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Rhizosphere paper
The rhizosphere as formative in advanced weathering-stage soils
Daniel D. Richter, Neung-Hwan Oh, Ryan Fimmen, & Jason Jackson
Duke University & Yale University
There are not many differences in mental habit more significant than that between thinking in discrete, well defined class concepts and that of thinking in terms of continuity, of infinitely delicate shading of everything into something else, of the overlapping of essences, so that the whole notion of species comes to seem an artifact of thought, not truly applicable to fluency, the so to say universal overlapping of the real world.
A.O. Lovejoy (1936)
Introduction
In many accounts, the rhizosphere is narrowly conceived in space and time. Since first described by Hiltner (1904), the rhizosphere is taken as the soil volume that interacts directly and immediately with living plant roots, an environment that is nanometers to centimeters in radial distance from the root surface. No doubt, rhizospheres are most remarkable microsites with a gaseous, solution, and surface chemistry that greatly affects microbial and plant productivity, nutrition, and physiology. The ready supply of photosynthetically derived organics drives a number of organic geochemical reactions with a variety of inorganic minerals and mineral surfaces.
The rhizosphere is not only an environment that transforms near-root chemistry and greatly affects plants and soil biota. Over time, rhizospheres affect a much larger environment, including much if not all of the so-called “bulk soil” itself. Although not frequently remarked, it can be said that rhizosphere processes have much to do with the ultimate formation of soils.
This chapter examines rhizospheres and their broad biological, physical, and chemical effects on soil formation; it is specifically focused on how rhizospheres contribute to the formation of advanced-weathering stage soils such as Ultisols and Oxisols, extremely weathered soils found in warm temperate zones and the lowland tropics. We hypothesize that rhizospheres, as the dynamic interface between biotic and mineral systems, are critical to the formation of advanced-weathering stage soils such as Ultisols and Oxisols.
In organization, this chapter opens with a discussion of general concepts: of the oft-used dichotomy of rhizosphere vs bulk soil, of rhizospheres as microsites within soil profiles, and of advanced-weathering stage soils. Subsequently, we evaluate physical and chemical effects of rooting on the soil. Throughout, we examine how the biota’s physical and chemical interactions with soils are concentrated in the rhizosphere, and that over pedogenic time these concentrated interactions transform soils across a wide range of spatial scales, from individual mineral grains to entire soil profiles. We conclude that rhizosphere processes are instrumental to the formation of advanced weathering stage soils such as Ultisols and Oxisols, including those with deep C horizons and saprolites.
A Review of Concepts
Rhizosphere vs Bulk Soil
Although variously defined, the rhizosphere is generally taken to be the soil adjacent to actively functioning roots. Compared with the soil as a whole, reactive organic reductants and microbial activity are concentrated in near-root environments. In contrast, the bulk soil is a generally oligotrophic environment with respect to supply of organic matter. Plant root systems are networks within the bulk soil, biological hotspots where respiration and gas exchange, and localized supplies of organic matter are relatively concentrated.
By convention, the rhizosphere has been characterized as having three components:
· rhizoplane, the immediate surface of the root,
· rhizosphere, the soil volume immediately surrounding the rhizoplane that
is directly affected by root activity, and
· bulk soil, the soil not directly affected by living roots.
This tripartite construct helps emphasize the special nature of the rhizosphere, but we suggest overemphasizes the dichotomy between near-root and bulk soils. The construct also appears simplistic given observations from high powered microscopy and results from molecular biology of root systems. The rhizoplane is far from a planer surface, and the radial influence of the rhizosphere is fundamentally ill-defined and ranges widely in spatial scale. Although rhizospheres have been variously understood (e.g., Rovira and Davey 1974), the neat tripartite concept of rhizoplane, rhizosphere, and bulk soil is difficult to align with our developing understanding of the complexity of root systems. Roots systems are best conceived as symbiotic system in which cells of fungi, bacteria, and plants are so intimately associated, both structurally and functionally, that it is difficult or impossible to isolate plant from microbe. The fact that fungi and bacteria colonize root tissues in an “endorhizosphere” suggests a slightly modified tripartite construct of the near-root environment:
· root-microbe system, which includes all cells of plant roots, mycorrhizal fungi, and closely associated non-mycorrhizal fungi and bacteria,
· active rhizosphere, the environment surrounding the root-microbe system which is immediately affected by active functioning of the root-microbe system. The volume of the active rhizosphere is a continuum and dependent on chemical reaction, chemical element, microorganism, and soil type.
· bulk soil, the soil not immediately affected by the functioning of roots, but which may well be transformed by the rhizosphere over pedogenic time.
Much rhizosphere research, however, including our own, has moved little beyond a dichotomous contrasting of characteristics and processes of the rhizosphere with those in the bulk soil. Whether the variable of interest is microorganism numbers or activity, organic compounds, biological or chemical reactions, or communication-signaling, “rhizosphere effects” are frequently indexed by R/S ratios, i.e., the ratio of an attribute in the rhizosphere with that in bulk soil (Katznelson 1946). For many soils, R/S ratios for microorganism numbers range from 5 to 20 to even much greater than 100 (Anderson et al. 2002, Richter and Markewitz 2001). The responsiveness of roots and soil biota is demonstrated by the very rapid stimulation of microorganism numbers in rhizospheres of young seedlings. Deep in the soil, in B horizons for example, active bacteria and fungi may be prolific in the rhizosphere but close to detection limits throughout the surrounding bulk soil (Table 1).
