Classification and geochemistry of arid
and semi-arid paleosols
Veronika Geißler, TU Bergakademie Freiberg
Abstract. Many authors (Mack et al. 1993, Retallack 1998) argued, that for the classification of
paleosols it is useful to operate with the existing soil taxonomy, because then analogies between
paleosols and recent soils can be drawn. Comparison is especially helpful for identifying conditions of
formation, which are precipitation, temperature, parent material, morphology of landscape and time.
These information can be employed for the reconstruction of paleo-environments.
Classification is needed for soil types (paleosols respectively), but today the huge amount in use
complicates communication and comprehension. The attempt of this paper is to give an overview of
the most important soil types (after Soil Survey Staff 1975) for semi-arid and arid climates, adding
common synonyms applied in other classifications.
Introduction
There are several approaches to paleosols. From a sedimentological point of view, they represent a gap
in sedimentation. Furthermore paleosols can be employed as lithostratigraphic markers or to define
paleo-morphology. In this paper special emphasis is laid on the possibilities paleosols offer for the
reconstruction of paleo-climate. The formation of soils is mainly influenced by agents as precipitation,
temperature, parent material, morphology of landscape and time. Therefore paleosols provide an
excellent reservoir of information on weathering and climate conditions.
Classification
Not all current classifications are suitable for the application on paleosols. It is necessary, that
classification is based on criteria, which can be observed in paleosols. Therefore U.S. Soil Taxonomy
seems to be an appropriate choice. The twelve orders differentiated by U.S. Soil Survey Staff are not
directly defined due to their climatic occurrence, but can be related to it (see Table 1). The marked soil
types will be discussed in detail in the following paragraphs.
Table 1: Soil types with description and relation to climate
Soil type
Description
(Soil Survey Staff 1975)
(Retallack 1998)
Aridisol
desert soil
Mollisol
grassland soil
Vertisol
swelling clay soil
Alfisol
fertile forest soil
Ultisol
base-poor forest soil
Spodosol
sandy forest soil
Oxisol
tropical deeply weathered soil
Histosol
peaty soil
Gelisol
permafrost soil
Entisol
incipient soil
Inceptisol
young soil
Andisol
volcanic ash soil
Climate
(Retallack 1998)
arid
semi-arid - subhumid
seasonal dry
semi-arid - subhumid
humid
humid
humid
humid
nival
no indication
no indication
no indication
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Veronika Geißler, TU Bergakademie Freiberg
Aridisol
Common synonyms for Aridisols are Cryids, Salids, Durids, Gypsids, Argids, Calcids, Cambids (all
United States Department of Agriculture, in the following abbreviated by USDA), Yermosol (Food
and Agricultural Organisation, in the following FAO), Xerosol (FAO), Calcixeralfs (USDA), Calcisol
and Gypsisol.
a)
Fig. 1: Aridisol: a) photography of typical Aridisol (from www.wikipedia.de ) b) class sketch, for legend see
Fig. 2, climate is represented by sun/ cloud symbols, time of formation by an hourglass (from Retallack 1997)
c) sketch of petrography and soil texture of horizon A in thin section (from Retallack 1997)
Fig. 2: Legend for class sketches (from Retallack 1997)
Aridisols are desert soils. Depending on the parent material, a relative short time of development is
required. They are common in arid and semi-arid climates, where as a consequence of the lack of
rainfall easy soluble salts can endure.
The surface horizon shows a light color. Its texture is soft and often vesicular (McFadden et al. 1998).
Salinization can occur, if salts accumulate on the surface (www.wikipedia.de). Characteristic for the
subsurface layers are shallow calcareous (Bk), gypsiferous (or anhydritic) (By) or salty (Bz) strata (see
Fig. 1), which represent the depth of wetting during occasional rainfall and can therefore be correlated
to the amount of mean annual precipitation (see also Depth to calcic horizon) (Retallack 1997). The
appearance of cemented horizons varies from needle-fibre calcite over nodules to distinct hardpan
calcretes (Wright and Tucker 1991) (See also Calcic horizon as indicator for the maturity of
paleosols). Not cemented or clayey subsurface layers are also common. The material for the formation
of either clayey, calcareous or salty horizons respectively derives from weathering or eolian
deposition. After the latter extremely fine grained dust of easily weatherable minerals flushes down
the profile and accumulates. On Aridisols only sparse vegetation grows, which includes prickly shrubs
and cacti. Therefore root traces are not particularly common in paleosols. In general Aridisols have a
very poor concentration of organic matter. Geomorphical settings for Aridisols are low lying areas,
since on slopes erosion would prevent soil maturation (Retallack 1990).
