Three major pools of iron can be distinguished in soils: (i) iron included in primary and secondary minerals, (ii) soluble iron (and the speciation of iron, i.e., the range of iron species that occur in solutions), and (iii) iron bound to organic matter. Iron availability in soils and rhizo- spheres is governed by its concentration in the soil solution and more importantly by its partitioning within the solid phases, and the ability of the latter to replenish the soil solution via dissolution/precipitation and dissociation/association of complex species.
2.1. Pools of iron in minerals
Iron is the fourth most abundant element in the earth’s crust after oxygen, silicon, and aluminium (Louet, 1986; Ma, 2005). Commonly, iron occurs in two oxidation states in soil minerals, designated as ferrous (II) and ferric (III) iron. These are found in most rocks in a range of primary minerals, and mostly in ferromagnesian silicates such as olivine, augite, hornblende, and biotite (Louet, 1986; Schwertmann and Taylor, 1989; Segalen, 1964). Pri- mary soil minerals, containing principally Fe(II), are generally unstable in soils and thus weather in the presence of water and atmospheric oxygen. The weathering of these minerals can be dramatically accelerated by the activities of living organisms, as shown for microorganisms (e.g.,Brantleyet al., 1999) and plants (e.g.,Hinsingeret al., 2001). During weathering, Fe(II) and Fe(III) ions are released by a range of dissolution and oxidation/reduction mechan- isms. However, in the presence of hydroxyl ions (OH), they rapidly hydrolyze and precipitate to form a range of poorly soluble secondary
Iron Dynamics in the Rhizosphere 187
minerals such as iron oxides (hematite, magnetite) and oxyhydroxides (goe- thite, lepidocrocite), as well as iron hydroxides and less organized minerals such as ferrihydrite (Cornell and Schwertmann, 2003; Segalen, 1964). Iron oxides may be associated with clays and sorb ions such as phosphate, leading to a decreased solubility of iron (Borggaard, 2002; Borggaard et al., 1990;
Hinsinger, 2001; Segalen, 1964). Depending on soil aeration, the predomi- nant iron oxides contain either Fe(II) or Fe(III), producing the characteristic colors of soils (Cornell and Schwertmann, 2003; Lindsay, 1979; Segalen, 1964). In oxic environments corresponding to the majority of cultivated soils, iron is mostly found as Fe(III) oxides, leading to soil colors varying between brown and yellow–red, whereas in reducing environments, such as waterlogged soils, marsh, and flooded zones, Fe(II) soluble species and oxides are predominant, giving rise to soils colors varying between grey and blue–
green. Mixed Fe(II)–Fe(III) hydroxides, commonly called green rusts such as fougerite, largely explain these colors of waterlogged soils or soil horizons (Federet al., 2005). In the most reducing soil conditions, ferrous iron sulfides (pyrite, FeS2) may ultimately form. Redox processes are reversible and these ferrous iron-bearing minerals may rapidly dissolve in the presence of oxygen, while iron oxides may dissolve in the presence of reducing compounds.
Other iron-bearing secondary minerals, ascarbonates (siderites, FeCO3), and more importantly phyllosilicates (clay minerals such as glauconite, which are iron-rich illites, and some smectites such as beidellite) are also present in soils (Louet, 1986; Schwertmann and Taylor, 1989).
Besides iron-bearing minerals, iron can be bound to mineral surfaces, especially clay minerals that are prone to adsorb metal cations in an exchangeable pool due to their cation exchange capacity. Because of the small concentrations of iron in soil solution in oxic environments (Fig. 1), iron usually makes up only a minor proportion of exchangeable cations.
This proportion, however, may increase under acidic conditions and more so under reducing conditions.
Progresses have been made in the knowledge of iron-bearing minerals and their surface reactivity over the years thanks to the fast development of spectroscopic methods (Mo¨ssbauer spectroscopy, synchrotron-based X-ray spectroscopies) and high resolution microscopic techniques (high resolution transmission electron microscopy and atomic force microscopy). The above-referred progresses were mostly achieved for transition metals and their metal oxides, but less for pure iron oxides [as reviewed by, for example,Fordet al.(2001)].