Approaches to the rhizosphere based on R/S ratios have been highly instructive in emphasizing the biological and chemical activity of the habitat of the near-root environment. Unfortunately, R/S ratios also tend to emphasize a dichotomy and lack of interaction between the rhizosphere and bulk soils. This may be highly significant considering that rhizosphere processes may contribute to a transformation of bulk soils over pedogenic time.
By arguing for a broader perspective of the rhizosphere, we by no means oppose traditional concepts of the rhizosphere and the rhizosphere’s significance to terrestrial ecosystems. We share concepts of the rhizosphere held by most ecological scientists: that rhizospheres represent a poorly defined volume of soil, adjacent to roots, and that steep microbiological and chemical gradients help define the rhizosphere environment. In fact, we recommend broadening our perspective of the rhizosphere precisely because the intensity of biological and chemical activity in the near-root environment can have profound effects on the whole soil when integrated over pedogenic time. In the rhizosphere, the higher order biological-chemical-physical interactions make research on these issues some of the most important to all of soil science, biogeochemistry, and ecosystem ecology.
Rhizospheres as Microsites Within Soil Profiles
Pedologists sometimes subdivide the soil profile into upper and lower systems (Brimhall et al. 1991, Richter et al. 1995). The upper soil system, i.e., the O, A, and at least the upper B horizons, is characterized by intense biological activity and extensive and thorough rooting. Roots and associated microorganisms control much about the physics and chemistry of the upper soil system. Not infrequently, the entire volume of upper-system soil is thoroughly explored by fine roots and their associated microbial communities. In soils that are prolifically rooted by fine roots and associated mycorrhizal fungi, much of the basal layer of the O horizon and all of the A horizon may part of the active rhizosphere.
Frequency of roots and root microbes, and concentrations of organic matter and bioavailable nutrients often diminish with increasing soil depth. In the soil’s lower system, deep within B and throughout C horizons, the near-root environment is nothing less than an oasis of resources compared with the surrounding subsoil. Rhizospheres in the lower soil system appear to have more in common with A horizons than they do with the B and C horizons that surround them (Table 1). In other words, R/S ratios for biologic and chemical properties may increase with increasing soil depth (Figure 1), a pattern that emphasizes the functional significance of rhizospheres in lower soil systems.
Advanced-Weathering Stage Soils
Because soils are open thermodynamic systems, over time soils proceed through a remarkable set of changes, as energy, chemical elements, and water are processed. Three stages of soil mineral weathering were used by Jackson and Sherman (1953) to generalize about the weathering of minerals and the formation of the earth’s soil (Table 2). Mineral weathering is stimulated under warm and humid climates, as minerals are weathered by physical, chemical, and biological interactions. Although new secondary minerals are formed during soil development, the soil’s overall acid-neutralizing capacity is consumed through time. Provided that the soil’s landform is geomorphically stable, e.g., on level or nearly level interfluves, weathering of soils proceeds through a sequence indicated in Table 2, with hydrologic leaching of chemical elements exceeding inputs via mineral decomposition and atmospheric deposition. Over pedogenic time, weathering consumes even large pools of primary minerals and advanced-weathering stage soils will result if hydrologic removals of solutes outpace renewals.
Our interest in this chapter is in linking rhizosphere processes to the advancement of mineral weathering, and specifically to the formation of the earth’s most weathered soils, the Ultisols and Oxisols, common soils found within temperate and tropical climates throughout the world. Ultisols and Oxisols are closely correlated to Acrisols and Ferrosols in the FAO-UNESCO classification; Red-Yellow Podzols and Latosols in older 20th century classifications; Sols Ferralitiques in French classifications; Kaolisols, Ferrisols, and Ferralsols in Belgian systems; Red soils in Chinese systems; and Podzolicos Vermelho Amarelo and Latosolos in Brazilian systems. Since the original starting materials have been completely transformed by weathering, these soils are composed of only the most insoluble chemical elements and recalcitrant minerals. Our point in this chapter is straightforward: that the development of advanced weathering-stage soils is a result of mineral weathering reactions that are greatly affected by rhizosphere processes.
In the humid temperate zones and tropics, geomorphically stable surfaces can develop enormously deep profiles, often 5 to >30-m deep above unweathered bedrock. It is not uncommon that soil weathering exhausts nearly all primary minerals and a number of chemical elements throughout these depths (Figure 2). Not atypical is an upper 1 to 3 m of O, A, and B horizons, below which is the C horizon of highly variable depth, all of which is acidic, extremely low in base cations and P, and depauperate of primary minerals.
Several calculations help emphasize the extreme state of weathering represented by such soils. In our long-term research site at the Calhoun Experimental Forest in the Piedmont of South Carolina, unweathered granite and gneiss underlies A, B, and C horizons in soil profiles that may total up to 25-m of unconsolidated material over actively weathering bedrock. The pH of the unweathered bedrock is 7.9 in water, yet the pH of the soil sampled throughout at least the upper 8 meters of A to C horizons ranges from 3.8 to 4.2 in 0.01 M CaCl2. Exchangeable acidity (in 1 M KCl) totals about 3000 kmolc ha-1 in the upper 6-m of soil profile, an enormous quantity of acidity. Even more impressive however is the quantity of acid that has been consumed during weathering of granitic-gneiss into the kaolinite-dominated, Fe-oxide/hydroxide-rich Ultisol. Transforming 1 m of granitic gneiss into kaolinite is estimated to require a minimum of 100,000 kmolc ha-1 (Richter and Markewitz 1995, 2001). Weathering 10 m of granitic gneiss to kaolinite thus requires about 106 kmolc ha-1. This extreme acidification raises questions about the sources and rates of acid inputs that have so thoroughly weathered this Ultisol, as well as advanced weathering-stage soils in general. In the next section of our chapter, we examine the physical and chemical interactions of rhizospheres on soil mineralogy which over pedogenic time lead to advanced weathering soil.
