical standpoint of this chapter, the most significant differences within the peatland environment are shown within the framework of the system Fe–Ca–K–S–CO2–H2O (Fig. 8.2). Our interpretation is based on the following observations developed from the foregoing discussion. Presence of a water table divides any peatland into an oxidized and a reduced zone. This of course is recognized in the acrotelm/catotelm vertical stratification in mires.
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Figure 8.7. The electron pump. This represents the overall movement of electrons at the macroscale of a bog, mire, or soil landscape from the principal source in decaying organic matter, to the principal sink of atmospheric oxygen. In a decaying peat, the oxidation state would rise in the direction of the heavy arrows. Oxidation of a peat deposit in the water-unsaturated acrotelm, would lead to conditions above the PO2 ¼ PH 2 contour, and as oxidation proceeded PO2 ¼ 100 kPa contours. Conditions in the watersaturated catotelm would normally rise no further than the PO2 ¼ PH 2 contour, unless there was an influx of oxygenated water. However, biological oxygen demand would determine that the oxygenated state did not persist.
The predominance of organic matter means that it is the principal source of H+ for acid–base reactions. The definitive distinction between ombrotrophic and minerotrophic peatlands is the addition of extraneous materials only via the atmosphere in the former case, and additionally (and more copiously) via ground and surface waters in the latter case. Ombrotrophic mires The geochemistry of the ombrotrophic environment is the simpler of the two types. Living and decaying vegetation (Sphagnum, Carex, and other genera) dominates the environment and pH is determined basically by the dissociation of carboxylic functional groups in the dead organic matter. Since the rate of addition of terrigenous minerals and the rate at which many of them dissolve in the acidic bog waters are too
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Figure 8.8. Several sources and sinks of electrons in the land surface. At a microscale in a bog, mire, or soil landscape, the movement of electrons is more complicated than the big picture shown in Figure 8.7. Many more sources (for example, ferrous iron, sulfides, native metals ) and sinks (for example, manganous oxides and hydroxides, and nitrates) exist, and will influence the detailed geochemistry of a natural system, but in a peatland the overall geochemistry will be dominated by the reactions between organic matter and oxygen as shown in Figure 8.7.
slow to consume the protons generated by CO2 and organic acids the pH can be expected to remain constant with depth as shown in Figure 8.9. Eh will vary independently, and in the catotelm may reach levels of reduction low enough to allow the generation of CH4. It is possible that oxygenated groundwater may enter the system below the water table, and rise up in a plume, but in general oxygen is unlikely to be the terminal electron acceptor within the zone of water saturation. The fate of the scarce wind-borne additions to ombrotrophic bogs have been discussed by Martı´ nez Cortizas et al. (2001a, b) in the context of the predominantly quartzo-feldpathic additions to the wetlands of Galicia, northwestern Spain (Fig. 8.10). The material is assumed to be added in amounts too small to have an effect on the overall pH of the environment.
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Figure 8.9. The ombrotrophic environment. The field of ombrotrophic peatlands will be approximately as shown in the gray, cigar-shaped area, with the relatively low pH determined by the dissociation of carboxyllic acid groups in organic matter. Generally, eolian inputs of inorganic dust will be too feeble to move the pH to higher values.
Minerotrophic mires The increased importance of extraneous additions to minerotrophic as opposed to ombrotrophic peatlands shows up in two ways. First, the addition of Fe-containing minerals (ferromagnesians, oxides or hydroxides) and their reaction with water tends to produce a redox profile that follows a contour parallel to the boundary of the ferric-hydroxide field of predominance. Second, hydrolysis of added minerals will raise pH, most dramatically if calcite or dolomite is in the mix. The boundary of the predominance field of calcite will provide an upper limit in this case, at least as long as carbonates persist. Where only primary silicates are added (and the most common will be feldspar) a raised pH will be more slowly established and will not normally reach the calcite field. The possibilities are shown in Figure 8.11. Discussion Mires, mainly ombrotrophic ones, are extensively used as archives of compounds deposited by wet and dry deposition from the atmosphere, heavy metals in particular.
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Figure 8.10. Reaction products of terrigenous dust in the ombrotrophic environment of Northwestern Spain, showing the local suite of input minerals, the PH 2 and PO2 ranges of the local blanket bogs and the resulting reaction products in six segments of the PH 2 PO2 framework, shown as mineral facies diagrams based on a triangular projection of the system Si–Al–Fe–O–H–S.
The basic idea underlying this kind of research is that dust and other compounds deposited on the mire will eventually be buried as plant remains accumulate. Any subsequent core may then be sectioned and analyzed to give a high-resolution record of changes in atmospheric deposition for up to a few thousand years. Elemental mobilities in the peat are a potential complicating factor, though most studies assume that elements of interest are immobile. Few papers actually address the geochemical conditions of the mire in terms of their effect on the stability of the mineral phases hosting an element (Shotyk, 1995; Steinmann and Shotyk, 1997) and thereby on the possibility of that element being released in mobile form. Four aspects are of particular interest in this regard: (1) the mineralogical composition of the deposited dust, (2) the geochemical environments of the mire (the acidic, oxidized acrotelm and the slightly less acidic but anoxic catotelm), (3) the rate of peat accretion, and (4) the interference of human activities on the natural composition of the atmospheric dust and aerosols. Materials deposited from the atmosphere onto the surface of the mire are incorporated initially into the acid/oxidized environment of the acrotelm. Most primary minerals are unstable in this environment, and at the extreme (pH below 4), most
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Figure 8.11. The minerotrophic environment. The lower pH values or the cigar-shaped field of the minerotrophic mines are defined by ferrous–ferric equilibria that partially delimit the predominance field of ferrihydrite, a precursor of the stable-phase goethite, shown in Figure 8.10, and formed as an initial product of the breakdown of ferromagnesians or magnetite. Depending on the amount and type of inorganic crustal material present in a mire, the upper pH limit of the minerotrophic environment will extend in the direction of the heavy arrows, towards the calcite-present area.
common minerals, primary or secondary, are unstable. Only if a given element is hosted in stable minerals may its concentration in the peat be assumed to reflect that of the original dust (though correction for the effect of the organic substrate may be necessary). Problems arise when the inorganic phase breaks down and an element is released as an ion, into solution. Once in solution it may leach (essentially through lateral movement) or diffuse through the mire water column following concentration gradients. Figure 8.12 exemplifies this by showing the profile of Fe concentrations in an ombrotrophic mire. The highest Fe concentrations occur in the aerated acrotelm of the upper 25 cm of the bog; below this, concentrations abruptly decrease to reach minimum values where the fluctuation of the water table takes place. Below this level (and particularly below 110 cm in the example shown) values increase again. A similar pattern was obtained in an experiment where cores from uncontaminated sites were transplanted to contaminated sites, and vice versa. Iron concentrations in the peat core of the contaminated mire decreased after being transplanted to the uncontaminated site, whereas the opposite was found for the peat core of the uncontaminated mire after it was transplanted to the contaminated one (Martin Novak, personal communication). This indicates that Fe was mobilized in or out of the
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Figure 8.12. Profile of Fe (ppm) concentrations in an ombrotrophic bog (Pena da Cadela, Xistral Mountains, NW Spain). The vertical scale is in centimeters, the horizontal scale of the Fe profile is in parts per million. Such a profile might be interpreted as being due to a greater influx of Fe in more recent dusts than in older ones (represented by deeper parts of the bog). However, in the upper 50 cm in which Fe varies 2.3-fold, normalization against a known immobile element (Ti) shows a variation of 3.2-fold. The fact that the two ranges are different is prima facie evidence that Fe is mobile. Extrapolating to the total profile suggests the possibility that Fe has also been mobilized at depth, a possibility supported by the transplant experiments described in the text.
