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
2
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.
3
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.
4
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 weatheringstage 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
5
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.
6
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
7
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
8
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 plantnutrient 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 electrondeficient 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
9
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
10
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 + 2¾H2O
Fe2+ + ¼O2 + 2½ H2O  Fe(OH)3(s) + 2H+
Goethite (FeOOH)
FeOOH(s) + ¼ CH2O + 2H  Fe + ¼ CO2 + 1¾ H2O
+
2+
Fe2+ + 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
11
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
12
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,
13
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
14
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.
15
References
April, R.
Anderson, T.A., D.P. Shupack, and H. Awata. 2002. Biotic and abiotic interactions in the
rhizosphere: organic pollutants. p. 439-455. IN P.M. Huang, J.-M. Bollag, and N. Senesi
(eds.) Interactions between Soil Particles and Microorganisms. J. Wiley & Sons,
Chichester, UK.
Buol, S.W., F.D. Hole, R.J. McCracken, and R.J. Southard. 1997. Soil Genesis and Classification,
4th Edition. Iowa State Univ. Press, Ames, IA.
Chadwick, O.A., G.H. Brimhall, and D.M. Hendricks. 1990. From a black to a gray box – a mass
balance interpretation of pedogenesis. Geomorphology 3: 369-390.
Coyne, M. 1999. Soil Microbiology. Delmar Publishers, Albany, NY.
Curl, E.A. and B. Truelove. 1986. The Rhizosphere. Springer-Verlag, Berlin.
Denison, R.F. 2001. Ecologists and molecular biologists find common ground in the rhizosphere.
Trends in Ecology and Evolution 16: 535-536.
Dexter, A.R. 1987. Compression of soil around roots. Plant and Soil 97: 401-406.
Hiltner, L. 1904. Über neuere Erfahrungen und Probleme auf dem Gebiet der Bodenbakteriologie
und unter besonderer Berücksichtigung der Gründüngung und Brache. Arbeiten der
Deutschen Landwirtschafts Gesellschaft 98:59-78.
Jackson, M.L. and Sherman. 1953. Advances in Agronomy
Katznelson, H. 1946. The rhizosphere effect of mangels on certain groups of micro-organisms.
Soil Science 62:343.
Lovejoy, A.O. 1936. The Great Chain of Being. Harvard University Press, Cambridge, MA.
Misra, R.K., A. R. Dexter, and A. M. Alston. 1987. Maximum axial and radial growth pressures
of plant roots. Plant and Soil 95: 315-326.
Priha, O., T. Hallantie, and A. Smolander. 1999. Comparing microbial biomass, denitrification
enzyme activity, and numbers of nitrifiers in the rhizospheres of Pinus sylvestris, Picea
abies, and Betula pendula seedlings by microscale methods. Biology Fertility Soils 30:1419.
Richter, D.D., B. Browne, K.H. Dai, T. Huekel, D. Markewitz, H. Stevens, and A. Stuanes. Soil
compaction by the tap root of a growing tree. Soil Science Society of America Journal
(submitted).
Richter, D.D. and D. Markewitz. 1995. How deep is soil? BioScience 45:600-609.
Rovira, and C. Davey. 1974. Biology of the rhizosphere. Pp. 153-204. IN E.W. Carson (ed.)
The Plant Root and Its Environment. University Press of Virginia, Charlottesville, VA.
16
Wheatley, R., K. Ritz, and B. Griffiths. 1990. Microbial biomass and mineral N transformations
in soil planted with barley, ryegrass, pea, or turnip. Plant Soil 127: 157-167.
Richter D.D. & Markewitz D. (2001) Understanding Soil Change; Soil Sustainability over
Millenia, Centuries, and Decades. Cambridge, United Kingdom.
Stone A.T. (1986) Adsorption of organic reductants and subsequent electron trnser on metal
oxide surfaces. In: Geochemical Processes at Mineral Surfaces (eds Davis J.A. & Hayes K.F.),
pp. 446-461. American Chemical Society, Chicago, USA.
Stumm W. & Morgan J.J. (1981) Aquatic Chemistry: an introduction emphasizing chemical
equilibria in natural waters, 2nd edn. John Wiley & Sons, Toronto, Canada.
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.
17
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
FDAactive
bacteria
Total
fungi
FDAactive
fungi
m
%
no. g-1
no g-1
m g-1
m g-1
Oe horizon
-
-
1.97 x 108
32.9 x 106
59160
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 FDAactive fungi are < 4 x 103 units g-1, < 0.3 cm g-1, < 0.3 cm g-1, respectively.
18
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
19
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 mmol L-1
%
0.036
5.65
0.0029
1.0
4.9
0.0145
5.0
4.6
0.036
10
4.4
0.046
100
3.9
0.145
20
Fungal Biomass (ug/g)
1
10
100
1,000 10,000
Soil Depth (m)
0
0-0.1
0.1-0.4
Entire Soil
Rhizosphere
0.4-0.8
0.8-1.5
1.5-2
2-2.5
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.
21
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).
22
Figure 3. ESEM images of rhizospheres (1.5 m depth) in Appling soil B horizon.
23
Vertical uplift (cm)
20
15
41cmTree
62cmTree
10
63cmTree
74cmTree
5
0
0 12 24 36 48 60 72 84 96
Radial distance from tree
(cm)
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.
24
Figure 5. Bulk density of soil surrounding two 70-year old loblolly pine
trees. Bulk density in g/cm3.
25
26
27
1.2 m ------
Rhizospheres (5Y
Munsell)
 Many fine-roots & Basidiomycete hypha
 Carbon rich (0.41 %)
 Iron poor (Fecyrst = 5.2 mg/g)
 Clay rich (73% c)
Bulk soil (2.5 YR Munsell)
1.5 m ------




Absence of fine roots & fungi
Carbon poor (0.086 %)
Iron rich (Fecryst = 37 mg/g)
Clay poor (26% c)
Figure 6. Pronounced soil mottling in B horizons at Calhoun Experimental Forest (1.2 to 1.5m depths).