69
Soil Development in Volcanic Ash
Fiorenzo C. UGOLINI1 and Randy A. DAHLGREN2
1
Dipartimento di Scienza del Suolo e Nutrizione della Pianta
Università Degli Studi, Piazzale delle Cascine, 15, 50144, Firenze, Italy
2
Department of Land, Air and Water Resources, University of California
Davis, CA 95616, USA
E-mail: [email protected]
Abstract
Volcanoes are revered and feared for their awesome and devastating eruptions that often obliterate
terrestrial ecosystems and bury landscapes with volcanic ash. Yet from these ashes of devastation arise
some of the most productive soils in the world with the capacity to sustain high human population
densities. This paper presents an overview of soil developmental processes occurring in soils derived
from volcanic materials. We examine the genesis of six volcanic soils placed in a developmental
sequence that involves climate and time. Climate, in conjunction with its influence on vegetation, and
time of exposure to weathering are the two primary factors regulating soil development pathways in
volcanic materials. Soils formed in volcanic ejecta have many distinctive morphological, physical and
chemical properties that are rarely found in soils derived from other parent materials. These distinctive
properties are largely due to the formation of noncrystalline materials (i.e., allophane, imogolite,
ferrihydrite, Al/Fe-humus complexes) and the accumulation of organic carbon, the two dominant
pedogenic processes occurring in volcanic soils. Formation of noncrystalline materials is directly
related to the properties of volcanic ejecta as a parent material, namely the rapid weathering of the
glassy particles. Andisols generally form rapidly in humid climates and alter to other soil orders as
soil age and the degree of weathering increase. In regions with intermittent deposition of volcanic ash,
each addition of new material rejuvenates soil developmental processes so that Andisols may be
maintained as a relatively stable soil condition.
Key words: allophane, Andisols, Andosols, soil genesis, soil solution, volcanic ash
1. Introduction
It was probably first said by Aristotle that to know
about the nature of objects you have to know their
origin. On the other hand, humans have made large
use of natural objects, such as soils, without really
knowing their origin. Soil has been used extensively
since human society changed from nomadism to farming, which occurred about 9,000 – 10,000 years ago in
the Middle East. Yet, in agriculture a very strong
pragmatism has pervaded, soil is seen as an object to
exploit and to make more productive. Therefore, it
does not really matter how it is functioning and how it
is originating as long as it is productive! Of course,
progress has been made in understanding soil forming
processes, but often more for the purpose of developing criteria for soil classification than for the shear
curiosity of wanting to know the mechanism of
formation.
Volcanic soils have somewhat escaped this trend,
thank to a small group of dedicated scientists that have
addressed the questions: Why do volcanic soils have
rather unique morphological, physical and chemical
properties? Why are certain secondary minerals in
these soils so prevalent? Why does organic matter
tend to persist? Why is their formation relatively
rapid? Why are young volcanic soils among the
most productive soils in the world? Answers to these
questions often lead to the influence of the parent
material on which these soils have formed.
Before Dokuchaev recognized that soil was the
result of the climate, biota, parent material, topography and time, the geological concept was pervading in
the study of soils. Soils were identified and classified primarily by the origin and nature of the parent
material: glacial soils, fluvial soils, residual soils, etc.
As a well-known English geologist said (Joffe, 1936) there must be as many soils as there are geological
formations! Yet even today it is common to find
articles and books with titles including volcanic soils,
volcanic ash soils, or soils derived from tephra. On
the other hand, there is a paucity of reports on soils
formed on other lithologies. This evokes the question: Why is there such an emphasis on volcanic
70
F. C. UGOLINI and R. A. DAHLGREN
deposits as a parent material, in spite of the fact that
volcanic soils occur across a wide range of climates,
vegetation and topographic positions? Volcanoes are
distributed from the cool, humid climate of Alaska to
the hot, humid tropical areas of Sumatra. While soils
formed from volcanic ejecta in these two regions are
different, they share several unique properties, especially during the early and mid stages of development.
Nevertheless, the soils developed in the humid temperate climate (e.g., Japan, New Zealand) are those that
display the most distinctive characteristics for which
volcanic soils are known (Shoji et al., 1993).
2. Dominant Pedogenic Processes and
Distinctive Properties of Volcanic Soils
Formation of noncrystalline materials (active Al
and Fe compounds) and accumulation of organic
matter are the dominant pedogenic processes occurring in most soils formed in volcanic materials (Shoji
et al., 1993). This combination of processes, occurring preferentially in soils formed in volcanic materials,
is termed “andosolization” (Duchaufour, 1977). In
contrast to podzolization, soil solution studies in
Andisols indicate that there is no significant translocation of Al, Fe and dissolved organic matter (Ugolini et
al., 1988). Formation of noncrystalline materials is
directly related to the properties of volcanic ejecta as a
parent material. The small particle size, the glassy
nature of the particles, and the high porosity and permeability of volcanic ejecta enhance chemical weathering rates (Lowe, 1986; Dahlgren et al., 1999).
Rapid weathering releases elements, such as Si, Al
and Fe faster than crystalline minerals can form, resulting in soil solutions becoming over-saturated with
respect to metastable, noncrystalline materials, such as
allophane, imogolite, opaline silica, and ferrihydrite.
Preferential precipitation of noncrystalline materials
occurs because nucleation of a more soluble phase
(noncrystalline materials) is kinetically favored over
that of a less soluble phase (crystalline minerals)
owing to the lower solid-solution interfacial tension of
the more soluble phase (Stumm, 1992). Over time,
crystalline minerals form at the expense of their metastable, noncrystalline precursors. Climatic conditions also play an important role in the formation of
crystalline minerals as crystallization is promoted as
the soil climate becomes warmer and dryer.