transplanted cores, following concentration gradients. Using a chemical speciation approach, Shotyk (1995) also suggested that there is no obvious chemical mechanism to retard the possible migration of As and Sb. However in a recent investigation of a peat core from a blanket bog from the Faroe Islands it was found that Sb and Pb
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records were highly correlated, and this correlation was taken as evidence of effective Sb immobilization (W. Shotyk, personal communication). Other authors have also indicated that changes in redox potential, dependent on water table fluctuations, are the most important mechanisms in the distribution of heavy metals (Damman, 1978; Clymo and Hayward, 1982; Damman et al., 1992). Steinmann and Shotyk (1997), on the other hand, found relatively high Fe3+ to Fe2+ ratios in bog-pore waters, despite the anoxic conditions, which they explained in terms of the higher thermodynamic stability of the organic complexes of Fe3+ compared to those of Fe2+. Unlike the case of a mineral soil, in ombrotrophic mires released elements are unlikely to fix by complexation to another inorganic phase, since no more than 2% of the total dry mass of the peat is inorganic matter. However, released elements may obviously become bound to organic matter, for example, Hg shows a marked affinity for humified organic matter (Benoit et al., 1994), with binding to reduced sulfur groups being preferred (Skyllberg et al., 2003). Complexation is the most important mechanism for metal fixation, as Shotyk (1995) has shown. He calculated that in bogs 99.9% of Cu, Pb, and Zn can be bound to an organic component, either in the solid phase or in solution (as dissolved organic carbon, DOC). So, organically bound elements can be retained by the peat or they can be leached from the system with the dissolved organic matter. In other words, the extent to which the original concentration of a given element will be changed subsequent to the addition of that element to a mire will depend on the degree to which the element is partitioned between the immobile solid phase and the mobile dissolved phase. This in turn will depend upon the kinetics of the reactions that produce the soluble chemical species. Finally, the hydrodynamics of the system plays an important role. If the mobile phase is definitively trapped within the mire, no compositional change will take place. If, as seems most likely, the mobile phase finds an exit route, then elements complexed to DOC will leave the system. In minerogenic mires the higher mineral content makes the role of the inorganic compounds more relevant in terms of the redistribution of metals. For example, Franze´n et al. (2004) analyzed a minerogenic mire and found enrichments in Hg associated to elevated Fe concentrations and the presence of goethite, suggesting Hg adsorption to Fe-oxyhydroxides probably as Hg-fulvo-acid complexes. One important aspect of the fate of organically bound elements is that their potential release is coupled to that of dissolved organic matter (DOM) and linked to climate changes. As indicated by Biester et al. (2006 – this book, Ch. 19) the release of organo-halogens is expected to increase during warm and wet periods and decrease during dry periods; this may also affect other organically bound elements (that is, metal-DOM complexes). Continuous growth of vegetation in the surface of the mire tends to keep the acrotelm at an almost constant thickness (Belyea and Clymo, 2001). By contrast, the catotelm grows upwards. This means that any given peat surface will reach the catotelm in a few hundreds to a thousand years (Martı´ nez Cortizas et al., 2001a, b). Thus minerals/elements progressively sink into a more anoxic environment as time goes by. Vertical water movement is largely impeded in the catotelm (Ingram, 1983)
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and the fate of an element in this section will mostly be governed by redox reactions and diffusion. Human activities can affect the amount and composition of the deposited dust in mires. Induced soil erosion due to forest clearance, fires, and agriculture have increased the amount of atmospheric dust and changed its composition (by increasing the surface area and types of exposed sources: soils and rocks) especially since the so called Neolithic Revolution; that is, from 10,000 to 5,000 years ago (Martı´ nez Cortizas et al., 2005). The most prominent changes in atmospheric composition occurred once the smelting of metals began. In the last 250 years, the changes have accelerated as a consequence of the Agricultural and Industrial revolution. Metallurgy, mining, fuel burning, waste incineration, and other human activities have released to the atmosphere compounds that are markedly different from those originating from soils and rocks. As with dusts of natural origin, the anthropogenic material is also subject to possible modification in the mire-environment. For example, Sonke et al. (2002) investigated metal pollution in a bog close to a Zn-smelter in the Kempen region (Belgium) and found that though oxidized smelter dusts were found in the nearby mineral soils, the bog showed extensive in situ FeS2 and ZnS, indicating that anoxic conditions were responsible for rapid diagenetic transformations of anthropogenic metal oxides into sulfides. Rausch et al. (2005a, b) investigated solid-phase and pore-water concentrations of some trace metals in uncontaminated and contaminated sites in Finland and found that Cu was generally retained more effectively than Ni, Co, Zn, or Cd, but in highly contaminated sites even Cu was substantially transferred from the solid phase to the pore waters. These authors indicate that ombrotrophic bogs may or may not function as reliable archives of atmospheric trace metals, depending on the concentration and chemical properties of the elements considered, the mineralogical form at the time of deposition, as well as the pH of the bog waters. As a general rule, they hypothesize that sites that are remote from point-pollution sources and receive metals exclusively by atmospheric deposition are more conducive to preserving the chronology of deposition. Not only element concentrations can be affected by ambient geochemical conditions, isotope ratios may also be modified. The bulk isotopic composition of peat arises from a mixture of sources (for example, soil dust, anthropogenic pollutants), which can have different isotope ratios. This characteristic is used, for instance, in tracing the source of Pb pollution, since the isotopic signature of Pb from an anthropogenic source such as gasoline, differs from that of Pb in soil dust. However, there may be considerable heterogeneity in the Pb from different soil pools (Emmanuel and Erel, 2002; Wong and Li, 2004). Consequently the materials added to a mire, whether they are from natural or from anthropogenic sources, and whether they are from finer or coarser fractions, from the primary mineral or secondary mineral, or other fractions of the soil, will show a range of geochemical stabilities in the peat environment. This will complicate attempts to decipher the geochemical history of a peatland. In addition, if the hydrodynamics of a given system is conducive to the leaching of the more labile weathering products, post-depositional changes of the original composition, including its bulk isotope ratio become possible, and use of the peat as an archive is compromised. The danger is that these changes may be interpreted as changes in Pb atmospheric composition through time.