Rhizospheres Are Where Ecosystem Concentrate Interactions with Soil Minerals
The extreme acidification and weathering state of Ultisols and Oxisols raise questions about the mechanisms by which these soils are so transformed over time. Since rooting affects both physical and chemical weathering in soils and rocks, in this section, we examine some physical effects of rooting on the rhizosphere environment, and subsequently examine prominent sources of rhizosphere acidity that stimulate weathering and eventually form Ultisol and Oxisol soils.
The Physical Attack
Growing roots and their mycorrhizal hyphae follow pores and channels that are generally not less than their own diameters (Figure 3). As roots grow, they expand in volume axially and radially. As roots increase in diameter, they exert enormous pressures on the surrounding soil by cylindrical expansion.
The pressure of growing roots can mechanically decompose minerals, by exerting pressures on individual mineral grains and whole soils, across spatial scales that range from micrometer to many decimeters and even meters (Misra et al. 1986, Dexter 1987, April and Keller 1990, Richter et al. submitted). Such pressures affect soils differently in the upper and lower systems.
In the upper soil system, growing roots can displace soil upward. Surrounding the root collars of large trees, for example, surface soils are uplifted in the surrounding rhizosphere. Individual soil particles can be moved on the order of a meter over the course of several decades (Figure 4). Over time, the uprooting of trees during storms causes particle abrasion and mixing of the upper soil system, increasing the soil’s surface area subject to chemical weathering. Over generations of trees, tree growth and uprooting facilitate physical weathering of minerals in surficial soil layers, no doubt making minerals more susceptible to chemical weathering attacks.
Lower in soil profiles, the pressures of growing tap roots are relieved by soil consolidation, a process that must have severe physical effects on individual soil particles, soil structure, and overall soil architecture. In contrast to the A horizon, root growth pressures can not be relieved by upward displacement in B and C horizons. Growing tap roots consolidate surrounding soils as they establish anchorage by expanding radially. On the Duke Forest, bulk density of B horizon materials adjacent to tap roots of 70-year old trees exceeded 1.9 Mg m-3, a consequence of tap roots consolidating soil for up to 50-cm radial distance from the growing root (Figure 5). These rhizosphere effects must cause severe abrasion and disintegration of individual soil particles; reduced porosity, hydraulic conductivity, and aeration; and greatly altered biogeochemical functioning. Such effects would appear to accumulate over time and may well be a significant and understudied process affecting physical and chemical weathering of forests. In sum, such mechanical processes would have impacted many Ultisols and Oxisols on numerous occasions, given their relatively great age.
Under other rooting conditions, even relatively consolidated rocks are susceptible to physical effects of roots. There are in the literature many examples root wedging, i.e., growing roots expanding rocks’ joints and fractures (their planes of weakness). All these physical effects accelerate the chemical weathering attack by increasing the mineral surface area that can be contacted by organic compounds, electrons, and protons.
The Chemical Attack
Root-microbe systems not only physically attack soil’s minerals, they chemically transform soils as well. We focus on four biological processes that affect soil acidification and weathering, and thereby eventually promote the formation of advanced weathering-stage soils such as Ultisols and Oxisols. The four biological processes that are particularly significant in rhizospheres include:
· root nutrient uptake of anions and cations which affects the production of protons and hydroxyls in the rhizosphere.
· partial oxidation of organic matter produces organic acids, many of which have relatively low pKa and contribute protons to the rhizosphere.
· oxidation of organic matter stimulates redox reactions of electron-deficient metals and consequently soil acid-base status.
· complete oxidation of organic matter and plant-root respiration produces CO2 and thereby carbonic acid.
In sum, these sources of acidity result from the vegetative production and decomposition of organic matter. Although other biogenic acid systems can affect soil acidification and weathering dissolution (most especially oxidation reactions involving nitrogen and sulfur), we focus on these four as widespread sources of protons across a wide range of soils and rhizospheres.
Root uptake of nutrient ions. A major source of acidity is derived from the uptake of nutrients by vegetation. Root uptake of nutrient cations and anions directly affects the soils acid-base chemistry because the physiological process of nutrient uptake is charge balanced: i.e., the uptake of cations and anions affects a release of H+ and OH-, respectively, into the rhizosphere. If vegetation takes up more nutrients as cations than anions, the plant accumulation of nutrients acidifies soil. A variety of scientific literature describes the acid balance of terrestrial ecosystems, including old-growth forests, aggrading secondary forests, and cultivated field crops (Pierre et al. 1970, Ulrich 1980, van Breemen et al. 1982, Driscoll and Likens 1982, Sollins et al. 1980, Binkley and Richter 1987, Johnson and Lindberg 1992, Markewitz et al. 1998).