Classification and geochemistry of arid and semi-arid paleosols
3
Mollisol
For Mollisols the following synonyms are used:
Albolls, Aquolls, Rendolls, Cryolls, Xerolls, Ustolls, Udolls (all USDA), Chernozem, Solonetz
(Canadian system of soil classification) Kastanozem, Phaeozem, Rendzinas, Greyzem and Chernozem
(all FAO).
a)
Fig. 3: Mollisol: a) photography of typical Mollisol (from www.wikipedia.de ) b) class sketch, for legend see
Fig. 2, climate is represented by sun/ cloud symbols, time of formation by an hourglass (from Retallack 1997)
c) sketch of petrography and soil texture of horizon Bk in thin section (from Retallack 1997)
Mollisols form in grassland environments. Therefore their occurrence in the geological history is
connected to the evolution of grass in Tertiary. Their development takes place in relative short time, of
course in dependence of their host material. For Mollisols base-rich parent materials such as clay, marl
or basalt are common. They are frequent under subhumid - semi-arid climate conditions.
The surface layer is base-rich and consists of a mixture of clay and organic matter, which is not
carbonaceous or coaly (Retallack 1990). This A horizons usually shows granular or crumb bed
structures (Mack et al. 1993). Other significant criteria of Mollisols are abundant fine root traces of
grassy vegetation and burrows of diverse populations of invertebrates. The subsurface can be either
argillic (Bt), calcareous (Bk) or gypsiferous (By), but mostly an A-Bk-profile is found (See Fig.3).
Mollisols form under a wide range of temperature in lowland to high mountain meadows (Retallack
1990).
Vertisol
For Vertisols a lot of local names exist as synonyms:
cracking clays (Australia), Adobe (Philippines), Shachiang (China), Black Cotton Soils (India),
Smolnitza (Bulgaria, Rumania), Tirs (Marocco), Makande (Malawi), Vleigrond (Southafrica),
Sonsosuite (Nicaragua), Margalite soils (Indonesia), Densinegra soils (Angola), Grumusols (United
States), in FAO also named Vertisol, Aquerts, Cryerts, Xererts, Torrerts, Usterts and Uderts (all
USDA)
Vertisols are defined as swelling clay soils. Depending to their parent material, which is commonly of
intermediate or basaltic composition, the time span for formation varies extremely between hundreds
and thousands of years. But they still require less time for generation than Aridisols and Mollisols.
Vertisols occur in subhumid – semi-arid climates with a pronounced dry season. Precipitation must be
enough to form clays, which means about 180-1520mm mean annual rainfall. Vertisols are
characterized by a uniform, at least half meter thick clayey horizon, which shows desiccation cracks
during the dry season (Retallack 1990). The shrinking and swelling results in constant mixing, which
causes Vertisols to have an extremely deep A horizon and no B horizon (Soil Survey Stuff 1999). The
4
Veronika Geißler, TU Bergakademie Freiberg
a)
Fig. 4: Vertisol: a) photography of typical Vertisol (from www.wikipedia.de ) b) class sketch, for legend see
Fig. 2, climate is represented by sun/ cloud symbols, time of formation by an hourglass (from Retallack 1997)
c) sketch of petrography and soil texture of horizon Bt in thin section (from Retallack 1997)
desiccation cracked surface forms a typical hummock-and-swale topography, also known as gilgai
microrelief (See Fig.9). The corresponding subsurface also displays the disrupted, festoon-shaped
texture (Mack et al. 1993). Slickensided profiles with internal deformations of horizons are frequent
(Retallack 1990) (See Fig.4). Responsible for this appearance is the high content of clays (50-70%)
and its swelling properties respectively. The clays consist predominantly of 2:1 and 2:2 layer clay
minerals. Their color ranges from grey to red or the more familiar black (Soil Survey Staff 1999).