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Conclusions Redox–pH diagrams provide a concise graphical way in which to depict Earthsurface environments in terms of the activities of electrons, the operators involved in the two commonest types of reaction in the weathering zone on land, i.e., redox and acid–base reactions. Figure 8.13 illustrates the following points by way of summary. Peatlands tend to fall on the acid side of the diagrams with ombrotrophic bogs showing least overlap with common soil-forming environments. Minerotrophic mires overlap the field of ombrotrophic bogs under relatively high redox conditions, though with depth they move along a line defined by ferrous–ferric equilibria, to a near-neutral pH. The same kind of equilibria also plays a controlling influence on ambient conditions in mineral soils at the low limit of their pH range. The upper limit of the minerotrophic field in peatlands, however, will depend on whether or not there are extraneous additions of carbonates. If there are ambient conditions, the peatland may extend to the calcite predominance field, where a pH approaching 8 is possible. Within this extended field lies the environment of histosol formation. Histosols may be looked upon as a kind of bridge between peatlands and mineral soils.
Figure 8.13. The geochemical environments of bogs and mires shown in the context of the Eh–pH range of common soils and of acid-sulfate soils.
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In the context of this volume, the most important conclusion is that the peatland environment covers a wide range of geochemical conditions and encompasses those under which most primary, and many secondary minerals would break down. If in fact inorganic additions to a peatland do break down, the possibility of the mobilization of the elemental components must be entertained. Consequently redox–pH diagrams offer a first line of defense against the hasty use of any given proxy element assumed to be immobile, in the reconstruction of past environmental conditions.
References Belyea, L. and Climo, R.S., 2001. Feedback control on the rate of peat formation. Proc. Roy. Soc. Lon. 268, 1315–1321. Benoit, J.M., Fitzgerald, W.F., and Damman, A.W.H., 1994. Historical atmospheric mercury deposition in the mid-continental U.S. as recorded in an ombrogenic peat bog. In: Huckabee, J. and Watras, C. (Eds), Mercury Pollution: Integration and Synthesis. Lewis Publishers, Chelsea, MI, pp. 187–202. Berner, R.A., 1970. Behaviour of iron during weathering, sedimentation and diagenesis. In: Wedepohl, K.H. (Ed.), Handbook of Geochemistry, Section 26G. Springer, Berlin, pp. 1–8. Biester, H., Martı´ nez Cortizas, A., and Keppler, F., 2006. Occurrence and fate of halogens in mires. In: Martini, I.P., Matı´ nez Cortizas, A., and Chesworth, W. (Eds), Peatlands: Evolution and Records of Environmental and Climatic Changes, Ch. 19. Elsevier, Amsterdam. Brookins, D.G., 1988. Eh–pH Diagrams for Geochemistry. Springer, Berlin, 176pp. Chesworth, W., 1992. Weathering systems. In: Martini, I.P. and Chesworth, W. (Eds), Weathering, Soils and Paleosols. Elsevier, New York, pp. 19–40. Clymo, R.S. and Hayward, P.M., 1982. The ecology of Sphagnum. In: Smith, A.J.E. (Ed.), Bryophyte Ecology. Chapman & Hall, London, pp. 229–289. Damman, A.W.H., 1978. Hydrology, development, and biogeochemistry of ombrogeneous peat bogs with special reference to nutrient relocation in a western Newfoundland bog. Can. J. Bot. 64, 384–394. Damman, A.W.H., Tolonen, K., and Sallanus, T., 1992. Element retention and removal in the ombrotrophic peat of Hadekeidas, a boreal Finnish peat bog, Suo 43, 137–145. Edmond, J., Corliss, J.B., and Gordon, L.I., 1979. Ridge crest hydrothermal metamorphism in the Galapagos spreading center and reverse metamorphism. In: Talwani, M., Hay, W.W., Ryan, W.B.F. (Eds.), Deep Drilling Results in the Atlantic Ocean : Continental Margins and Paleoenvironment. Proceedings of the 2nd Maurice Ewing Symposium, American Geophysical Union, Washington, DC, pp. 383–390. Emmanuel, S. and Erel, Y., 2002. Implications from concentrations and isotopic data for Pb partitioning processes in soils. Geoch. Cosmochim. Acta 66, 2517–2527. Franze´n, C., Kilian, R., and Biester, H., 2004. Natural mercury enrichment in a minerogenic fenevaluation of sources and processes. J. Environ. Monit. 6, 466–472. Garrels, R.M. and Christ, C., 1965. Solutions, Minerals and Equilibria. Harper & Row, New York, 450pp. Ingram, H.A.P., 1983. Hydrology. In: Gore, A.J.P. (Ed.), Ecosystems of the World 4A: Moors: Swamp, Bog, Fen and Moor. Elsevier, Amsterdam, pp. 67–224. Krumbein, W.C. and Garrels, R.M., 1952. Origin and classification of chemical elements in terms of pH and oxidation–reduction potentials. J. Geol. 60, 1–33. Martı´ nez Cortizas, A., Chesworth, W., and Garcı´ a-Rodeja, E., 2001b. Dina´mica geoquı´ mica de las turberas de Galicia. In: Martı´ nez Cortizas, A., Garcı´ a-Rodeja, E., (Eds.), Turberas de Montan˜a de Galicia, Coleccio´n Te´cnica de Medio Ambiente. Xunta de Galicia, pp. 141–148. Martı´ nez Cortizas, A., Garcı´ a-Rodeja, E., and Chesworth, W., 2001a. Geochemistry of Si–Al–Fe Minerals in Mire Environments of Northwestern Spain, ICOBTE (International Conference on the Biogeochemistry of Trace Elements), Proceedings, Guelph, Ontario, Canada. p. 491. Martı´ nez Cortizas, A., Mighall, T., Pontevedra-Pombal, X., et al., 2005. Linking changes in atmospheric dust deposition, vegetation change and human activities in northwest Spain during the last 5,300 years. The Holocene 15, 698–706.