Plant species exert large effects on soil acidity due in part to differences in plant-nutrient uptake. Deciduous trees, such as oaks and hickories, have calcium uptakes that are two to five-fold greater than conifers such as pines, and thus much more potential to promote acidity throughout full soil profiles. Alban (1982) demonstrated this with comparisons of acidity throughout A and B horizons of pine, two species of spruce, and aspen, and we hypothesize that such differences are affected most intensively in rhizospheres. Richter (1986) estimated H+ budgets of five forest stands and illustrated the wide variation in net cation uptake and potential for soil acidification.
Within the rhizosphere, low pH has been measured with plant systems having large net cation uptake (Lynch 1990). As much as two pH-unit depressions have been measured in rhizospheres compared with bulk soil, conditions that affect not only cation exchange in the rhizosphere, but dissolution of weatherable minerals as well.
Organic acid production. Organic acids play significant and varied roles in soil acidification and mineral weathering, contributing protons directly to acidification, serving as ligands that complex metals, and stimulating redox reactions of electron-deficient metals (Buol et al. 1989, Duchaufour 1982, Brimhall et al. 1991, Qualls and Haines 1991, Boyle and Voigt 1973). A very wide variety of organic compounds have acid functional groups, mainly carboxylic or phenolic, and these originate not only from products of decomposition and carbon oxidation but also can be exudates from plant roots and associated microbes (Lapeyrie et al. 1987, Herbert and Bersch 1995). Organic acids range widely in molecular weight from relatively small compounds such as oxalic and citric acids to much larger humic compounds with combinations of carboxylic and phenolic functional groups.
Organic acids are weak acids with values for pKa that range widely from as low as 3 to as high as 9. Many carboxylic functional groups have a relatively low pKa, with oxalic, citric, malic, and formic acids, all with pKs <4.0. Such acids readily contribute protons to the soil system under a wide range of pH conditions. In addition to being a source of protons, many organic acids are effective ligands that complex metal cations such as Al and Fe, thereby greatly facilitating metal translocation within soils and enhancing mineral weathering.
In general, organic acids are typically highest in concentration in O horizons and decrease sharply with depth into the mineral soil (Herbert and Bertsch 1995, Fox and Comerford 1990, Richter and Markewitz 1995b). For example, in our Calhoun Experimental Forest, collections of soil water within soil profiles that support pine forests have soluble organic acids that decrease from about 115 umolc/L in O horizon waters, to 73 umolc/L in waters of A horizons, and are below detection at 60-cm and deeper (Markewitz et al. 1998).
The commonly observed decrease in organic concentrations with soil depth belies the fact that organic acids significantly affect acidification and weathering in the lower soil system. Moreover it is in the rhizosphere where these interactions are most concentrated. Rhizospheres in lower soil systems develop in pedogenic aggregate macropores, solution channels, and deep pedo-geological planes of weakness (Herbert and Bersch 1995), all of which are environments periodically supplied with organic acids given the presence of active roots and associated biota. Such inputs ensure that organic acids are highly significant to soil systems throughout soil profiles.
Redox cycling of electron-deficient metals. Acid-base consequences of redox reactions have been studied in anaerobic soils subject to high water tables (e.g., Brinkman 1970, van Breemen 1988). In seasonally waterlogged soils such as paddies and other wetlands (van Breeman 1988), redox cycles of reductive dissolution and oxidative precipitation are separated in time: during wet seasons and high water tables, Fe3+ is reduced and acidity is consumed; during dry seasons, Fe2+ is oxidized and acidity produced. Few studies have considered whether these wetland redox processes are related to redox processes in soils that are only infrequently subject to low redox potential.
In fact, the extensive occurrence of mottling in many soils suggests that redox reactions may be significant to acid-base reactions in many upland soils. In humid climates, all but the most well drained soils experience at least temporary periods of saturation during which electron-deficient Fe and Mn oxides and hydroxides can function as electron acceptors in microbially mediated reactions. The close correspondence of rhizospheres and soil mottling in a number of upland soils (Fimmen 2004) gives rise to the hypothesis that rhizosphere-stimulated reductive dissolution initiates a chemistry that results in a significant input of acidity to bulk soil environments.
Soil mottles are the outcome of Fe reduction, translocation, and oxidative precipitation (Ref). Two rhizosphere processes facilitate reductive dissolution of Fe and thereby mottling: the ready supply of organic reductants from root and microbial activities, and the consumption of O2 by respiration in the near-root environment which readily lowers the redox potential of the near-root environment. Ferrous iron is relatively soluble and can be mobilized in the liquid phase until it encounters soluble O2, at which time it is rapidly oxidized and precipitated as Fe3+. This redox cycling of Fe can be represented by:
Reductive Dissolution Oxidative Precipitation
Amorphous Fe(OH)3
Fe(OH)3(s) + ¼ CH2O + 2H+ ( Fe2+ + ¼ CO2 + 2¾H2OFe2+ + ¼O2 + 2½ H2O ( Fe(OH)3(s) + 2H+
Goethite (FeOOH)
FeOOH(s) + ¼ CH2O + 2H+ ( Fe2+ + ¼ CO2 + 1¾ H2OFe2+ + ¼ O2 + 1½ H2O ( FeOOH(s) + 2H+
Thus, reductive dissolution of Fe3+ in the rhizosphere consumes protons but upon translocation to adjacent but more oxidized microsites, oxidative precipitation of ferric Fe produces protons and which facilitate cation exchange and mineral weathering in the bulk soil environment (Figure 6). Reaction kinetics of adsorbed Fe2+ is relatively rapid compared to soluble Fe2+ (Wherli 1990), and we hypothesize that oxidation of adsorbed Fe2+ rapidly yields co-adsorbed hydrogen ions, which protonate pH-dependent exchange sites and stimulate mineral dissolution. Thus, the rhizosphere processes that drive reductive dissolution of Fe oxides and hydroxides can lead directly to oxidative precipitation, acidification, and mineral weathering in the bulk soil environment.