Andisols typically have their colloidal fraction
dominated by Al-humus complexes or allophane/imogolite in humid weathering environments. In contrast, halloysite often dominates in climates with a
distinct dry season or in buried soil layers with
imperfect drainage. Under humid weathering conditions, the composition of the colloidal fraction forms a
continuum between pure Al-humus complexes and
pure allophane/imogolite, depending on the pH and
organic matter characteristics of the weathering environment (Mizota and van Reeuwijk, 1989). Andisols
are often divided into two groups based on their mineralogical composition: allophanic Andisols dominated by allophane and imogolite, and nonallophanic
Andisols dominated by Al-humus complexes and 2:1
layer silicates (Shoji and Ono, 1978; Shoji, 1985;
Shoji et al., 1985). Al-humus complexes form preferentially in pedogenic environments that are rich in
organic matter and have pH values of 5 or less (Shoji
and Fujiwara, 1984). In this pH range, organic acids
are the dominant proton donor, lowering pH and aqueous Al activities through formation of Al-humus complexes. Under these conditions, humus effectively
competes for dissolved Al, leaving little Al available
for co-precipitation with silica to form aluminosilicate
materials. Allophane and imogolite form preferentially in weathering environments with pH values in
the range of 5 to 7 and a low content of complexing
organic compounds (Ugolini and Dahlgren, 1991).
Soil solution studies in Andisols show that allophane
and imogolite form in situ, primarily as a result of
weathering by carbonic acid (Ugolini et al., 1988).
Ferrihydrite is typically found as the dominant iron
oxyhydroxide in Andisols. Because Fe has a greater
stability in oxyhydroxides as compared to humus
complexes, concentrations of Fe-humus complexes
are typically low (Wada and Higashi, 1976).
Accumulation of organic matter is another characteristic property of Andisols. Organic matter is protected against biodegradation by Al toxicity to microorganisms (Tokashiki and Wada, 1975), sorption of
degradation enzymes and organic matter substrate
(Wada, 1977; Tate and Theng, 1980), physical protection within abundant microaggregates (Gregorich and
Janzen, 2000), and/or phosphorus deficiency of microorganisms caused by high P retention (Brahim, 1987).
Organic matter stabilization may occur through formation of Al-humus complexes and sorption to allophane,
imogolite, and ferrihydrite. Burial of soils by intermittent additions of volcanic ash is also a very important factor contributing to subsurface accumulation of
organic matter in volcanic soils. Fire, with the production of charred fragments that eventually are
degraded into black humic acids, has been implicated
for the formation of the typical black A horizon in
Andisols (Shindo et al., 2002). These processes of
organic matter stabilization plays a major role in the
formation of melanic and fulvic surface horizons (Soil
Survey Staff, 1999).
Noncrystalline materials and humus contribute to
the unique chemical and physical properties of
Andisols, such as variable charge, high water-holding
capacity, high phosphate retention, low bulk density,
high friability, and formation of stable soil aggregates
(Shoji et al., 1993). They also influence the productivity of Andisols through their positive role in
retaining and supplying nutrient elements, retaining
plant-available water, and development of a favorable
rooting zone, as well as their potentially negative attributes of P fixation, low exchangeable base concentra-
Soil Development in Volcanic Ash
tions, strong acidity and Al toxicity that develop with
increased weathering intensity.
3. Soil developmental sequence
The primary objectives of this paper are to present
an overview of soil developmental processes in soils
derived from volcanic materials and to briefly discuss
the environmental implication of these processes.
Research on volcanic soils would fill volumes and
several compilations on volcanic soils have been
previously published (e.g., Ugolini and Zasoski, 1979;
Theng, 1980; Yoshinaga, 1983; Wada, 1985; 1986;
Mizota and van Reeuwijk, 1989; Shoji et al., 1993;
Kimble et al., 2000). To distinguish this work from
past reviews, we have taken the approach of examining the genesis of six volcanic soils placed in a
developmental sequence that involves time, but most
of all climate. Time and climate combine to determine the relative degree of weathering and pedogenic
development. Selected soil and site characterization
data for the developmental sequence are shown in
Table 1 and representative soils are shown in Figure 1.
For depicting the genesis of these soils we make use
of both soil solution (when available) and solid-phase
characterization data. Soil solution has the advantage of providing the contemporaneity of the processes, while the solid phase integrates the sum total
of all processes that have occurred since the inception
of soil formation.
71
3.1 Initial stages of soil formation
The Mt. St. Helens pyroclastic flow was deposited
in May 1980 and is used to demonstrate the initial
stages of soil formation in volcanic materials (Table 1).
Soil solutions collected between 1984 and 1986 were
dominated by the cations, Ca2+ and Na+, and the
anions, SO42–, Cl– and NO3– (Fig. 2). To properly
interpret the soil solution chemistry and therefore the
initial processes in these developing soils, it is essential to take into account the events prior to, during,
and after deposition of the volcanoclastic material.
Prior to deposition, volcanic gases emanating from the
vent reacted with the airborne tephra. Vapors of sulfuric, hydrochloric, and nitric acids condensed on
tephra particles during atmospheric transport (Fruchter
et al., 1980). These acids induced intense chemical
weathering of the cation-rich volcanic glass resulting
in the formation of Ca and Na salts with the strong
acids anions. These soluble salts (Na+ and Ca2+ with
SO42–, Cl– and NO3–) began to leach as soon as the
fresh deposits were exposed to precipitation events
(Hinkley and Smith, 1987; Dahlgren and Ugolini,
1989a). While Na+ and Cl– salts were rapidly leached, the leaching of Ca2+ and SO42– salts persisted for 4
years in the humid environment (Fig. 2; Dahlgren et
al., 1999).
Table 1 Selected soil and site characterization data for a developmental sequence of soils forming on volcanic materials.