Because of the dry climate and the sparse vegetation alkaline reactions and good reserves of
exchangeable cations can be maintained (Driese et al. 2000). As geomorphological setting mainly flat
regions and the feet of gentle slopes are possible (Retallack 1997).
Alfisol
Fig. 5: Alfiisol: a) photography of typical Alfisol (from www.wikipedia.de ) b) class sketch, for legend see
Fig. 2, climate is represented by sun/ cloud symbols, time of formation by an hourglass (from Retallack 1997)
c) sketch of petrography and soil texture of horizon Bt in thin section (from Retallack 1997)
Classification and geochemistry of arid and semi-arid paleosols
5
Synonyms for Alfisols are:
Luvisols, Nitosols, Solonetz (all FAO), Aqualfs, Cryalfs, Udalfs, Ustalfs and Xeralfs (all USDA).
Alfisols are defined as fertile forest soils. The prefix alf derives in this case not from the word pedalfer
(Marbut, 1935, see also Pedocals and pedalfers), but refers to aluminium (Al) and iron (Fe). For
generation Alfisols require relative short time. Their parent material is mainly base-rich as the soil
itself. Alfisols are common in subhumid – semi-arid climates.
The surface layer is of light color. The subsurface horizon is mostly argillic (Bt) (See Fig. 5).
Especially base-rich clays (smectites) are abundant (Retallack 1990). Alfisols represent the youngest
forest soil order. Because of this they are less leached, and have a greater than 35% base saturation
(www.wikipedia.de). Therefore the soil contains lots of exchangeable cations. A geochemical
characteristic of Alfisols are molecular weathering ratios of alumina/ bases of less than two.
Vegetation of Alfisols ranges from grassy woodland to open forest. The topography, in which Alfisols
form, varies (Retallack 1990).
Common subordinate modifiers of paleosols (from Mack et al. 1993)
A finer classification of paleosols is not only possible by employment of subgroups, but also by
application of subordinate modifiers (see Table 2).
Table 2: Common subordinate modifiers of paleosols (Mack et al. 1993)
Presence of eluvial (E) horizon
albic
Presence of allophone or other Si and Al compounds
allophanic
Presence of illuvial clay
argillic
Presence of pedogenic carbonate
calcic
Presence of dark organic matter, but not coal
carbonaceous
concretionary Presence of glaebuls with concentric fabrics
Low base status as indicated by the paucity of chemically unstable grains such as feldspar
dystric
and volcanic rock fragments
High base status as indicated by the abundance of chemically unstable grains such as
eutric
feldspar and volcanic rock fragments
Presence of iron oxides
ferric
Subsurface horizon that was hard at the time of soil formation (for example root traces and
fragic
burrows terminate or are diverted at this horizon)
Evidence of periodic waterlogging, such as drap hues; mottles of drap color and yellow, red
gleyed
or brown; or presence of pedogenic pyrite or siderite
Presence of vadose gypsum or anhydrite
gypsic
Presence of glaebuls with an undifferentiated internal fabric
nodular
Presence of light-colored A horizon
ochric
Presence of pedogenic salts more soluble than gypsum
salic
Presence of pedogenic silica
silicic
Presence of decimeter-scale desiccation cracks, wedge-shaped peds, hummock and swale
vertic
structures, slickensides, or clastic dikes
Vitric
Presence of relict or actual glass shards of or pumice
Calcic Horizon as indicator for the maturity of paleosols
Soil profiles of arid and semi-arid climates often contain pedogenic calcretes (Bk). They typically
occur within Aridisols, Vertisols and Mollisols.
In the following table pedogenic calcretes are classified according to their morphology, with the
system devised by Netterberg (1967, 1980), which was developed for geotechnical surveys.
The morphology of calcretes relates to stages seen in the development of calcrete profiles (Netterberg,
1980). Manchette has provided the most comprehensive sequence and has recognized six stages of
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Veronika Geißler, TU Bergakademie Freiberg
development for gravelly parent material. Retallack extended this sequence for sandy and finer grained
parent material (see Fig. 6).