2.2. Solubility of iron oxides
In aerated soils, iron-bearing secondary minerals mostly contain Fe(III).
Their solubility is controlled by dissolution–precipitation equilibria such as the following for, for example, iron hydroxide: Fe(OH)3⇆ Fe3þþ3OH
188 A. Robinet al.
(Table 1). This reaction is fully dependent on soil solution pH as shown in Fig. 1 for the case of a ubiquitous iron oxyhydroxide in soils (goethite), according to: FeOOHþ3Hþ⇆Fe3þþ2H2O. For this mineral, which is less soluble than Fe(OH)3 and the so-called ‘‘soil iron oxide’’ in Lindsay
Table 1 Information on parameters related to the iron status in soils and rhizospheres
Parameter
Stability constant
and concentration Reference
Solubility of iron Lindsay (1974, 1979);
Lindsay and Schwab (1982 )
Fe3þþ3OH!Fe(OH)3
KspFe(OH)3ẳ[Fe3ỵ] [OH]3 Kspẳ1038M
[Fe3ỵ]ẳ107M (at pH 3.5) [Fe3ỵ]ẳ1023M
(at pH 8.5)
Stability constants with iron
Enterobactin 1052 Pollack and Neilands
(1970)
Ferrioxamine 1032 Berneret al.(1988)
Pyoverdine 1032 Meyer and Abdallah
(1978)
Fusarinine 1029 Scher and Baker (1982)
Phytosiderophore 1018 Sugiuraet al.(1981)
Organic acids Jones (1998)
Malate 107
Citrate 1011
Oxalate 108
EDDHA 1033 Lindsay (1979)
EDTA 1025 Lindsay (1979)
Humic acid 1013 Takahashiet al.(1997)
Concentrations in rhizospheres
Siderophores 107to 108M Powellet al.(1980) Phytosiderophores 103M Ro¨mheld (1991)
Organic acids 5105M to
9103M
Dinkelakeret al.(1989);
Joneset al.(1996) Iron requirement for optimal growth
Plants 104to 109M Guerinot and Ying
(1994)
Microbes 105to 107M Loper and Buyer (1991)
Iron Dynamics in the Rhizosphere 189
(1979), the activity of Fe3ỵranges from about 109M at pHẳ3.5 to about 1024M at pHẳ8.5, these concentrations being lower than the plant and microbe requirements. The solubility diagram presented in Fig. 1 shows that Fe3þis never the dominant Fe(III) species in solution over this range of pH since Fe(OH)2ỵappears to be dominant up to pHẳ8 and Fe(OH)4at higher pH values. This diagram also shows that total soluble iron (sum of all iron species in equilibrium with goethite) reaches a minimum at pH ẳ8, value which is close to that of calcareous soils. Accordingly, the total Fe concentration in the soil solution is dramatically decreased when pH increases up to 8 (Lindsay and Schwab, 1982).
The solubility of iron oxides is also very much dependent on redox conditions. When oxidizing conditions prevail, which is the case in most soils, the activity of Fe(II) species is lower than that of Fe(III) species at pH values commonly found in soils, except possibly in the most alkaline con- ditions (see Fe2þ vs Fe3þinFig. 1) (Lindsay, 1979). Thus, at atmospheric pO2 (peỵpH ẳ20.61), Fe2ỵconcentration is rather negligible (close to 1020M at pHẳ7) and contributes for little to the total soluble iron in soil solution. However, Fe2þ concentration will increase tenfold when the pe drops by 1 unit. Thus, under reducing conditions, when peþpH reaches 9, Fe2þ makes up a major contribution to total soil Fe (its concentration getting close to 109M at pHẳ7). However, this does not apply to oxic conditions, and will then not be further discussed here.
The solubility of iron oxides is also much affected by their particle size.
When the particle size of goethite and hematite decreases, their usually low solubility increases to reach that of the more soluble ferrihydrite (Trolard and Tardy, 1987). This is relevant to soils because goethite (and hematite) particles of nanometric size are ubiquitous in oxic soil environments (Cornell and Schwertmann, 2003). The particle size not only affects iron solubility but also more importantly the kinetics of iron oxide dissolution (Kraemer, 2004).