Soil name
Location
Soil
classification
Parent material
MAP
mm
MAT Elevation
m
°C
Vegetation
Literature
references
St. Helens
Mt. St. Helens
pyroclastic flow,
Washington, USA
Entisol
Dacitic pyroclasitic
flow (4-6 y)
2250
5.6
1063
Barren except for
lupine patches
Nuhn, 1987
Ugolini et al.
1991
Yunodai
Southern Hakkoda
Mountains, Aomori
Prefecture, Japan
Melanudand
Dacitic ashes; Towadaa ash (1000 B.P.)
Chuseri ash (5000 B.P.)
1200
8
410
Miscanthus sinensis;
Fagus crenata;
Weigela hortensis;
Sasa kurilensis
Shoji et al. 1988b;
Ugolini et al.
1988
Tsutanuma
Southern Hakkoda
Mountains, Aomori
Prefecture, Japan
Fulvudand
Dacitic ashes; Towada-a
ash (1000 B.P.)
Chuseri ash (5000 B.P.)
1200
8
430
Fagus crenata;
Sasa kurilensis
Takahashi and
Shoji, 1988; Shoji
et al. 1988b, 1993
Findley
Lake
Western Cascades,
Washington, USA
Andic
Humicryod
Dacitic tephra; Mt. St.
Helens tephra (3500
and 450 B.P.); Mt.
Mazama tephra
(7000-6700 B.P.)
overlying andesitic
glacial till
2300
5.5
1150
Abies amabilis;
Tsuga mertensiana
Ugolini et al.
1977; Ugolini and
Dahlgren, 1987;
Dahlgren and
Ugolini,
1989a,b,c; 1991
Santa Maria Island of Santa
Maria, New
Herbides, Vanuatu
Hydrudands
Basaltic cinders (2000
to 4000 B.P.)
>2000
>23
400-650 Calophyllum;
Hernandia;
Kermadecia
Quantin, 1992
Erromango
Dystrudepts
tending
toward
Oxisols
Basaltic, Holocene
cinders overlying
Plio-Pleistocene
basaltic flows
2500
to
3500
22
to
26
200-400 Agathis obtusa
Calophyllum
Quantin, 1992
Island of
Erromango, New
Herbides, Vanuatu
Fig. 1
Yunodai soil profile; Melanudand
Courtesy of S. Shoji
Mt. St. Helens pyrochlastic flow following
1980 eruption. Note microtopography.
Selected soil profiles representing soil development under
contrasting climate and time of development.
Tsutanuma soil profile; Fluvudand
Courtesy of S. Shoji
Soil profile on Mt. St. Helens
pyroclastic flow.
Hilo soil profile; Hydrudand, Hawaii
Courtesy of H. Ikawa
Findley Lake Spodosol.
Andic Humicryod
72
F. C. UGOLINI and R. A. DAHLGREN
Soil Development in Volcanic Ash
Fig. 2
73
Charge balance, dissolved aluminum and silicon concentrations and pH for precipitation (PPT), canopy
throughfall TF and soil solutions of selected soils comprising the soil developmental sequence formed
on volcanic materials. The anion deficit was assumed to be the contribution of organic anions.
Following deposition, the pyroclastic flow surface
was subjected to additional acidification from acidic
precipitation (pH 3.6 - 5.2). The precipitation acidity
originated from gaseous emissions during the
dome-building phase in the crater following the
eruption.
The acidic deposition was largely
neutralized by weathering reactions within the upper
5 cm of soil (Fig. 2) (Nuhn, 1987). Due to rapid
weathering that consumed H+ and the absence of
organic acids to chelate metals, transport of soluble Al
and Fe was generally less than detection. Soil
solutions were dominated by Si along with the major
cations (Ca2+, Mg2+, Na+, K+) leaching with strong
acid anions (SO42–, Cl–, and NO3–) released by
weathering and solubilization of salts lodged in the
vesicles of glass shards. The proton donors that
actuate weathering by providing protons and
conjugate bases for transport are primarily derived
from the volcanic vent. Lack of biota (plants and
microbes) results in little CO2 production rendering
Therefore, soil
H2CO3 weathering insignificant.
genesis within the plume of gaseous emanations is
conditioned by volcanic activity, prior to, during, and
after the emplacement of the deposits.
This geochemical scenario is somewhat modified
by the microtopography of the pyroclastic flow, which
displayed an undulating surface with approximately
30 cm of relief between mounds and depressions.
Weathering was more intense in depressions because
they collected more water than adjacent mounds.
Detrital 2:1 layer silicate minerals present in the
original deposits (smectite-saponite and trioctahedral
vermiculite) (Pevear et al., 1982; LaManna and
Ugolini, 1987) were degraded within five years to a
poorly crystalline kaolin (Nuhn, 1987; Ugolini et al.,
1991). A small accumulation of noncrystalline materials, as revealed by oxalate extractable Al (∼2 – 3 g
kg–1), was also detected in the upper 14 cm of soils in
the depressions. In contrast, the detrital 2:1 layer
silicates were present and the kaolin mineral absent in
the mound portion of the landscape.
Nitrogen was the major nutrient limiting establishment of plants in the vicinity of Mt. St. Helens
following the eruption. Plants were able to rapidly
regenerate in locations where their roots could tap
nitrogen pools in the buried soil or where erosion
exposed the buried soil. In contrast, thick pyroclastic
flow surfaces were colonized by nitrogen fixing lupine
(Lupinus spp.). The presence of lupine resulted in
dramatic, short-term changes in soil genesis. It is
commonly believed that nitrogen fixing plants tend to
acidify soils (e.g., Ugolini, 1968; van Miegroet and
Cole, 1985); however, soils colonized by lupine in the
vicinity of Mt. St. Helens had a higher pH than
adjacent barren sites (Table 2). The increase in pH
was due to neutralization of acidic deposition by the
74
F. C. UGOLINI and R. A. DAHLGREN
Table 2 Selected soil characterization data for a developmental sequence of soils forming on volcanic materials.
pH
Organic C
g kg–1
Alp
g kg–1
Alo
g kg–1
Sio
g kg–1
Feo/Fed
Allophane &
Imogolite
g kg–1
4.5 → 5.7
5.6 → 7.2
0.9
17.1
<0.2
<0.2
0.3 → 2.2
0.3 → 1.2
<0.1
0.1 → 0.7
N.D.†
N.D.