Table 3: Morphological classification of calcretes based on Netterberg (1967, 1980) and Goudie (1983) with
classification of pedogenic calcretes based on stages of development after Machette (1985). Modified after Quast
(2003)
Calcrete types
Maturation
Description (after Netterberg 1980 and Goudie 1983)
stages (after
Machette
1985)
Calcareous soil
Very weakly cemented or uncemented soil with small carbonate accumulations as
grain coatings, paches of powdery carbonate including needle-fibre calcite
(pseudomycelia), carbonate-filled fractures and small nodules
Calcified soil
A firmly cemented soil, just friable; few nodules. 10-50% carbonate
Stage 1
Powder
A fine, usually loose powder of calcium carbonate as a continuous body with little
calcrete
or no nodule development
Pedotubule
All, or nearly all, the secondary carbonate forms encrustations around roots
calcrete
or fills root or other tubes (tubules)
Nodular
calcrete
Stage 2
Honeycomb
calcrete
Stage 3
Hardpan
calcrete
Laminar
calcrete
Boulder/
cobble calcrete
Stage 4
Stage 5
Stage 6
(syn. glaebular calcrete of Netterberg, 1980) Discrete soft to very hard concretions
of carbonate-cemented and/ or replaced soil. Concentrations may occur as
laminated coatings to form pisoids
Partly coalesced nodules with interstitial areas of less indurated material between
(syn. petrocalcic horizon) An indurated horizon, sheet-like. Typically with
a complex internal fabric, with sharp upper surface, gradational lower surface
Indurated sheets of carbonate, typically undulose. Unsually, but not always, over
hardpans or indurated rock substrates
Disrupted hardpans due to fracturing, dissolution and rhizobrecciation (including
tree-heave). Not always boulder grade. (Clasts are rounded due to dissolution)
Fig. 6: Stages of the morphology of carbonate accumulations in soils a) carbonate stages of Machette for
gravelly parent material b) soil development stages of Retallack for sandy and finer grained parent material
(from Retallack 1997)
Classification and geochemistry of arid and semi-arid paleosols
7
Implications of paleoclimate (after Retallack 1990)
Indicators of precipitation
The availability of water plays a major role in many soil-forming chemical reactions, which proceed in
dilute solutions. Reacted products are removed or concentrated. Therefore it is no wonder that freely
drained soils of humid climates show more profound chemical weathering than those of dry climates.
Many chemical and mineralogical characteristics of soils can be associated to precipitation, but here
only the most reliable and important are presented.
Pedocals and pedalfer
In soils of dry climates, calcium carbonate is an abundant component, whereas in humid climate it is
dissolved, transported away and therefore not found. Marbut (1935) used this observation to divide
soils into two main groups. Pedocals have free carbonate, while pedalfers lack it. Although calcareous
paleosols are a reliable indicator of dry climate and noncalcareous paleosols of wet climates, it is not
possible to specify a dividing line accurately without information on other aspects of paleo-climate
(mean annual temperature, season of main rainfall, etc.).
Depth to calcic horizon
Free carbonate in soils usually forms a distinct calcareous layer or calcic (Bk) horizon. The position of
this horizon within the soil profile reveals the depth of wetting of the soil by available water.
Therefore in dry climates the calcic layer is closer to the surface than in wetter ones (see Fig. 7). But
this rule applies only for soils of moderate development, which include nodules of carbonate rather
than wisps or layers, in unconsolidated parent material and for soils of seasonal warm climates. And
here still some difficulties appear. Erosion before burial falsifies the depth of the calcic horizon, as
well as compaction does. A third problem are higher CO2 levels in the past, which lead to a deeper
calcic horizon. But only in extreme greenhouse periods of the Jurassic-Cretaceous, OrdovicianSilurian and perhaps also early Precambrian this is a severe problem (Ekart et al. 1999).
Fig. 7: The relationship between mean annual rainfall and depth of the calcic horizon in 317 surface soils from
all continents (from Retallck 1998)
Chemical composition
Weathering and soil formation leave characteristic imprints in the chemical composition of soils.
Molecular weathering ratios reveal changing chemical proportions as a result of processes or
properties like salinization (Na2O/ K2O), calcification ((CaO + MgO)/ Al2O3), clayeyness (Al2O3/
SiO2), base loss (Al2O3/(CaO + MgO + Na2O + K2O)), leaching (Ba/ Sr) and gleization (FeO/ Fe2O3).
Fig. 8 shows an example how molecular weathering ratios can contribute to a more precise image of
the conditions during soil fomation.