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Pourbaix, M., 1974. Atlas of electrochemical equilibria in aqueous solutions. Natatinal Association. Corrosion Engineering, Houston, 644pp. Rausch, N., Nieminen, T., and Ukomaanaho, L., 2005a. Comparison of atmospheric deposition of copper, nickel, cobalt, zinc, and cadmium recorded by Finnish peat cores with monitoring data and emission records. Environ. Sci. Technol. 39, 5989–5998. Rausch, N., Ukomaanaho, L., and Nieminen, T., 2005b. Porewater evidence of metal (Cu, Ni, Co, Zn, Cd) mobilization in an acidic, ombrotrophic bog impacted by a smelter, Harjavalta, Finland and comparison with reference sites. Environ. Sci. Technol. 39, 8207–8213. Shotyk, W., 1988. Review of the inorganic geochemistry of peats and peatland waters. Earth-Sci. Rev. 25, 95–176. Shotyk, W., 1995. Natural and anthropogenic enrichments of As, Cu, Pb, Sb, and Zn in ombrotrophic versus minerotrophic peat bog profiles, Jura Mountains, Switzerland. Water Air Soil Pollut. 90, 375–405. Skyllberg, U., Qian, J., Frech, W., et al., 2003. Distribution of mercury, methyl mercury and organic sulphur species in soil, soil solution and stream of a boreal forest catchment. Biogeochemistry 64, 53–76. Sonke, J.E., Hoogewerff, J.A., van der Lanna, S., and Vangronsveld, J., 2002. A chemical and mineralogical reconstruction of Zn-smelter emissions in the Kempen region (Belgium), based on organic pool sediment cores. Sci. Tot. Environ. 292, 101–119. Steinmann, P. and Shotyk, W., 1997. Chemical composition, pH, and redox state of sulfur and iron in complete vertical porewater profiles from two sphagnum peat bogs, Jura Mountains, Switzerland. Geoch. Cosmoch. Acta. 61, 1143–1163. Truesdell, A.H., 1968. The advantage of using pe rather than Eh in redox equilibrium calculations. J. Geol. Educ. 16, 17–20. Weaver, C.E. and Pollard, L.D., 1973. The Chemistry of Clay Minerals. Elsevier, New York, 213pp. Wong, C.S.C. and Li, WX.D., 2004. Pb contamination and isotopic composition of urban soils in Hong Kong. Sci. Total Environ. 319, 185–195. Woods, T.L. and Garrels, R.M., 1987. Thermodynamic Values at Low Temperature for Natural Inorganic Materials. Oxford University Press, New York, 266 pp.
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The redox–pH approach to the geochemistry of the Earth’s land surface, with application to peatlands W. Chesworth, A. Martı´ nez Cortizas and E. Garcı´ a-Rodeja
Introduction To a first approximation, any scheme that seeks to define the broad geochemical evolution of the land surface of the Earth needs to take the following features into account. The ubiquitous presence and importance of water. Geochemical change in materials in soils and other weathering systems depends primarily on water: its supply, abundance and residence time in the system. The water acts physically as an agent of transport, as a medium through which reactants diffuse to and from sites of reaction, and in freezing exerts a pressure capable of disintegrating rocks along discontinuities in their structure from the scale of a grain boundary to that of the jointing system. In addition it exerts a partial pressure that ranges from the saturation value to lower, but never completely zero, values in the arid regions. Chemically, water acts as a solvent, as a reactant in all important reactions in the weathering zone, a chemical constituent of many secondary minerals, especially clays and hydroxides, and as a chemical buffer fixing the thermodynamic activity of soil solution at one, or close to it, except in arid systems and episodes such as in the formation of sal/sodic soils. The ubiquitous presence and importance of organic matter. Organic materials may range from total dominance in the system as in ombrotrophic peatlands to a minor or insignificant role in very young or incipient soils such as lithosols. In the A horizon of a soil, organic matter undergoes decay, and organic acids and complexants are produced. These promote the weathering of inorganic constituents and the mobilization of inorganic ions in solution. The concentration of organic material attenuates with depth, and is generally taken to be negligible, or at least not obvious in the C horizon of a soil. The nature of the inorganic substrate. The dominant minerals in virtually all soils are silicate or aluminosilicate in composition. Over time they tend to be stripped of the so called basic cations, and to be converted into some combination of quartz, 1:1 sheet silicates, and hydroxides of Al and Fe, thereby producing acid soils of low ISSN: 0928-2025
DOI: 10.1016/S0928-2025(06)09008-0
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cation exchange capacity. If calcite is present, it reacts more rapidly than the (alumino)silicates and a pH between 7 and 8 is established. A rarer occurrence is the presence of sulfides, pyrite being the most likely. It too weathers more rapidly than the (alumino)silicates, and gives rise to waters with a pH of 3 or less. Of the elemental components in the inorganic substrate, Fe plays an extremely important role. It is the most abundant element with more than one valence, and takes part in reactions between the soil solution and ferric-hydroxy phases that relatively quickly set the overall redox state of the soil. The physical nature of the substrate is also of importance, particularly its hydrodynamic properties. The latter determine rates of drainage of the material and hence the degree of saturation with respect to water. The partial pressure of O2. Initial values (20 kPa) are set by atmospheric O2. In the vadose zone of the weathering system it will be essentially the same near the surface, but with depth and the increased biological oxygen demand of aerobic decay of organic matter the value will be less, and as microbial organisms proliferate and use up all the available resource of uncombined oxygen, will tend to reach the limit where PO2 is equal to PH 2 , and the system as a whole changes from aerobic to anaerobic. This limit will normally coincide with the water table, below which any pore space will be saturated with water to the exclusion of all oxygen except the small amounts actually dissolved in the water right at the water table. In some cases, lateral inflow of oxygenated water, may result in anomalously high levels of oxygen within the saturated zone. The partial pressure of CO2. Again the initial value is set by the atmospheric partial pressure: 101.5 kPa (103.5 atm). Respiration in the root zone, and decay of organic matter may raise this up to a maximum of about 100 times the atmospheric value. PCO2 determines the pH of the H2O–CO2 (carbonic acid) system, with 5.7 as the common value in unpolluted rainwater. In the absence of organic acids, carbonic acid is the most important one in acid–base reactions in the weathering mantle. This means that its chemical importance in soil tends to increase with depth as the importance of the organic component diminishes. Clearly there is ample opportunity for feedback among these factors, and taken together, they determine that the commonest types of chemical reaction within a weathering system are of acid–base (including cation-exchange) and redox type. Consequently, serviceable choices of master variables for the soil environment are those that are related to the thermodynamic activities of the proton and the electron. Regarding the activity of protons in a system, the universal variable of choice is pH. Regarding electrons, the equivalent parameter pe may be chosen, but the classic variable employed is Eh, which has the advantage of being measurable both in laboratory and field. As shown in Figure 8.1, there is a simple relationship between pe and pH. The application of this type of diagram to weathering systems generally and to peatlands in particular will now be described.
Physico-chemical background Figure 8.1 is variously referred to as a redox–pH diagram, an Eh–pH diagram, a pe–pH diagram, or simply a Pourbaix diagram, after the metallurgist who devised this
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Figure 8.1. Pourbaix diagram. The sloping, dashed line separating oxidizing from reduced conditions is for PH 2 ¼ PO2 in the equilibrium H2O ¼ H2+0.5 O2.
type of Cartesian graph for the purposes of examining corrosion (Pourbaix, 1974). In earth science their use was pioneered by Krumbein and Garrels (1952), and a good introduction is provided by Garrels and Christ (1965). A compilation of basic diagrams useful to the geochemist may be found in Brookins (1988). Truesdell (1968) suggested the adoption of pe rather than Eh as the redox parameter as a means of simplifying the calculation of redox–pH diagrams from fundamental thermodynamic data. Berner (1970) details a number of practical limitations related to the construction and interpretation of Eh–pH diagrams. At the most fundamental level, there is the problem of the quality of the thermodynamic data upon which the diagrams are based. In this regard, Woods and Garrels (1987, p. ix) refer to ‘‘the perils of indiscriminate selection of values from various sources’’. A major difficulty, and one which clearly has a determining influence on any ultimate conclusions to be drawn from the diagrams, lies in choosing what minerals to show on them, whether to consider metastable phases, whether to consider solid solution in condensed phases, and what to do about bacterial reductions. Furthermore, it is assumed that equilibrium (stable or metastable) holds for a particular diagram constructed. Finally, there is the major difficulty of obtaining good measurements of Eh in the field (Shotyk, 1988). In spite of all this, a consensus exists that the diagrams are useful in indicating general conditions and general tendencies in a system. Even for more parochial considerations, redox–pH diagrams may prove useful provided that matters of scale, and the possibilities of local reactions are kept in mind.