Carbonic acid system. Respiration is a central process of ecosystems, and organic-matter decomposition and plant-root respiration elevate belowground CO2 greatly. Soil’s elevated CO2 can stimulate significant cation exchange and weathering due to interactions of carbonic acid with mineral surfaces (Reuss and Johnson 1986, Amundson and Davidson 1990, Richter and Markewitz 1995b). Carbonic acid weathering involves all three phases of the soil system: CO2 in the gas phase, carbonic acid and associated ions in the liquid phase, and cation exchange and mineral surfaces in the solid phase. Carbonic acid weathers minerals throughout soil profiles, but since partial pressures of CO2 typically increase with soil depth, B and C horizons are subject to the main brunt of this acid system’s attack. Since rhizospheres are the source of nearly all of the subsoil’s CO2, we expect that CO2 is most greatly elevated in subsoil rhizospheres.
Since H2CO3* is a very weak acid with a pKa1 of 6.36 (Stumm 1996), the carbonic acid weathering system is widely conceived to be self-limiting in its effects on acidification and weathering (Reuss and Johnson 1986). However, H2CO3* (the sum of dissolved and hydrated CO2) can be an effective acidifying agent even at relatively low pH, as pure H2CO3 (hydrated CO2) is a much stronger acid than H2CO3* and even has a reported pKa1 of 3.76 at 25 oC (Snoeyink 1980). This apparently little appreciated but critical feature of the chemistry of carbonic acid is critical to soil weathering, especially as CO2 commonly ranges between 1 and 10% in soil atmospheres. The elevated partial pressures of soil CO2 help ensure relatively high concentrations of H2CO3* in solution and that protons of even a small fraction of H2CO3* will dissociate, despite low pH, due to the low pKa1 of pure H2CO3. Equilibrium calculations indicate that between 1 to 10% CO2, in situ pH of dilute soil waters can be depressed to 4.9 to 4.4, respectively (Table 3) and that with dilute solutions, HCO3- will be about 15 and 46 umol L-1 in soil water, very close to what is measured by titration in soil water collections of extremely acid Ultisols (Markewitz et al. 1998). Although the pH of soil solution may rarely be lowered below 4.5 due to CO2 and will not be able to mobilize Al from the soil profile (Reuss and Johnson 1986), elevated subsoil and rhizosphere CO2 ensures that carbonic acid affects mineral dissolution and can create Al-saturated soils.
Two distinct lines of evidence appear to corroborate these conclusions. First, many Ultisols and Oxisols underlain by deep saprolites or C horizons have incredibly acidic C horizons. We consider that the carbonic acid system is the most likely candidate for pushing such systems to such extreme states of weathering. Remarkably, some of these profiles are 10s of meters in depth with low base saturations throughout, perhaps well below the active zone of rooting. Given that most subsoil CO2 originates from the rhizosphere, we are again impressed that rhizosphere processes may affect such enormous volumes of soils.
A second line of evidence comes from recent laboratory studies in which solutions equilibrated with varying pressures of CO2 were used to extensively leach soils that had a wide range of cation exchange capacities and weatherable minerals. Cation exchange was the dominant mechanism supplying cations to solution and greatly diminished base saturation of all soils. All exchangeable base cations were displaced from two soils, and in one, even 1% CO2 displaced all exchangeable base cations from the soil and even elevated Al in soil solution. We should not underestimate the potential for soil CO2 to acidify and weather soil minerals.
Overview of the Rhizosphere as Interface of the Ecosystem’s Weathering Attack
Whether the perspective is one of mechanics or of chemistry, the rhizosphere is an interface between biology and geology that has very broad consequences for earth’s biogeochemistry as well as soil formation.
We started this chapter by noting that the preponderance of scientific literature on the rhizosphere is narrowly focused in space and in time. While the narrow definition of the rhizosphere has helped emphasize that actively growing roots create unique and special environments with great consequence for plants and microbes, rhizosphere environments also have a wide range of highly significant effects on soil formation and biogeochemistry. Because the rhizosphere is the interface where roots exert intense physical pressures and remarkable chemical dissolution reactions on minerals, rhizospheres are fundamentally important to soil formation, including the formation of the earth’s most extremely weathered soils, the Ultisols and Oxisols.
A general hypothesis is presented for how processes in rhizospheres weather soil minerals physically and chemically and with great intensity and consequence. To describe this hypothesis, we subdivide the soil profile into upper and lower soil systems (Brimhall et al. 1991, Richter et al. 1994a). The upper system includes O, A, and B horizons (the horizons of the traditional solum). The lower soil system includes C horizon or saprolite.