Conclusions for climate can be drawn from the degree of depletion of alkali (Na+, K+) and alkali earth
(Ca2+, Mg2+) elements normalized to silica and alumina, since in dry climates a lack of water
prevents chemical weathering and transport of these elements.
8
Veronika Geißler, TU Bergakademie Freiberg
Fig. 8: Example of selected molecular weathering ratios for four paleosols in the Upper Hell Creek Formation,
Montana (from Retallck 1994)
Strong correlations between the molecular ratio of bases/alumina and mean annual rainfall can be
observed in soil subsurface (Bt) horizons (Marbut 1935). But this relationship doesn’t hold for
paleosols of either very wet (with common kaolinite) or dry climates (with evaporites) and for very
strong or very weak developed paleosols, or those altered during burial by illitization.
Clay minerals
Clays, regarded as products of hydrolysis of weatherable minerals, can be linked to the amount of
precipitation available for soils (Folkoff & Meetenmeyer 1987). Clay composition in soils is
controlled by the grain size and mineral composition of parent materials, temperature, seasonality of
rainfall and time for formation of a soil. For paleosols additional problems have occur, that is to say
whether the clay formed during soil diagenesis or is inherited from parent material or altered after
burial. With theses cautions in mind some generalized statements can be made.
In wet climates clay are more likely to possess a 1:1 rather than a 1:2 layer structure. They obtain
fewer cations and are lower in general weathering sequence of clay sized minerals. Especially in
Vertisols, or generally in dry climates, swellable clays like sodium smectites produce a distinctive soil
structure of domed columnar peds extending throughout much of the solum (natric horizon: McCahon
& Miller 1997). In Table 4 indicative clay minerals are presented.
Table 4: Indicative clay minerals
Dominant clay mineral
Indicative for
Palygorskitea, sepiolitea
Smectite
Kaolinite
Iron oxides and alumina
a
Very arid climate
Mean annual rainfall < 1000mm
Mean annual rainfall between 1000 - 2000mm
Mean annual rainfall > 2000mm
These clay minerals are easily dissolved and not stable during burial (Botha, Hughes 1992)
Classification and geochemistry of arid and semi-arid paleosols
9
Evaporate minerals
Evaporate minerals are widespread in climates, in which evaporation exceeds precipitation. Due to the
lack of water, evaporate minerals and mainly salts can accumulate respectively, when a source of salts
is present. Gypsum is very abundant, but also halite, sylvite or mirabilite. Like in calcic strata, the
depth of the gypsic horizons gives a hint to the rate of rainfall.
Salts as gypsum are easily soluble in groundwater. As residue of many ancient evaporate horizons
only pseudomorphs of crystals or zones of breccia are left, where the overlying rock has collapsed into
the dissolved layer zone (Bowles, Braddock 1963).
Desert pavement
When climates is too dry for constant vegetation soils are covered by a natural plaster of stones,
developed due to eolian deflation. In such environments windsculpted stones (ventifacts), carbonate
collars (rims of calcite around stones at ground level), rock varnish (armorphous iron-manganese
crusts on top of stones) and vesicular structure are common (McFadden et al. 1998).
Indicators of temperature
Even though water is the agent for most soil-forming processes, the rate of reactions depends on the
temperature. Refering to Van t’Hoff’s rule for temperature, for every 10°C increase in temperature the
rate of reaction is doubled or even tripled. This coincides well with the observation that in warm or hot
regions soils are much more deeply weathered than in cold areas (Birkeland 1992).
Oxygen isotopic composition
With rising temperature the activity of the heavy oxygen isotope 18O increases relative to the light
isotope 16O. In soil minerals or fossil biominerals the ratio of these isotopes (δ18O) is preserved. But
the temperature of formation can only be concluded from the δ18O-proxy, if the isotopic composition
of the water from which the mineral precipitated is known. This original water composition is
conserved in aragonite shells of bivalve or gastropod. But even if isotopic composition does not reveal
absolute temperatures, it can still be useful to detect relative changes in temperature through time.
Still caution is necessary, since the oxygen isotopic composition is not stable in the face of elevated
temperatures and fluid migration during burial and metamorphism.
Indicators of seasonality
Seasonality produces varying amounts of rainfall, of heat, of dust influx and other agencies of soil
formation that are essential to clayeyness, redness and base saturation of soils. But also other factors
contribute to these features, therefore it is difficult to tease out the role of seasonality. Here only the
most diagnostic features are reviewed.