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Figure 8.2. Predominance fields in the system Fe–Ca–K–S–CO2–H2O. Conditions: 100 kPa total pressure, 251C, activities of all components except S, 106. S is undefined in terms of a specific activity in order to show the sulfide field at its maximum extent.
Predominance fields Figure 8.2 is a redox–pH diagram of the type used in the geochemical interpretation of low-pressure/temperature environments. Areas called predominance fields are labeled in terms of the minerals that would be expected to form under the Eh and pH combinations of those fields. The actual positions of the boundaries between fields depend on the choice of ionic activities made in calculating the appropriate mineralogical reactions for the system. A total pressure of 100 kPa and a standard temperature of 25 oC are usually chosen as representative ambient conditions. This diagram is of special interest in the context of the geochemistry of peatlands, as will be made clear later.
Geochemical fences A notable contribution of Krumbein and Garrels (1952) to the usefulness of the Eh–pH approach in earth science is the concept of the geochemical fence: a relatively narrow, linear zone in redox–pH space defined by a specific chemical reaction or by a related group of reactions. Their original diagram, developed to help in the
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Figure 8.3. The original system of geochemical fences devised by Krumbein and Garrels (1952).
interpretation of sedimentary environments, is shown in Figure 8.3. It requires modification if the full scope of soil conditions are to be covered, especially below pH 7. In practical terms this amounts to changing existing geochemical fences, and adding new ones. The modified fences, shown in Figure 8.4, are: Organic fence. Within the upper organic-rich part of the solum, microbially mediated breakdown of organic matter will give a range of possible Eh values for a given pH, up to the lower limit of oxidation by O2 in soil. The latter will likely coincide with the water table in the soil system as a whole, with the exclusion of oxygen to microsystems being controlled by details of micromorphology, for example by the presence of clay cutans protecting a mineral surface from oxidation. Above the water table the microbial biomass, using dead organic matter as its carbon source, will generally experience exponential growth up to the point where free O2 is fully utilized. Again, this will obviously coincide with the lower limit of oxidation by O2. It should therefore be placed at the Eh–pH contour along which PO2 is equal to PH 2 in the gaseous dissociation reaction of H2O, rather than where Eh is zero, the value chosen by Garrels and Christ (1965). In theory, the upper limit would be marked by the partial pressure of oxygen in the atmosphere. In practice, the biological oxygen demand in virtually all soils depletes oxygen to lower levels than this.
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Figure 8.4. The pedogenic grid (after Chesworth 1992), an extension of the Krumbein and Garrells diagram (Fig. 8.3) to include new fences and fences modified from the diagram. The extended diagram is applicable to pH values lower than Krumbein and Garrells consider, including values found in bogs, many mires, and acid soils (see Figs 8.9, 8.11, 8.13).
The width of the organic fence shown in Figure 8.7 is consistent with measurements in such materials. Neutral fence. Garrels and Christ (1965) give no mineralogical justification for a neutral fence, and in the ubiquitous presence of atmospheric and respiratory CO2 in a weathering system, it is unlikely that the ionic dissociation of water will buffer a natural system to a pH of 7. However, an average smectite (Weaver and Pollard, 1973), with so-called basic cations (Mg, Ca, Na, K) in exchange positions, marks a near-neutral fence according to the reaction: Mg0:2 ðSi3:81 Al1:71 Mg0:29 O10 ðOHÞ2 ðmontmorilloniteÞ þ 0:98 Hþ þ 0:22e þ 0:33H2 O ¼ 0:86 Al2 Si2 O5 ðOHÞ4 ðkaoliniteÞ þ 0:22 FeOOH ðgoethiteÞ þ 2:1SiO2 ðquartzÞ þ 0:49Mg2þ Because of minor substitution of Fe in the octahedral layer it will not be a vertical fence as in the Garrels and Christ diagram (Garrels and Christ, 1965). It should also be realized that this particular fence does not mark the lower pH limit of a smectite predominance field, since there exists a spectrum of smectites down to beidellite that is stable at low pH. Fe hydroxide fence. Garrels and Christ (1965) based their Fe-oxide fence on the solubility of Fe2O3. In the weathering regime Fe-hydroxy phases are overwhelmingly
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more important. Goethite is generally accepted as the stable hydroxide, though ferrihydrite is probably its precursor. The new fences in Figure 8.4 are : An aluminosilicate fence. This is defined by reactions in the system SiO2– Al2O3–H2O. For example by the congruent dissolution of kaolinite: Al2 Si2 O5 ðOHÞ4 ðkaoliniteÞ þ 6Hþ ¼ 2Al3þ þ 2SiO2 ðquartzÞ þ 5H2 O ðpH4SiO4 o4:7Þ by the incongruent dissolution of kaolinite: Al2 Si2 O5 ðOHÞ4 ðkaoliniteÞ þ 5H2 O ¼ 2AlðOHÞ3 ðgibbsiteÞ þ 2H4 SiO4 ðquartzÞ ðpH4SiO4 o4:7Þ and by dissolution of gibbsite: AlðOHÞ3 ðgibbsiteÞ þ 3Hþ ¼ Al3þ þ 3H2 O A sodium carbonate fence. This defines the high-pH extreme of the soil environment, and lies in the pH range of feldspar dissolution, if the latter were to reach a saturated equilibrium, as well as the range in which zeolites form authigenically under earth-surface conditions. In addition, two fences on the original Krumbein and Garrels diagram are left essentially in their original form. These are: The limestone fence. Renamed the calcite fence, on the high pH side of which solid calcite persists in a soil system. The sulfate– sulfide fence. Name simplified to sulfide fence, marking the upper limit of Eh of the field wherein sulfides, and in particular pyrite, may form in a soil. Other geochemical fences may be useful in specific or local situations. For example, in ombrotrophic bogs with low-atmospheric input of particulate matter, the lowpH limit could be marked by a fence defined by the dissociation of H+ from carboxylic functional groups in the organic component of the soil system. An even lower pH fence can be defined for the interpretation of acid-sulfate soils by the breakdown products of pyrite. Where neither organic matter nor pyrite is common, the system H2O–CO2 will be useful in defining a fence from about pH 5.7 to 5.5 depending upon the balance between atmospheric and respiratory inputs. Under extreme reduction a CH4 fence comes into play.