Besides pH, pe, and particle size, ligands, which make complex species with Fe(II) and Fe(III), influence the dissolution/precipitation equilibrium and kinetics (Fig. 2).
2.3. Complexes and chelates with organic matter
The solubility of soil iron is also considerably affected by the complexation or chelation of Fe3þby organic ligands that make up the DOM. According to Van Hees and Lundstro¨m (2000), more than 95% of Fe in soil solution is likely to be chelated. The corresponding organic ligands belong to microbial siderophores (Neilands, 1981), plant root exudates, including phytosidero- phores (Takagiet al., 1984), and more complex and diverse macromolecules constitutive of humic substances (Stevenson, 1994). The contribution of microorganisms and plants to the solubilization of iron through siderophores
190 A. Robinet al.
and other ligands will be presented in a separate section. Surprisingly, com- pared with other metals such as copper, the binding of iron by DOM or humic substances is much less documented (Tipping et al., 2002). These authors successfully modelled the data fromLiu and Millero (1999)and confirmed that Fe(III) concentration could be increased by about 100-fold (reaching values above 109M) upon addition of humic acid (at concentrations below 1 mg dm3and pHẳ8). Concentrations of Fe(III) above 108M were recorded when adding more than 4 mg dm3at pHẳ8, while micromolar concentra- tions were attained at acidic pH (below 5).Tippinget al.(2002)even showed that the concentration of iron bound to a specific DOM (fulvic acid), at pH values between 6.5 and 8.5, was more than twice that of all inorganic iron species. They also reported that, whatever the pH value was, the species of iron bound to fulvic acid consisted mostly in Fe(OH)2þ, Fe3þ being hardly not bound except at the most acidic pH values.
The high content of oxygen-containing functional groups of humic substances is favorable to the formation of stable complexes with Fe (Chen, 1996; Stevenson, 1994). Binding of iron to humic substances occurs mostly at high soil pH values. These complexes are protected from possible precipitation and subsequent crystal growth processes that would decrease iron solubility (Cesco et al., 2000; Schwertmann, 1991; Varanini and Pinton, 2001). The importance of the complexing capacity of DOM on the availability of micronutrients including iron has long been acknowl- edged (Hodgson, 1969), and humic substances were shown to provide a pool of iron available for plants (Cescoet al., 2002; Chen and Aviad, 1990;
Lobartini and Orioli, 1988; Pintonet al., 1998, 1999; Varanini and Pinton, 1995, 2001). This was illustrated in soilless cultures, whereas contradictory
Log dissolution rate (mol m-2 s-1) -11 -10 -9 -8 -7
2
pH
Protonation Complexation
Reduction
3 4 5 6 7
Figure 2 Diagram showing the impact of different mechanisms—protonation, com- plexation (with oxalate), and reduction (with ascorbate)—on the dissolution of an iron oxide (goethite, FeOOH) as a function of pH. The rate of dissolution as affected by these mechanisms is clearly pH dependent; and appears to decrease as the pH increases (fromStumm and Furrer, 1987, with permission of J. Wiley and sons, NY).
Iron Dynamics in the Rhizosphere 191
information has been reported for soils. Humates were shown to increase the extractability of iron from soils, even in calcareous soils, and to improve iron nutrition (Olmoset al., 1998; Pandeyaet al., 1998). However, in other studies, plant iron uptake was not enhanced by such substances (Alva and Obreza, 1998; Kumar and Prasad, 1989). The discrepancy between the observations made in different experimental conditions may be related to differences in the complexing ability of humic substances in different environments. The stability of Fe–humic complexes is indeed influenced by parameters such as pH—with a maximum stability at pH 8—and the Fe/
humic substances ratio (Garcı´a-Minaet al., 2004). Furthermore, modifica- tion of iron solubility upon introduction of organic matter may also result from indirect effects such as adsorption of plant and microbial siderophores (Crowleyet al., 1991), and many others via the alteration of soil’s physical, chemical, and biological properties (Kraemeret al., 2006).