Over millennial time scales, acidity due to root uptake of nutrient ions affects the upper soil system most especially, and has an important if secondary role in rhizospheres of the lower system as well. This latter acidity depends on root distribution with soil depth, and on spatial patterns of nutrient uptake. Organic acids contribute significant amounts of acidity to the upper soil system. Such acids are closely associated with organic matter decomposition and are produced in rhizospheres as well, and as a consequence are highest in concentration in O horizons, in upper A horizons, and in the rhizosphere. Organic acids have important effects in the lower soil system, specifically in rhizospheres. Organic acids in subsoils acidify minerals directly, complex metal cations such as Fe and Al, or along with other organic reductants serve as an electron source for electron-deficient metals. Given the variety of organic compound input, rhizospheres are sites of reductive dissolution particularly of Fe and Mn, which are translocated out of the rhizosphere only to precipitate and oxidize on contact with soluble O2. Given that rhizosphere-mottling involves a spatial separation of microsites of reduction and oxidation, protons are produced in oxidative microsites and can affect cation exchange and weathering dissolution.
Lastly, carbonic acid is positively correlated with soil depth and is suggested to be elevated in rhizospheres as well, given that rhizospheres are hot spots of root and microbial respiration. Carbonic acid is influential in both upper and lower soil systems although the concentration gradient of CO2 leads to highest concentrations of CO2 in subsoils and presumably in rhizospheres. Recent evidence suggests that the potential for carbonic acid to weather soils and acidify even already acidic soils has been underestimated and closer examination of CO2 and carbonic acid weathering in rhizosphere’s is very much needed.
In concert, a large number of ecosystem processes are concentrated in the rhizosphere where they attack soil minerals throughout the soil profile. Chemical elements are released by the combined effects of mechanical and chemical weathering, taken up by plants and microbes to meet nutritional requirements, adsorbed to electrostatically charged surfaces, complexed by ligands, recombined into secondary clay minerals, and leached to groundwaters, rivers, wetlands, lakes, and eventually to the ocean. Over pedogenic time, on stable terre firme landforms, the ultimate products of such weathering are advanced weathering-stage soils, such as Ultisols and Oxisols. Only a remarkably few chemical elements, such as Zr and Ti, are insoluble enough to resist transportation from weathering environments, despite the physical and chemical effects of the rhizosphere.
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van Breemen N. (1988) Effects of seasonal redox processes involving iron on the chemistry of periodically reduced soils. In: Iron in Soils and Clay Minerals (eds Stucki J.W., Goodman B.A. & Schwertmann U.), pp. 797-810. D. Reidel Publishing Company, Dordrecht, Holland.
Van Ranst E. & De Coninck F. Evaluation of ferrolysis in soil formation. European Journal of Soil Science 53, 513-519. 2002.
Wehrli B. (1990) Redox reactions of metal ions at mineral surfaces. In: Aquatic Chemical Kinetics: reaction rates of processes in natural waters (ed Stumm W.), pp. 311-336. John Wiley & Sons, Inc., New York, New York.
Table 1. Chemistry and microbial properties of bulk soil (conventional 6-cm dia core samples) in four soil horizons, and in rhizosphere soils (<2-mm distance from roots) sampled at 2 to 3 m depth in the pine-forest soil of the Calhoun Experimental Forest. Soil microbial data courtesy of Dr. Elaine Ingham, Oregon State University, Corvallis.
Soil
material
Soil
depth
Total
carbon
Total bacteria
FDA-active bacteria
Total
fungi
FDA-active
fungi
Oe horizon
m
-
%
-
no. g-1
1.97 x 108
no g-1
32.9 x 106
m g-1
59160
m g-1
906
A horizon
0-0.075
0.70
1.44 x 108
23.8 x 106
18140
653
BE horizon
0.6-1.0
0.24
1.59 x 108
1.47 x 106
294
5.5
B horizon
2.0-3.0
0.073
1.23 x 108
0*
0*
0*
Rhizosphere
soil in B
2.0-3.0
0.42
3.17 x 108
3.54 x 106
1467
65.8
* Detectable concentrations for FDA-active bacteria, total fungi, and FDA-active fungi are < 4 x 103 units g-1, < 0.3 cm g-1, < 0.3 cm g-1, respectively.
Table 2. Soils are open systems and they proceed through developmental stages of weathering and formation as energy, water, and chemical elements are processed over time. Three general weathering stages of soil mineral weathering were used by Jackson and Sherman (1953) to illustrate soil formation. Table 2 illustrates the implications of this dynamism for common soil minerals and soil orders. This paper will illustrate the fundamental importance of rhizosphere processes to formation of advanced weathering stage soils.
Jackson-Sherman (1953) soil weathering stage
Attribute
Early
Intermediate
Advanced
Soil Taxonomy orders (Soil Survey Staff 1998)
Entisol
Andisol
Inceptisol,
Mollisol,
Alfisol
Ultisol
Oxisol
Common soil minerals
Gypsum
Calcite
Olivine
Biotite
Feldspar
Feldspar,
Muscovite,
Vermiculite,
Smectite
Kaolinite
Gibbsite
Fe oxide/hydroxides
Table 3. Solution pH of low ionic strength solutions in equilibrium with CO2 at different partial pressures. The soil atmosphere at >1-m depths of many soils ranges up to 5 to 10% CO2, and in atmospheres of rhizospheres may exceed 10%.