The pedogenic grid The resulting adaptation (Fig. 8.4) of the Krumbein and Garrels (1952) diagram is referred to as the pedogenic grid (Chesworth, 1992) and covers the broad range of conditions in common mineral soils. Certain organic soils and peatlands, and less common soils such as acid-sulfate soils, lie outside this range. Parallels exist between the pedogenic grid and the pressure–temperature (PT) diagram of the petrologist. The invariant point of a PT diagram, where pressure and temperature are held constant, is similar to those regions on the grid where two
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fences cross. Chesworth (1992) called such crossing points nodes. They act as quasiinvariant points where the pH and Eh of a system is buffered and poised, respectively. A further analogy arises from the fact that the essentially univariant (or pseudounivariant) fences of the grid delimit separate ranges of conditions with each area being characterized by its own individual assemblage of minerals. This is implicit in the diagram of Krumbein and Garrels (Fig. 8.3).
Geochemical trends in the weathering zone Three general trends are followed in the geochemical evolution of surface materials in the weathering zone (Fig. 8.5). Essentially they are the result of the pumping of protons and electrons between sources and sinks in the land environment. Proton and electron pumps A consequence of considering the activities of the proton (H+) and electron (e) as the major driving forces in effecting chemical change in soils (H+ in hydrolysis, and acid-base reactions generally, e in oxidation–reduction reactions) is that conceptually both can be considered as being pumped between a source (or sources) and sink (or sinks) in the weathering regime. First consider H+. On the largest scale, weathering may be described as the titration of acids from atmospheric and organic sources, against bases represented by the aluminosilicate, carbonate, and other mineralogical constituents of the earth’s surface. Figure 8.6 illustrates this within Eh–pH space as a proton pump between source and sink. In a humid climate, with adequate atmospheric precipitation, the land surface is ‘‘inevitably over-titrated, acting as a sink for protons’’ (Edmond et al., 1979, p. 21). Given enough time, all so-called base cations (Na, K, Ca, Mg) will be stripped from solid phases and removed by leaching. Starting from calcareous materials, e.g., solution of calcite, would buffer the system between pH 7 and 8 until it was completely dissolved. Then progressive overtitration would lead to hydrolysis and the formation of a residuum made up of such minerals as kaolinite, gibbsite, and goethite, together with resistant minerals, of which quartz is the commonest. Eventually (and with the proviso that no metal sulfides were present) a steady (or quasi-steady) state would be set up, buffered by the aluminosilicate fence of the pedogenic grid. How close a given weathering system will approach this steady state will depend largely on the availability of water, and the hydrodynamics of the landscape. A close approach is favored by high rainfall in warm climates, and with a good drainage to carry away dissolved cations. In addition, the persistence of phases such as calcite and smectites saturated with the so-called basic cations Ca, Mg, Na, and K will buffer pH and retard the progress of acidification. In humid conditions, and taking the long-term geological perspective, this can only be considered a transient state, the lifespan of which will depend on such factors as landscape position (shorter lifespan on slopes and topographic highs, a longer one in less well-drained topographic lows)
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Figure 8.5. The three major geochemical trends in the weathering zone on the land surface of the Earth. The arrows indicate an acid trend characteristic of regions of humid climate, an alkaline trend characteristic of dry climates, and a reduced (or hydromorphic) trend found, for example, under waterlogged conditions in peatlands and gleysolic soils.
and texture of the weathering materials, coarse, well-drained deposits leading to a shorter persistence of the basic minerals, than finer, less well-drained ones. Semi-arid and arid conditions will lead to an opposite trend; that is, a trend of undertitration wherein it is the basis in the system that predominates over the additions of H+, and the system is buffered at an alkaline pH. The development of sodicity, and in the limit the precipitation of salts of the alkali metals, is the likely result. Figure 8.7 is a similar interpretation regarding e. The most reduced material in the biosphere is the living biomass. When death and decay intervene, organic materials become a prolific source of electrons in soil and in effect pump electrons to any available sink, the largest of which is atmospheric oxygen. For Malthusian reasons (the tendency of organisms to use resources to the maximum), any resource they depend upon will normally mean that the overall Eh of the soil will be poised along the contour in Figure 8.1 that represents this limit (PO2 ¼ PH 2 ). For example, if the ready diffusion of oxygen into the soil is impeded, for reasons of water logging, heavy texture, or compaction , microbial activity will be fuelled by
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Figure 8.6. The proton pump, showing the movement of hydrogen ions from sources such as dissociating carboxylic groups in decaying organic matter, and from carbonic acid from atmospheric precipitations, to sinks represented by the inorganic materials of the Earth surface, ranging from carbonate to silicate materials. In effect, this is a natural titration, which in humid climates leads to overtitration (and low pH values at the Earth surface) and in dry climates to undertitration (and higher values at the Earth surface).
the movement of electrons from the organic matter to some other available sink. The range of operation of these alternative electron acceptors is well known, and is shown in Figure 8.8. At finer scales, particularly in the contact zones between phases, and on the reactive interface between solid and soil solution, an immense number of transfers of H+ and e are possible. On the whole, however, and on the macroscopic scale of the solum presented here, the general tendencies outlined are reasonable ones, at least semi-quantitatively. The peatland environment From the geochemical standpoint of this chapter, the most significant differences within the peatland environment are shown within the framework of the system Fe–Ca–K–S–CO2–H2O (Fig. 8.2). Our interpretation is based on the following observations developed from the foregoing discussion. Presence of a water table divides any peatland into an oxidized and a reduced zone. This of course is recognized in the acrotelm/catotelm vertical stratification in mires.
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Figure 8.7. The electron pump. This represents the overall movement of electrons at the macroscale of a bog, mire, or soil landscape from the principal source in decaying organic matter, to the principal sink of atmospheric oxygen. In a decaying peat, the oxidation state would rise in the direction of the heavy arrows. Oxidation of a peat deposit in the water-unsaturated acrotelm, would lead to conditions above the PO2 ¼ PH 2 contour, and as oxidation proceeded PO2 ¼ 100 kPa contours. Conditions in the watersaturated catotelm would normally rise no further than the PO2 ¼ PH 2 contour, unless there was an influx of oxygenated water. However, biological oxygen demand would determine that the oxygenated state did not persist.