Humic substances not only contribute to increase iron bioavailability through their iron-chelating properties but also have redox-reactive prop- erties (Weberet al., 2006). These properties are related to phenolic groups contributing to Fe(III) reduction (Chen et al., 2003; Deiana et al., 1995;
Szilaˆgyi, 1971). The chemical reduction of Fe(III) by humic substances is strongly pH dependent, the highest reduction capacities occurring at pHẳ 3 (Chenet al., 2003). As pH increases, humic substances are more frequently bound to metal cations and therefore have a decreased reducing ability (Chenet al., 2003).
2.4. Iron bioavailability
Bioavailable iron can be defined as the portion of total iron that can be easily assimilated by living organisms, according to the general definition of bioavailability given by Harmsen et al. (2005). Iron bioavailability is expected to be governed by the kinetics of iron dissolution from iron- bearing minerals (Kraemer, 2004; Reichardet al., 2005), desorption of iron from exchangeable forms (Crowley et al., 1991), iron speciation in soil solution, and, to a less extent, iron concentration. Chemical analyses of soil iron are commonly based on chemical extraction and quantification, giving access to the exchangeable or extractable iron (Borggaard, 1976;
Cornell and Schwertmann, 2003; Haynes, 1983; Lindsay and Norvell, 1978). These methodologies (Baize, 2000) do not provide information on the associations of iron with soil constituents (solid phases) and even less on the bioavailability of iron. These limitations call for alternative methods to characterize these associations in more detail, particularly for examining the stability kinetics of associations between iron and soil constituents (Bermond et al., 2005). Recent techniques, including the use of stable
192 A. Robinet al.
iron isotopes, bring new prospects for analyzing the behavior of Fe in soils, in particular the interactions between the various reservoirs of Fe in soils (Emmanuelet al., 2005).
None of these chemical methodologies, as common as they are, can replace the final application of living organisms to test their ability to use iron in the soil environment. Iron stress conditions for microorganisms can be assessed in soil indirectly by characterizing their susceptibility to iron starvation. This susceptibility can be evaluated by determining the minimal concentrations of a strong iron chelator (8-hydroxyquinoline) inhibiting the growthin vitroof bacterial isolates from different soil condi- tions. This strategy unlighted the lower susceptibility to iron starvation of populations from rhizosphere than from bulk soil (Lemanceauet al., 1988b), and from the rhizosphere of a transgenic tobacco line overexpressing ferritin than from the rhizosphere of a control nontransformed plant (Robinet al., 2006b), indicating the lower availability of iron in these environments.
Another strategy relies on so-called iron biosensors based on the use of reporter genes which are under the control of the promoters of iron- regulated genes, as those encoding siderophore synthesis in fluorescent pseudomonads (Loper and Lindow, 1997). Such constructs were made in Pseudomonas fluorescens Pf-5 (Loper and Lindow, 1994) and in P. putida WCS358 (Duijff et al., 1994a) by fusing a promoter-less ice-nucleation activity gene (inaZ) to an iron-regulated promoter regulating the produc- tion of fluorescent siderophores. Expression of ice-nucleation activity from this construct is inversely related to the iron concentration. Use of the Pf-5 construct indicated that the bacterial cells were mildly iron-stressed in the rhizosphere (Loper and Lindow, 1994) and that the ice-nucleation activities were similar in different root zones (Marschner and Crowley, 1997). Ice- nucleation activity was shown to decline with time, indicating that iron bioavailability increased during plant growth (Loper and Henkels, 1997).
As stressed byHarmsenet al.(2005), the bioavailability of a nutrient or a toxic substance varies upon the living organisms according to their different acquisition pathways/capabilities. Variation between organisms may also result from a so-called bioinfluenced zone in which each organism interacts with its environment, thereby altering the availability of the nutrient or toxic substance (Harmsenet al., 2005); for plants, this bioinfluenced zone typically corresponds to the rhizosphere. These variations according to the organisms necessarily stress the limits of biosensors based on the use of a given organism for assessing the bioavailability of soil iron, since it would be relevant only to those which are close enough to the species used as biosensor. For instance,Bontideanet al.(2004)designed a bacterial biosen- sor for assessing mercury bioavailability. However, the biosensor response appeared to not match with the bioavailability of mercury to a higher plant species (common bean,Phaseolus vulgaris).
Iron Dynamics in the Rhizosphere 193