Not detectable
Not detectable
Yunodai
A1
2Bw
4.9
5.7
107
36
11
6
14
33
2
15
0.89
0.75
14
107
Tsutanuma
A1
2Bw
4.6
5.9
124
26
10
4
12
25
2
11
0.68
0.88
14
78
Findley Lake
E
Bs1
3.8
4.9
20
45
1
11
1
49
1
12
0.15
0.72
Not detectable
85
Santa Marie
A1
B
~6
5.5 → 6.0
140 → 200
11 → 23
N.D.
N.D.
3 →10
3 →10
N.D.
N.D.
N.D.
N.D.
150 → 300
150 → 300
47 → 99
4 → 10
N.D.
N.D.
N.D.
N.D.
N.D.
N.D.
N.D.
N.D.
<10
<5
Location/
Horizon
Mt. St. Helens
Barren
Lupine
Erromango
A1
4.0 → 5.0
B
4.4 → 5.4
†
N.D. = not determined
lupine foliage. As acidic deposition interacted with
the foliage, H+ was exchanged for base cations from
the leaves. This process does not prevent soil acidification altogether, but changes the distribution of
protons added to the soil. In the barren areas, buffering of the proton load occurs nearly exclusively in the
surface layer (0-5 cm), while in the lupine patches the
buffering occurs throughout the rooting zone. Uptake of base cations from the rooting zone to replenish
the foliage results in release of protons by the roots.
Consequently, plant invasion on soils developed near
an acidic gas-producing volcano alters the acidification pattern of the developing soil. A similar situation was reported in an area affected by natural acid
rain produced by Massaya volcano, Nicaragua
(Johnson and Parnell, 1986). In addition to their role
in altering acidification patterns, individual lupine
plants create “islands of soil fertility” by fixing nitrogen, incorporating soil organic matter into the soil,
and trapping nutrient-rich eolian materials. As a
result, soil solutions beneath the lupine contained
higher concentrations of dissolved organic carbon,
H2CO3/HCO3–, NH4+ and NO3– compared to barren
sites (Nuhn, 1987).
3.2 Soil development in humid-temperate climates
The second case study demonstrates soil development along a biosequence under a humid-temperate
climate in the southern Hakkoda region of northeastern Japan (Table 1). Forest vegetation is the
natural vegetation in this region, however, grass
vegetation results from human activities, such as fires
and grazing (Shoji et al., 1993). Melanudands
(Yunodai profile) form under grass vegetation
(Miscanthus sinensis) while Fulvudands (Tsutanuma
profile) form under forest vegetation (Fagus crenata).
Both soils are Andisols and andosolization is the soil
process leading to their formation (Shoji et al.,
1988b).
Precipitation, throughfall and soil solutions collected by tension lysimeters were used to decipher soil
processes in the Melanudand. Cations in the precipitation were dominated by base cations and NH4+,
while anions were dominated by Cl– and SO42–, with
lesser concentrations of NO3– (Fig. 2). The SO42–,
NH4+ and NO3– were derived from a combination of
anthropogenic pollution, ammonia volatilization from
paddy fields and gaseous emission (SO2) from fumarolic vents. Acidic components lower the pH of precipitation to values near 5. Interaction of precipitation with M. sinensis foliage effectively neutralizes the
acidity resulting in a pH of about 6.4 for throughfall.
As throughfall enters the soil, soluble components
such as DOC, base cations (Ca2+, Mg2+, K+), and P are
retained in A horizons. Soil solution exiting A horizons registers a reduction of all major cations, except
Na+. This implies active uptake of base cations by
roots to sustain the loss of base cations used by the
foliage to neutralize precipitation acidity. Concentrations of soluble Fe, Al and DOC are all very low in
Soil Development in Volcanic Ash
soil solutions exiting A horizons indicating that translocation of these components is minimal from A
horizons (Ugolini et al., 1988). Retention of Fe, Al
and organic C in A horizons is the essence of andosolization and provides a strong contrast to the
podzolization process where these components are
translocated from A/E to B horizons.
Solutions exiting the Bw horizon show a further
reduction in soluble constituents including a sharp
attenuation of silica. Retention of SO42– and DOC in
the Bw is favored by the presence of allophane/
imogolite that strongly adsorbs these constituents.
Carbonic acid is the dominant proton donor in both A
and B horizons of the Melanudand. Carbonic acid
weathering leads to incongruent dissolution reactions
in which the aluminum-rich residues interact with
silica to form allophane/imogolite. These data document that allophane/imogolite in B horizons is formed
primarily by in situ weathering as opposed to
translocation from the upper soil horizons.
The solid phase of the Melanudand has pH values
that range from 4.9 to 6.2 and only a small amount of
exchangeable Al. Extractable Al in A horizons is
mostly organically complexed, while inorganic forms
of Al (allophane/imogolite) tend to prevail in B horizons (Shoji et al., 1988b). Laminar opaline silica is
found in surface horizons where humus preferentially
retains Al (Al-humus complexes) preventing formation of aluminosilicates and allowing Si to accumulate
until it becomes supersaturated with respect to opaline
silica. M. sinensis vegetation plays a key role in
formation of Melanudands due to neutralization of
acidic precipitation by foliage and incorporation of
large amounts of organic matter from the
below-ground biomass (Shoji et al., 1990b). The
thick, dark A horizon has high organic C concentration
(Table 2) and is dominated by A-type humic acids.