The predominance of organic matter means that it is the principal source of H+ for acid–base reactions. The definitive distinction between ombrotrophic and minerotrophic peatlands is the addition of extraneous materials only via the atmosphere in the former case, and additionally (and more copiously) via ground and surface waters in the latter case. Ombrotrophic mires The geochemistry of the ombrotrophic environment is the simpler of the two types. Living and decaying vegetation (Sphagnum, Carex, and other genera) dominates the environment and pH is determined basically by the dissociation of carboxylic functional groups in the dead organic matter. Since the rate of addition of terrigenous minerals and the rate at which many of them dissolve in the acidic bog waters are too
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Figure 8.8. Several sources and sinks of electrons in the land surface. At a microscale in a bog, mire, or soil landscape, the movement of electrons is more complicated than the big picture shown in Figure 8.7. Many more sources (for example, ferrous iron, sulfides, native metals ) and sinks (for example, manganous oxides and hydroxides, and nitrates) exist, and will influence the detailed geochemistry of a natural system, but in a peatland the overall geochemistry will be dominated by the reactions between organic matter and oxygen as shown in Figure 8.7.
slow to consume the protons generated by CO2 and organic acids the pH can be expected to remain constant with depth as shown in Figure 8.9. Eh will vary independently, and in the catotelm may reach levels of reduction low enough to allow the generation of CH4. It is possible that oxygenated groundwater may enter the system below the water table, and rise up in a plume, but in general oxygen is unlikely to be the terminal electron acceptor within the zone of water saturation. The fate of the scarce wind-borne additions to ombrotrophic bogs have been discussed by Martı´ nez Cortizas et al. (2001a, b) in the context of the predominantly quartzo-feldpathic additions to the wetlands of Galicia, northwestern Spain (Fig. 8.10). The material is assumed to be added in amounts too small to have an effect on the overall pH of the environment.
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Figure 8.9. The ombrotrophic environment. The field of ombrotrophic peatlands will be approximately as shown in the gray, cigar-shaped area, with the relatively low pH determined by the dissociation of carboxyllic acid groups in organic matter. Generally, eolian inputs of inorganic dust will be too feeble to move the pH to higher values.
Minerotrophic mires The increased importance of extraneous additions to minerotrophic as opposed to ombrotrophic peatlands shows up in two ways. First, the addition of Fe-containing minerals (ferromagnesians, oxides or hydroxides) and their reaction with water tends to produce a redox profile that follows a contour parallel to the boundary of the ferric-hydroxide field of predominance. Second, hydrolysis of added minerals will raise pH, most dramatically if calcite or dolomite is in the mix. The boundary of the predominance field of calcite will provide an upper limit in this case, at least as long as carbonates persist. Where only primary silicates are added (and the most common will be feldspar) a raised pH will be more slowly established and will not normally reach the calcite field. The possibilities are shown in Figure 8.11. Discussion Mires, mainly ombrotrophic ones, are extensively used as archives of compounds deposited by wet and dry deposition from the atmosphere, heavy metals in particular.
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Figure 8.10. Reaction products of terrigenous dust in the ombrotrophic environment of Northwestern Spain, showing the local suite of input minerals, the PH 2 and PO2 ranges of the local blanket bogs and the resulting reaction products in six segments of the PH 2 PO2 framework, shown as mineral facies diagrams based on a triangular projection of the system Si–Al–Fe–O–H–S.
The basic idea underlying this kind of research is that dust and other compounds deposited on the mire will eventually be buried as plant remains accumulate. Any subsequent core may then be sectioned and analyzed to give a high-resolution record of changes in atmospheric deposition for up to a few thousand years. Elemental mobilities in the peat are a potential complicating factor, though most studies assume that elements of interest are immobile. Few papers actually address the geochemical conditions of the mire in terms of their effect on the stability of the mineral phases hosting an element (Shotyk, 1995; Steinmann and Shotyk, 1997) and thereby on the possibility of that element being released in mobile form. Four aspects are of particular interest in this regard: (1) the mineralogical composition of the deposited dust, (2) the geochemical environments of the mire (the acidic, oxidized acrotelm and the slightly less acidic but anoxic catotelm), (3) the rate of peat accretion, and (4) the interference of human activities on the natural composition of the atmospheric dust and aerosols. Materials deposited from the atmosphere onto the surface of the mire are incorporated initially into the acid/oxidized environment of the acrotelm. Most primary minerals are unstable in this environment, and at the extreme (pH below 4), most
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Figure 8.11. The minerotrophic environment. The lower pH values or the cigar-shaped field of the minerotrophic mines are defined by ferrous–ferric equilibria that partially delimit the predominance field of ferrihydrite, a precursor of the stable-phase goethite, shown in Figure 8.10, and formed as an initial product of the breakdown of ferromagnesians or magnetite. Depending on the amount and type of inorganic crustal material present in a mire, the upper pH limit of the minerotrophic environment will extend in the direction of the heavy arrows, towards the calcite-present area.
common minerals, primary or secondary, are unstable. Only if a given element is hosted in stable minerals may its concentration in the peat be assumed to reflect that of the original dust (though correction for the effect of the organic substrate may be necessary). Problems arise when the inorganic phase breaks down and an element is released as an ion, into solution. Once in solution it may leach (essentially through lateral movement) or diffuse through the mire water column following concentration gradients. Figure 8.12 exemplifies this by showing the profile of Fe concentrations in an ombrotrophic mire. The highest Fe concentrations occur in the aerated acrotelm of the upper 25 cm of the bog; below this, concentrations abruptly decrease to reach minimum values where the fluctuation of the water table takes place. Below this level (and particularly below 110 cm in the example shown) values increase again. A similar pattern was obtained in an experiment where cores from uncontaminated sites were transplanted to contaminated sites, and vice versa. Iron concentrations in the peat core of the contaminated mire decreased after being transplanted to the uncontaminated site, whereas the opposite was found for the peat core of the uncontaminated mire after it was transplanted to the contaminated one (Martin Novak, personal communication). This indicates that Fe was mobilized in or out of the
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Figure 8.12. Profile of Fe (ppm) concentrations in an ombrotrophic bog (Pena da Cadela, Xistral Mountains, NW Spain). The vertical scale is in centimeters, the horizontal scale of the Fe profile is in parts per million. Such a profile might be interpreted as being due to a greater influx of Fe in more recent dusts than in older ones (represented by deeper parts of the bog). However, in the upper 50 cm in which Fe varies 2.3-fold, normalization against a known immobile element (Ti) shows a variation of 3.2-fold. The fact that the two ranges are different is prima facie evidence that Fe is mobile. Extrapolating to the total profile suggests the possibility that Fe has also been mobilized at depth, a possibility supported by the transplant experiments described in the text.