CO2
pH
HCO3 -
%
0.036
5.65
mmol L-1
0.0029
1.0
4.9
0.0145
5.0
4.6
0.036
10
4.4
0.046
100
3.9
0.145
1
10
100
1,000
10,000
0
0-0.1
0.1-0.4
0.4-0.8
0.8-1.5
1.5-2
2-2.5
Soil Depth (m)
Fungal Biomass (ug/g)
Entire Soil
Rhizosphere
Figure 1. Fungal biomass in bulk and rhizosphere soil at an Appling soil at the Calhoun Experimental Forest, South Carolina. Soil supported a 47-year-old loblolly pine (Pinus teada) forest.
Figure 2. Calcium loss from three deep soil profiles (Tarrus and Cecil are Ultisols, Enon is an Alfisol). Tau expresses the estimate of the original Ca that has been lost during soil formation (e.g., -0.5 indicates that 50% of the Ca in the primary minerals has been lost to weathering).
Figure 3. ESEM images of rhizospheres (1.5 m depth) in Appling soil B horizon.
0
5
10
15
20
0
12
24
36
48
60
72
84
96
Radial distance from tree
(cm)
Vertical uplift (cm)
41cmTree
62cmTree
63cmTree
74cmTree
Figure 4. Soil microtopography surrounding four 70-year-old loblolly pine (Pinus teada) trees in the Duke Forest, North Carolina. Diameters are given adjacent to each tree.
Figure 5. Bulk density of soil surrounding two 70-year old loblolly pine trees. Bulk density in g/cm3.
1.2 m ------
1.5 m ------
Figure 6. Pronounced soil mottling in B horizons at Calhoun Experimental Forest (1.2 to 1.5m depths).
Rhizospheres (5Y Munsell)
Many fine-roots & Basidiomycete hyphae
Carbon rich (0.41 %)
Iron poor (Fecyrst = 5.2 mg/g)
Clay rich (73% c)
Bulk soil (2.5 YR Munsell)
Absence of fine roots & fungi
Carbon poor (0.086 %)
Iron rich (Fecryst = 37 mg/g)
Clay poor (26% c)
1.4m
PAGE
27
Chart1
00
0-0.10-0.1
0.1-0.40.1-0.4
0.4-0.80.4-0.8
0.8-1.50.8-1.5
1.5-21.5-2
2-2.52-2.5
Sheet1
Soil Depth (m)Entire SoilRhizosphereSoil Depth (m)R/ES
02,9628,19602.77
0-0.15591,3640-0.12.44
0.1-0.42095700.1-0.42.73
0.4-0.837.682.40.4-0.82.19
0.8-1.520.746.60.8-1.52.25
1.5-29.5740.91.5-24.27
2-2.53.8270.22-2.518.35
Sheet1
00
00
00
00
00
00
00
Sheet2
Sheet3
Chart1
0000
4444
8888
12121212
16161616
20202020
24242424
28282828
32323232
36363636
40404040
44444444
48484848
52525252
56565656
60606060
64646464
68686868
72727272
76767676
80808080
84848484
88888888
92929292
96969696
100100100100
104104104104
RTCLRVOL surface profiles
41cmTree62cmTree63cmTree74cmTree
08.815.41919.6
47.614.118.517.7
85.613.414.217.1
123.811.513.815.8
163.29.213.515.3
202.07.711.213.5
241.35.98.812.1
281.24.47.89.9
321.12.96.58.1
360.42.05.57.1
40-0.11.55.26.4
440.21.04.26.1
48-0.00.33.55.3
520.5-0.02.85.0
561.84.5
600.84.0
642.33.9
682.33.9
72
76
800.72.1
841.3
88
921.0
96
1000.5
1040.2
108
RTCLRVOL surface profiles
RTCLRVOL
RTCLRVOL.WK1
TREE1DBH44.45TREE2DBH48.26TREE333.02TREE452.58
CM UNITSDBAH60.96CM UNITSDBAH62.74CM TO IN UNITS41.40CM TO IN UNITS73.91
TREE1 ROOT COLLAR DIAMETER = 60.96 CM; DBH = 44.45 CM.TREE2 ROOT COLLAR DIAMETER = 62.74 CM; DBH = 48.26 CM.TREE3 ROOT COLLAR DIAMETER = 41.4 CM; DBH = 33.0 CM.TREE4 ROOT COLLAR DIAMETER = 73.9 CM; DBH = 52.6 CM.