Fire has been implicated as an important process
contributing to formation of A-type humic acids since
upon oxidative degradation the charred plant fragments are converted to black humic acids (Shindo et
al., 2002). Stable Al-humus complexes attenuate Al
migration and hinder organic matter decomposition.
High root and microbial respiration lead to high concentrations of H2CO3 that promotes incongruent dissolution and favors formation of allophane/imogolite.
The Fulvudand (Tsutanuma profile) formed under
a beech forest (Fagus crenata) (Table 1). Only
solid-phase characterization data were available for
this soil. The contrast between forest and grassland
vegetation demonstrates the role of vegetation in soil
forming processes. In contrast to M. sinensis that
accumulates a large portion of dead biomass (roots)
within the A horizon, litterfall from the forest
vegetation provides a large input of organic matter to
the soil surface. This is morphologically evident by
the absence of a thick dark colored A horizon (melanic
epipedon) commonly found under M. sinensis;
however, the humus content in A horizons under F.
75
crenata is still high (120 g C kg–1). Also, the
composition of the humus is different between M.
sinensis and F. crenata. Under F. crenata, fulvic acid
prevails and the humic to fulvic acid ratio of A
horizons ranges from 0.24 to 0.33 compared to 1.54 to
2.11 in A horizons under M. sinensis. Further, the
humic acid is dominated by P-type, indicating a lower
degree of humification (Kumada et al., 1967). The
pH(H2O) with the exception of a low value (4.6) at the
soil surface ranges from 5.6 to 6.0, similar to the
Melanudand profile. Enhanced acidification at the
surface is reflected in higher exchangeable Al
concentrations. Extractable Al in A horizons occurs
dominantly as Al-humus complexes, while inorganic
forms (allophane/imogolite) prevail in B horizons.
Opaline silica is also favored in A horizons due to low
Al activity (Al is complexed by organic matter), high
silica activity and seasonal desiccation which leads to
supersaturation of soil solutions with respect to
opaline silica.
In summary, the genesis of Fulvudands along this
biosequence is determined by the presence of forest
vegetation. Relative to grass vegetation, the forest
vegetation contributes larger quantities of litterfall to
the soil surface to form an O horizon, while providing
relatively lower inputs of detrital materials into the
mineral soil horizons (A horizons). Additionally,
organic matter from F. crenata has a lower humic to
fulvic acid ratio as compared to organic matter from
M. sinensis. This results in formation of a lighter
colored A horizon that meets the definition for a fulvic
horizon (Soil Survey Staff, 1999). Additionally,
there is a reduction in cycling of basic cations because
these cations are stored in the large volume of woody
above-ground biomass. This results in increased
acidity and exchangeable Al concentrations in the
surface mineral soil layer. In this same horizon,
complexation of Al by the organic matter attenuates
the availability of aluminum hindering formation of
allophane/imogolite, but fostering the synthesis of
opaline silica.
In the absence of soil solution data, but by knowing the composition of the solid phase and the distribution of the secondary minerals one can presume the
major proton donors involved in the genesis of this
Fulvudand. There are two contrasting chemical
compartments in the Fulvudand as opposed to just one
in Melanudands (carbonic acid). The upper weathering compartment (A horizons) is dominated by organic acids that are responsible for complexing Al and
depressing dissociation of carbonic acid. In contrast,
the lower compartment has pH values greater than 5.6
and is dominated by carbonic acid that favors incongruent weathering and in situ formation of allophane/
imogolite in B horizons. Thus, the influence of organic acids becomes apparent as forest vegetation
replaces grassland vegetation
(Dahlgren et al.,
1991).
76
F. C. UGOLINI and R. A. DAHLGREN
3.3 Soil development in humid-cold climates
The next example represents a further evolution of
soil development on pyroclastic materials in a
cold-humid region with coniferous forests in the
central Cascades of Washington, USA (Table 1).
Parent material consists of about 35 cm of Holocene
tephra overlying andesitic glacial drift. The soils are
dominated by Spodosols (Andic Humicryods).
There is no appreciable acidic deposition; however,
the precipitation composition is strongly influenced by
interaction with the tree canopy. On average, pH of
throughfall is about 0.5 units lower than the incoming
precipitation due to leaching of organic acids from the
canopy (Fig. 2; Ugolini et al., 1977). Further lowering of pH (pH < 4) occurs as the solution penetrates
into the forest floor (O horizons). High and low
molecular weight organic acids are primarily responsible for acidification along with storage of base
cations in woody above-ground biomass (Dawson et
al., 1978; Sletten et al., 1988). Organic acids are
responsible for weathering and metal complexaton,
and depressing the dissociation of carbonic acid.
Sequestration of metals, especially Al and Fe, hinders
their reaction with Si (Huang and Violante, 1986) and
prevents the formation of secondary aluminosilicates,
such allophane/imogolite (Dahlgren and Ugolini,
1989b,c). These soluble organo-metal complexes
remain in solution in the upper profile where the
C/metal ratio is high (Dahlgren and Marrett, 1991).
Mobile organo-metal complexes are adsorbed and/or
precipitated in the upper B horizon to form Bhs
horizons. Hence, the entire compartment starting
from the top of the canopy to the bottom of the Bhs
horizon is dominated by indigenous proton donors of
organic origin. Removal of organic acid acidity in
the Bhs horizon results in an increase of the solution
pH to values greater than 5 in the lower B horizons
(Dahlgren and Marrett, 1991). As a result, carbonic
acid becomes the dominant proton donor in the lower
portion of the soil profile.