transplanted cores, following concentration gradients. Using a chemical speciation approach, Shotyk (1995) also suggested that there is no obvious chemical mechanism to retard the possible migration of As and Sb. However in a recent investigation of a peat core from a blanket bog from the Faroe Islands it was found that Sb and Pb
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records were highly correlated, and this correlation was taken as evidence of effective Sb immobilization (W. Shotyk, personal communication). Other authors have also indicated that changes in redox potential, dependent on water table fluctuations, are the most important mechanisms in the distribution of heavy metals (Damman, 1978; Clymo and Hayward, 1982; Damman et al., 1992). Steinmann and Shotyk (1997), on the other hand, found relatively high Fe3+ to Fe2+ ratios in bog-pore waters, despite the anoxic conditions, which they explained in terms of the higher thermodynamic stability of the organic complexes of Fe3+ compared to those of Fe2+. Unlike the case of a mineral soil, in ombrotrophic mires released elements are unlikely to fix by complexation to another inorganic phase, since no more than 2% of the total dry mass of the peat is inorganic matter. However, released elements may obviously become bound to organic matter, for example, Hg shows a marked affinity for humified organic matter (Benoit et al., 1994), with binding to reduced sulfur groups being preferred (Skyllberg et al., 2003). Complexation is the most important mechanism for metal fixation, as Shotyk (1995) has shown. He calculated that in bogs 99.9% of Cu, Pb, and Zn can be bound to an organic component, either in the solid phase or in solution (as dissolved organic carbon, DOC). So, organically bound elements can be retained by the peat or they can be leached from the system with the dissolved organic matter. In other words, the extent to which the original concentration of a given element will be changed subsequent to the addition of that element to a mire will depend on the degree to which the element is partitioned between the immobile solid phase and the mobile dissolved phase. This in turn will depend upon the kinetics of the reactions that produce the soluble chemical species. Finally, the hydrodynamics of the system plays an important role. If the mobile phase is definitively trapped within the mire, no compositional change will take place. If, as seems most likely, the mobile phase finds an exit route, then elements complexed to DOC will leave the system. In minerogenic mires the higher mineral content makes the role of the inorganic compounds more relevant in terms of the redistribution of metals. For example, Franze´n et al. (2004) analyzed a minerogenic mire and found enrichments in Hg associated to elevated Fe concentrations and the presence of goethite, suggesting Hg adsorption to Fe-oxyhydroxides probably as Hg-fulvo-acid complexes. One important aspect of the fate of organically bound elements is that their potential release is coupled to that of dissolved organic matter (DOM) and linked to climate changes. As indicated by Biester et al. (2006 – this book, Ch. 19) the release of organo-halogens is expected to increase during warm and wet periods and decrease during dry periods; this may also affect other organically bound elements (that is, metal-DOM complexes). Continuous growth of vegetation in the surface of the mire tends to keep the acrotelm at an almost constant thickness (Belyea and Clymo, 2001). By contrast, the catotelm grows upwards. This means that any given peat surface will reach the catotelm in a few hundreds to a thousand years (Martı´ nez Cortizas et al., 2001a, b). Thus minerals/elements progressively sink into a more anoxic environment as time goes by. Vertical water movement is largely impeded in the catotelm (Ingram, 1983)
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and the fate of an element in this section will mostly be governed by redox reactions and diffusion. Human activities can affect the amount and composition of the deposited dust in mires. Induced soil erosion due to forest clearance, fires, and agriculture have increased the amount of atmospheric dust and changed its composition (by increasing the surface area and types of exposed sources: soils and rocks) especially since the so called Neolithic Revolution; that is, from 10,000 to 5,000 years ago (Martı´ nez Cortizas et al., 2005). The most prominent changes in atmospheric composition occurred once the smelting of metals began. In the last 250 years, the changes have accelerated as a consequence of the Agricultural and Industrial revolution. Metallurgy, mining, fuel burning, waste incineration, and other human activities have released to the atmosphere compounds that are markedly different from those originating from soils and rocks. As with dusts of natural origin, the anthropogenic material is also subject to possible modification in the mire-environment. For example, Sonke et al. (2002) investigated metal pollution in a bog close to a Zn-smelter in the Kempen region (Belgium) and found that though oxidized smelter dusts were found in the nearby mineral soils, the bog showed extensive in situ FeS2 and ZnS, indicating that anoxic conditions were responsible for rapid diagenetic transformations of anthropogenic metal oxides into sulfides. Rausch et al. (2005a, b) investigated solid-phase and pore-water concentrations of some trace metals in uncontaminated and contaminated sites in Finland and found that Cu was generally retained more effectively than Ni, Co, Zn, or Cd, but in highly contaminated sites even Cu was substantially transferred from the solid phase to the pore waters. These authors indicate that ombrotrophic bogs may or may not function as reliable archives of atmospheric trace metals, depending on the concentration and chemical properties of the elements considered, the mineralogical form at the time of deposition, as well as the pH of the bog waters. As a general rule, they hypothesize that sites that are remote from point-pollution sources and receive metals exclusively by atmospheric deposition are more conducive to preserving the chronology of deposition. Not only element concentrations can be affected by ambient geochemical conditions, isotope ratios may also be modified. The bulk isotopic composition of peat arises from a mixture of sources (for example, soil dust, anthropogenic pollutants), which can have different isotope ratios. This characteristic is used, for instance, in tracing the source of Pb pollution, since the isotopic signature of Pb from an anthropogenic source such as gasoline, differs from that of Pb in soil dust. However, there may be considerable heterogeneity in the Pb from different soil pools (Emmanuel and Erel, 2002; Wong and Li, 2004). Consequently the materials added to a mire, whether they are from natural or from anthropogenic sources, and whether they are from finer or coarser fractions, from the primary mineral or secondary mineral, or other fractions of the soil, will show a range of geochemical stabilities in the peat environment. This will complicate attempts to decipher the geochemical history of a peatland. In addition, if the hydrodynamics of a given system is conducive to the leaching of the more labile weathering products, post-depositional changes of the original composition, including its bulk isotope ratio become possible, and use of the peat as an archive is compromised. The danger is that these changes may be interpreted as changes in Pb atmospheric composition through time.
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Conclusions Redox–pH diagrams provide a concise graphical way in which to depict Earthsurface environments in terms of the activities of electrons, the operators involved in the two commonest types of reaction in the weathering zone on land, i.e., redox and acid–base reactions. Figure 8.13 illustrates the following points by way of summary. Peatlands tend to fall on the acid side of the diagrams with ombrotrophic bogs showing least overlap with common soil-forming environments. Minerotrophic mires overlap the field of ombrotrophic bogs under relatively high redox conditions, though with depth they move along a line defined by ferrous–ferric equilibria, to a near-neutral pH. The same kind of equilibria also plays a controlling influence on ambient conditions in mineral soils at the low limit of their pH range. The upper limit of the minerotrophic field in peatlands, however, will depend on whether or not there are extraneous additions of carbonates. If there are ambient conditions, the peatland may extend to the calcite predominance field, where a pH approaching 8 is possible. Within this extended field lies the environment of histosol formation. Histosols may be looked upon as a kind of bridge between peatlands and mineral soils.
Figure 8.13. The geochemical environments of bogs and mires shown in the context of the Eh–pH range of common soils and of acid-sulfate soils.
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In the context of this volume, the most important conclusion is that the peatland environment covers a wide range of geochemical conditions and encompasses those under which most primary, and many secondary minerals would break down. If in fact inorganic additions to a peatland do break down, the possibility of the mobilization of the elemental components must be entertained. Consequently redox–pH diagrams offer a first line of defense against the hasty use of any given proxy element assumed to be immobile, in the reconstruction of past environmental conditions.
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