Tree1aTree1bTree1cTree1dTree1Tree2aTree2bTree2cTree2Tree3aTree3bTree3cTree3dTree3POSITNTree4aTree4bTree4cTree4dTree4
DfrTcmdHtcmdHtcmdHtcmdHtcmMndHtMean1a1bMean1c1dMean1a1bMean1c1dDfrTcmdHtdHtdHtMndHtMean2a2bMean2cMean2a2bMean2cDfrTcmdHtcmdHtcmdHtcmdHtcmMndHtdHtcmdHtcmdHtcmdHtcmMndHt
014.014.015.714.615.414.015.716.014.331414.018.514.018.5015.715.015.419.60
414.516.516.815.914.114.516.615.513.4715182218.314.211.022.021.510.51.319.120.319.78.87.616.017.018.318.017.317.77.6
815.714.518.318.016.613.415.118.214.911.81116192118.713.813.521.019.011.55.320.319.323.920.120.97.611.615.717.519.618.817.917.111.6
1217.317.019.320.318.511.517.119.812.910.21516202119.013.515.521.017.011.59.324.422.925.119.322.95.615.617.818.021.119.819.215.815.6
1620.320.320.621.820.89.220.321.29.78.81920232121.311.219.521.013.011.513.325.425.128.419.824.73.819.618.819.121.619.319.715.319.6
2021.623.921.322.422.37.722.721.87.38.22323272123.78.823.021.09.511.517.325.425.430.220.325.33.223.620.120.325.120.321.513.523.6
2424.125.423.623.424.15.924.823.55.26.52724292124.77.825.521.07.011.521.326.726.431.521.626.52.027.621.120.826.723.122.912.127.6
2925.426.225.924.925.64.425.825.44.24.63226312126.06.529.021.03.511.525.327.225.933.022.927.21.331.621.823.629.225.925.19.931.6
3325.428.428.226.227.12.926.927.23.12.83626332227.05.531.022.01.510.530.327.425.733.023.127.31.236.623.125.730.528.226.98.136.6
3725.429.730.226.728.02.027.628.42.41.64026332327.35.233.023.0-0.59.534.327.426.432.523.427.41.140.623.626.932.029.227.97.140.6
4224.929.533.026.728.51.527.229.82.80.24527342428.34.236.024.0-3.58.538.328.427.433.023.628.10.444.624.925.933.830.028.66.444.6
4724.630.233.327.929.01.027.430.62.6-0.65029332529.03.539.525.0-7.07.543.329.026.733.025.728.6-0.149.624.925.934.330.528.96.149.6
5224.130.535.628.429.70.327.332.02.7-2.05530332629.72.842.526.0-10.06.548.328.726.432.825.428.30.254.625.127.234.332.029.75.354.6
5624.930.735.828.730.0-0.027.832.32.2-2.35931332830.71.845.028.0-12.54.553.329.026.733.025.428.5-0.059.625.428.234.532.030.05.059.6
6125.131.235.829.230.46433332931.70.848.529.057.328.427.933.528.229.563.625.429.234.333.030.54.563.6
6725.431.035.130.230.47035333032.762.329.228.434.530.230.668.626.929.034.533.531.04.068.6
7626.932.035.131.031.27938333234.368.329.230.534.830.231.228.234.031.13.9
8427.933.535.132.032.18738343435.377.327.932.033.529.730.883.630.530.236.334.532.92.183.6
9128.733.035.833.332.79438363436.085.329.032.534.328.431.191.630.531.536.636.133.71.391.6
9729.533.537.134.033.510037363536.092.327.734.034.326.230.598.630.033.036.636.334.01.098.6
10331.237.338.634.835.510638373737.398.327.734.533.824.630.2104.630.534.537.335.634.50.5104.6
11134.036.839.934.336.311438373737.3104.327.735.333.824.130.2110.630.735.337.635.634.80.2110.6
12233.037.640.133.336.012539393738.3112.327.233.833.826.230.2118.630.035.839.637.835.8
12933.037.841.133.036.313240393738.7123.327.232.833.327.930.3129.628.236.342.238.636.3
13432.337.641.932.536.113740403739.0130.327.431.834.529.530.8136.628.236.843.438.636.8
14231.536.644.532.036.114539413939.7135.326.931.834.830.230.9141.628.436.843.738.136.8
14931.535.645.731.836.115239424040.3143.327.231.034.529.230.5149.629.036.844.738.137.1
15631.836.847.831.537.015940414040.3150.327.429.734.328.430.0156.629.736.345.737.837.4
16428.436.648.833.336.816740403939.7157.325.929.234.828.229.5163.631.035.846.738.938.1
17327.437.149.834.337.117640393839.0165.325.929.235.627.929.7171.632.336.648.039.139.0
18436.348.342.318741383738.729.233.831.5180.6
19027.247.533.035.9193424242.032.829.531.1191.635.334.848.037.839.0
20826.937.347.033.036.121142433740.731.828.430.1
21726.937.845.233.835.9443740.532.028.230.1
32.832.8
Tree1Tree2Tree3Tree4Tree4
015.40191.38.8019.619.6
414.1318.55.37.6417.718.6
813.4714.29.35.67.617.117.7
1211.51113.813.33.811.615.817.1
169.21513.517.33.215.615.315.8
207.71911.221.32.019.613.515.3
245.9238.825.31.323.612.113.5
294.4277.830.31.227.69.912.1
332.9326.534.31.131.68.19.9
372.0365.538.30.436.67.18.1
421.5405.243.3-0.140.66.47.1
471.0454.248.30.244.66.16.4
520.3503.553.3-0.049.65.36.1
56-0.0552.857.30.554.65.05.3
60591.859.64.55.0
64640.863.64.04.5
68702.368.63.94.0
723.9
762.1
80790.782.02.11.3
8483.61.31.0
880.5
9291.61.00.2
96
10098.60.5
104104.60.2
110.6
Radii@1cm47623598
DBsH cm6162.741.473.9
StumpVolcc45006571211184684069
TOTRtInfZoneVOLcc29058650646785772992367
VolRtCollar cc12279022467336963454149
Log5.095.354.575.66
MassSoil (g)10437219097231418445104
@0.85g/cc
OriginalSolVolcc7455113640922442317932
@1.4g/cc
60.7142857143
Current
6161
62.762.7
41.441.4
73.973.9