These proton donors create two different weathering regimens within Spodosols: congruent dissolution
with organic acids and incongruent dissolution with
carbonic acid. This broad division of weathering
compartments is reflected by the mineralogy with
dissolution of primary minerals in surface horizons
and synthesis of secondary minerals in lower horizons.
The mineral assemblage of the upper compartment (E
and Bhs) dominated by organic acid weathering
includes smectite, kaolinite, hydroxy-Al interlayered
smectite (HIS), and chlorite. Opaline silica forms in
the E horizon due to summer desiccation. While
Al/Fe-humus complexes, ferrihydrite and goethite are
virtually absent from the E horizon, they are present in
the Bhs horizon. It is important to remember that
formation and growth of the Bhs horizon is due to the
accumulation of organo-metal complexes at the upper
surface of the Bs horizon. Consequently, the minerals present in the Bhs horizon have the memory of
being formed in the Bs horizon. This is particularly
the case for HIS. In the lower compartment (Bs, BC
and C) dominated by carbonic acid, HIS, chlorite and
kaolinite are the dominant layer silicate minerals
along with allophane/imogolite, ferrihydrite, and
goethite. Most of these minerals were shown to be
thermodynamically stable at the activities of Al and Si
present in the soil solution (Dahlgren and Ugolini,
1989c; Zaboswki and Ugolini, 1992).
In summary, podzolization becomes the dominant
soil forming process as soil temperatures decrease and
coniferous vegetation becomes prevalent. Organic
acids become the dominant proton donor in the upper
soil horizons promoting congruent dissolution and
translocation of weathering products to the upper B
horizons. Following removal of organic acids by
sorption/precipitation reactions, H2CO3 becomes the
dominant proton donor resulting in incongruent dissolution with minimal transport of Al and Fe. The
removal of complexing organics and pH values in the
5 to 7 range promote in situ formation of allophane/
imogolite in the lower soil horizons.
Findley Lake Spodosols have been subjected to
repeated additions of airfall tephra throughout the
Holocene. We simulated the initial stages of chemical weathering in air-fall tephra by artificially applying tephra (5 and 15 cm) from the 1980 eruption of Mt.
St. Helens to the soil surface. The primary proton
donor in the surficial tephra layer was carbonic acid
(Dahlgren et al., 1999). Solutions leached from the
tephra layer indicated incongruent dissolution resulting in formation of a cation-depleted, silica-rich
leached layer on the glass and mineral surfaces. Due
to near neutral pH values and low concentrations of
complexing organic ligands, Al and Fe were relatively
insoluble and accumulated in the tephra layer rather
than leaching. The majority of the CO2 in the tephra
layer originates from upward transport of CO2 from
the organic-rich soil horizons of the buried soil.
Elevated concentrations of CO2 beneath the tephra
layer originate from biological respiration (e.g., roots
and microbes) and from protonation of HCO3– leaching from the overlying tephra layer. Cations released
by weathering in the tephra layer (pH ≈ 6-7) leach
downward with bicarbonate to the acidic organic
horizons (pH ≈ 4) where H2CO3 reequilibrates with
the high pCO2. The gaseous CO2 diffuses upwards
and is then available to take part in another cycle of
weathering (H2CO3) and transport (HCO3–).
This example of weathering in airfall tephra
deposits demonstrates a unique weathering pathway in
which the buried organic-rich horizon pumps protons
upward to the tephra layer which then acts as an
alkaline trap for CO2. In addition, base cations
leaching from the tephra layer exchange with H+ on
the cation exchange capacity of buried organic horizons resulting in increased base saturation, somewhat
higher pH levels, and enhanced plant availability of
nutrient cations. Rapid release of nutrients from
Soil Development in Volcanic Ash
intermittent airfall tephra additions is an important
factor maintaining the nutrient status and high productivity of volcanic soils. As an organic (litter) layer
accumulates on the surface of the tephra layer, the
H2CO3 weathering regime will gradually be replaced
by an organic acid weathering regime that promotes
leaching of organo-metal complexes from the tephra
layer, transforming the tephra layer into an E horizon
(Dahlgren et al., 1997).
3.4 Soil development in perhumid and
humid-tropical climates
The final case study demonstrates soil development on basaltic cinders in a tropical climate (Table 1).
Information on these soils comes from a detailed
study of tropical volcanic soils in the New Herbides
by Quantin (1992). Soil development on these
islands is determined by the age of the volcanic
deposits and climate, the climate being modulated by
exposure of the mountain flanks to rain-bearing trade
winds. In this example we compare moderately and
highly developed soils formed under natural forest
vegetation.
First we present examples from Santa Maria Island
to characterize an advanced stage of andosolization
under tropical climates with and without a short dry
season. We chose an Andosol Désaturé Perhydratés
Typique (Santa Maria profile; Table 1) that occupies
windward and summit areas and is equivalent to
Hydrudands (Soil Survey Staff, 1999). The surface
of these soils has been rejuvenated by intermittent
volcanic activity. Soils are very friable with low
bulk density (0.2 to 0.5 g cm–3) and clay contents of
200 to 300 g kg–1 in A and B horizons. Soil organic
matter content is high with 140 to 200 g kg–1 in surface horizons and 11 to 23 g kg–1 at one meter depth
(Table 2). Water content is extremely high, reaching
values of 380% by weight! On a volume basis, water
is the principal constituent occupying from 83 to 91%
of the total soil volume. Upon air drying, water
retention may irreversibly drop to values less than
70% of the non-dried values. Irreversible changes in
physical and chemical properties of hydric Andisols is
a common feature and is believed to be due to the
presence of noncrystalline inorganic materials (e.g.,
allophane, imogolite, ferrihydrite and Fe/Al hydroxide
gels) and the microaggregation that these materials
impart to the soil structure. These noncrystalline
materials have large surface areas, ranging from 150
to 260 m2 g–1 for the <2-mm fraction, which contributes to high moisture retention.