Mukkara and gilgai
Gilgai microrelief and mukkara subsurface structure (see Fig.9) occur in swelling clay soils as
indicators of a climate with pronounced dry and wet season. These structures are not found in
extremely arid or humid climates, since at least some moisture is needed to provoke clay formation
and provide a contrasting wet season. In humid climates clays become too deeply weathered and lose
their swelling properties.
Fig. 9: Sketch of gilgai microrelief and mukkara subsurface structure (Retallack 1997)
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Veronika Geißler, TU Bergakademie Freiberg
Concretions and argillans
Seasonal differences in the chemical condition of the soil may be reflected in concentric bands of concretions and strong banding of clay skins (argillans) (Nahon 1991). Especially in well-drained soils of
subtropical, monsoonal climates calcareous nodules, which develop under alkaline conditions of dry
climates, and ferruginous concretions occur, which in contrast formed in the wet season (Sehgal,
Stoops 1972).
Patterns of root traces
A characteristic root pattern for seasonally dry climates is recognizable in paleosols. Grasses and trees
possess a profuse surficial network of roots that supplies them with water during the wet part of the
year. In the dry period grasses weather back to their root stock, whereas some trees obtain moisture
through especially stout, deeply penetrating roots, termed “sinkers” (van Donselaar-ten Bokkel
Huinink 1966). Also soils of dry swamps show a similar bimodal distribution pattern.
Carbon isotopic composition
In soils two different carbon isotope compositions are present. The first derives from soil organic
matter, which reproduces the annual average composition of vegetative biomass (δ13Corg). The second
reflects the amount of CO2 dissolved during dry-season from root respiration and microbial decay of
organic matter (δ13Ccarb). The difference between both values (δ13Corg - δ13Ccarb) within one soil profile
shows the extent of seasonal change in precipitation, especially in strongly seasonal monsoonal
climates. The difference is also connected to (i) diffusion of atmospheric CO2 in correlation to
atmospheric partial pressure and (ii) equilibrium fractionation of CO2 for the temperature of formation.
Seasonal changes can be detected as relative changes in isotopic values from succession of paleosols
or modeled by assuming rates for diffusion and fractionation.
Indicators of greenhouse atmosphere
Methane and CO2 can be used as indicators for greenhouse atmosphere. Both gases were found in
bubbles within amber (Landis et al. 1996) in paleosols, but these observations are rare due to the
scarcity of amber. But their distinctive isotopic trace in the carbon isotopic composition of paleosols is
detectable. Minerals like calcite, etc. conserve the isotopic composition, therefore it is possible to
reconstruct former atmospheric contents of CO2 by assuming typical isotopic values of atmospheric
and soil gases. Atmospheric methane usually leaves an isotopic trace in paleosols that is extraordinary
depleted or light in isotopic values (less than -36‰ to as low as -80‰).
Alteration after burial (after Retallack 1990)
Some of the features used for interpreting paleo-environments from paleosols such as the amount of
organic carbon, the oxidation state and the composition of clay minerals, are prone to alteration after
burial. Fortunately, there remains an impressive residue of soil features that can be used for paleoenvironmental reconstruction, e.g. physical features such as nodules, root traces and burrows are particularly robust in the face of alteration after burial. In any case, for a correct interpretation it is necessary to distinguish between primary soil forming processes and those which occurred during burial.
The most important alterations after burial are:
Burial decomposition of organic matter, burial gleization of organic matter, burial reddening of iron
oxides and hydroxides, cementation of primary porosity, compaction by overburden, illitization of
smectite, zeolitization and celadonation of volcanic rocks, coalification of peat, kerogen maturation
and cracking, neomorphism of carbonate and metamorphism
Conclusion
Paleosols are found in many terrestrial setting. Their generation was mainly influenced by agents as
precipitation, temperature, parent material, morphology of landscape, time and later possibly alteration
after burial. Therefore particular caution is necessary to distinguish primary and secondary features of
soil formation. But as long as we are able to differentiate between both and interpret soil forming
factors, a lot of information for the reconstruction of paleo-climate can be gained from the study of
paleosols.
Classification and geochemistry of arid and semi-arid paleosols
11
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