The high intensity of chemical weathering in
tropical environments is reflected by low insoluble
residue (10 to 30%) remaining after perchloric acid
treatment. The residue decreases with depth (i.e.,
more highly weathered with depth), contrary to what
is generally expected. The inversion of this trend is
due to rejuvenation of the surface soil by intermittent
ashfall. Depth trends in the degree of weathering are
77
also reflected by SiO2/Al2O3 (0.9 to 1.4) and SiO2/
Al2O3+Fe2O3 (0.6 to 1.0) ratios of the <2-mm fraction,
which are highest at the soil surface. The low values
indicate appreciable desilication and accumulation Fe
and Al, a process similar to ferrallitization. In spite
of the high degree of weathering and leaching, the soil
pH is between 5 and 6, evidently buffered by the
presence of noncrystalline materials. These materials display a high degree of variable charge (∆CEC)
and high phosphorous retention. Secondary minerals
are dominated by noncrystalline materials that increase in concentration with depth. These materials
are rich in Al and Fe and were referred to as Al and Fe
allophanes by Quantin (1992). Primary minerals are
abundant only in the upper horizons due to accretion
of fresh ash. Allophane dominates in the upper soil
layers with lesser concentrations of imogolite and
layer silicate clays. Below the A horizon, imogolite,
allophane and Fe/Al hydroxide gels dominate, whereas gibbsite and layer silicates (halloysite, kaolinate
and smectites) are present at trace levels. Because of
the dominance of Fe hydroxide gels, only trace
amounts of crystalline Fe oxides and hydroxides (lepidocrocite, goethite and hematite) are present.
Soils developed on the same parent material but on
the leeward side of the island are exposed to a short
dry season, typically 1 to 2 months. Compared to
their windward soil counterparts, these soils display
less desilication and consequently less enrichment of
Al and Fe. Formation of Fe-rich 2:1 layer silicates
increases the Fe content in the clay fraction. These
2:1 clay minerals are poorly crystallized and may
form mixed-layer clays with halloysite (Delvaux and
Herbillon, 1995). Nonetheless, layer silicate clays
prevail over allophane and Fe oxides. At lower elevations where the dry season is even more pronounced
there is an increase in the prevalence of dehydrated
halloysite and Fe-rich smectite. Thus, the occurrence
of a distinct dry season impedes the intensity of
chemical weathering and enhances formation of crystalline layer silicate minerals at the expense of noncrystalline materials.
The highest level of alteration recorded by Quantin
(1992) for basaltic material in the New Hebrides is
represented by Ferrallitic soils. These soils form
under a perhumid climate regimen and also on older
parent material. The example presented is from the
Island of Erromango, and soils are classified as ferrallitic strongly unsatureted (Ferrallitiques fortement
désaturé; Erromango profile; Table 1). These soils
approach Oxisols in Soil Taxonomy but CEC values
exceed Oxisol criteria resulting in their classification
as Inceptisols. These soils have experienced intense
alteration to the extent that primary minerals have
completely disappeared. The residue after perchloric
acid treatment in B and BC horizons is less than 1%.
Total elemental analyses show that SiO2 is between 33
and 46% while Al2O3 and Fe2O3 are 27 to 40% and 16
to 31%, respectively. Base cations (Ca, K and Na)
78
F. C. UGOLINI and R. A. DAHLGREN
have virtually disappeared while Mg is present in
trace amounts.
The surface horizon has between 40 to
100 g C kg–1 and a strongly acidic pH of 4 to 5 (Table
2). Cation exchange capacity ranges between 8 and
20 cmolc kg1, while the base saturation ranges from 4
to 19%. Clay mineralogy in the upper horizons is
dominated by halloysite (1.0 nm) that partially dehydrates to 0.75 nm, gibbsite and poorly crystallized
hematite, kaolinite and goethite. The clay-size fraction in the B and BC horizons consists of disordered
kaolinite and halloysite with goethite and hematite.
Formation of halloysite in the surface horizons is
believed to be due to the input of fresh cinders. As
expected, no appreciable allophane or imogolite was
detected in these highly weathered soils. Thus, as
volcanic soils become progressively more weathered,
metastable noncrystalline materials are transformed to
more thermodynamically stable mineral phases. This
transformation results in loss of andic soil properties
and the conversion of many tropical Andisols to Inceptisols. Formation of Oxisols in the tropical environment primarily occurs on highly weathered stable
landscapes that have escaped rejuvenation by periodic
ashfalls or erosion (Beinroth, 1982).
4. General Trends in Soil Development on
Volcanic Materials
Climate, in conjunction with its influence on vege-
tation, and time of exposure to weathering are the two
primary factors regulating soil development pathways
in volcanic materials (Table 3). Andisols generally
form rapidly in humid climates and alter to other soil
orders as soil age and degree of weathering increases.
However, in regions with intermittent deposition of
volcanic ash each addition of new material rejuvenates
soil developmental processes so that Andisols may be
maintained as a relatively stable soil condition. Soil
development is stalled in the Entisols/Vitrands stage
under dry climate conditions as weathering is limited
by moisture. Under cool to cold-humid conditions,
Spodosols are often the dominant soil as cooler temperatures favor coniferous vegetation and predominance of the organic acid weathering regime.
Spodosols may form directly from fresh volcanic
deposits or as an alteration product from Andisols.
Mollisols form in volcanic materials under dry conditions in temperate and tropical environments and
under wetter conditions in the present of basic volcanic ash. Warm/dry conditions promote formation of
crystalline layer silicates rather than noncrystalline
materials and leaching is limited leading to high base
saturation. Vertisols may also form in warm regions
having a distinct dry season as volcanic ash weathers
preferentially to smectite and vermiculite rather than
noncrystalline materials. Andisols are the typical
product of weathering in both temperate and tropical
environments with sufficient moisture. With increased weathering intensity, metastable noncrystalline
Table 3 Examples of soil genesis pathways in contrasting climates extrapolated from the literature.
Climatic Regime
Warm/Arid & Semi-arid
Arid (400 mm):
Entisols
Semi-arid (700 mm): Ustivitrands
Literature Source
Dubroeucq et al. 1998
Dubroeucq et al. 1998
Cold/Dry
Entisols Æ Vitricyrands
Arnalds and Kimble, 2001
Cold/Humid
Entisols Æ Spodosols
Entisols Æ Andisols Æ Spodosols
Ping et al. 1988; 1989; Shoji et al. 1988a
Ping et al. 1988; 1989; Shoji et al. 1988a
Humid/Temperate
Shoji et al. 1990a
Warm/dry:
Entisols Æ Andisol Æ Mollisols
Numerous reports
Warm/moist: Entisols Æ Andisols
Warm/moist: Entisols Æ Andisols Æ Inceptisols Æ Alfisols/Ultisols Southard and Southard, 1987; Takahashi et al., 1993;
Chen et al. 2001
Shoji et al. 1988b; Takahashi et al. 1989
Cool/moist:
Entisols Æ Andisols Æ Spodosols
Ugolini et al. 1977; Parfitt and Saigusa, 1985;
Cold/moist:
Entisols Æ Spodosols
Shoji et al. 1988b
Tropical
Warm dry/moist: Entisols Æ Mollisols
Entisols Æ Andisols Æ Mollisols
Warm dry:
Entisols Æ Vertisols
Warm/moist: Entisols Æ Andisols Æ Inceptisols Æ Alfisols/Ultisols
Warm/moist: Entisols Æ Andisols Æ Inceptisols Æ Oxisols
Wielemaker and Wakatsuki, 1984;Yerima et al. 1987;
Otsuka et al. 1988; Chadwick et al. 1994;
Yerima et al. 1987
Martini, 1976; Delvaux et al. 1989
Quantin, 1992; Chadwick et al. 1994; Van Ranst et al.
2002
Martini, 1976; Beinroth, 1982; Kimble and Eswaran, 1988
Soil Development in Volcanic Ash
materials are gradually consumed by transformation to
more stable crystalline minerals. This often leads to
the alteration of Andisols to Inceptisols. On stable
landscapes in the perhumid tropics, volcanic soils may
transform to Oxisols. In the xeric moisture regime of
California, clay translocation in Inceptisols leads to
formation of Alfisols/Ultisols. Clay translocation is
not a prevalent process in Andisols due to the difficulty in dispersing noncrystalline materials. However, once noncrystalline materials are altered to layer
silicate minerals, translocation is greatly enhanced,
especially in the xeric moisture regime where repeated
wetting and drying leads to slaking and mobilization
of clays.
5. Environmental Implications
Formation of volcanic soils assumes the deposition
of volcanic ejecta on the landscape. The addition of
fresh material not only counteracts soil erosion but
also provides a new substrate to rejuvenate soil processes and sustain productivity of terrestrial ecosystems. Volcanic soils are among the most productive soils for agriculture and forestry. The high potential of these soils for agricultural production is
illustrated by the fact that many of the most productive agricultural regions of the world are located
near active or dormant volcanoes and the most densely populated areas in regions, such as Indonesia, are
found near volcanoes (Shoji et al., 1993). Intermittent additions of volcanic ash renew the long-term
fertility status of terrestrial ecosystems by providing a
source of nutrients from the rapid weathering of
volcanic ejecta (Dahlgren et al., 1999). Following
the 1980 eruption of Mt. St. Helens, wheat fields in
eastern Washington (USA) recorded a bumper crop
after deposition of a few centimeters of tephra that
acted as a mulch by killing weeds and conserving
moisture. Similarly forest productivity throughout the
semi-arid western United States is enhanced by
surficial deposits of volcanic ash that act as a mulch
and weather to soils with high plant-available water
retention (Meurisse, 1988).
Volcanism is probably the most obvious mechanism of recycling large amounts of geological material
and gases (e.g., CO2, SO2) on the planet Earth. By
providing a source of easily weatherable materials,
this natural phenomenon represents an effective sink
for carbon dioxide through carbonic acid weathering
(e.g., CO2 (g) + H2O = H+ + HCO3–). Soils formed
in volcanic ejecta also contain the largest
accumulations of organic carbon among the mineral
soil orders (Eswaran et al., 1993). Organic matter
preservation results from burial of soils by repeated
additions of volcanic ash, chemical interactions with
noncrystalline inorganic materials (e.g., allophane,
imogolite, ferrihydrite) and physical protection from
the microaggregation that these materials impart to the
soil structure. Thus, volcanism plays an important
79
role in the global carbon cycle, representing a primary
source and sink for carbon.
The mineralogy of volcanic soils is commonly
dominated by noncrystalline materials having high
surface areas and highly reactive, variable charge
surfaces that may contribute both cation and anion
exchange capacity. Because of this unique colloidal
assemblage, volcanic soils can often mitigate the adverse effects of anthropogenic pollution. Acidic deposition is effectively neutralized by rapid chemical
weathering reactions, pH buffering by variable charge
materials and sorption of SO42–. Similarly, noncrystalline materials along with high concentrations of
soil organic matter have a high capacity to retain
heavy metals, trace elements (cations and anions) and
organic compounds of both natural and anthropogenic
origins. These properties of volcanic soils attenuate
many pollutants, hence preventing their transmission
to surface and ground waters.
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(Received 3 September 2002, Accepted 10 October 2002)