Bergaya F. Handbook of Clay Science

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disordered illite–smectite minerals. One of these has about 50% smectite layers; the

other has 20% smectite and is less abundant. Both types co-exist in the same ho-

rizon, with beidellite and montmorillonite as the expanding layers. Smaller amounts

of illite, kaolinite, and vermiculite are also present, but rarely chlorite. Irrespective of

climate (continental) and parent material, a convergence in mineralogy towards

mixed-layer phases seems to occur. That is, there is a tendency towards the same

chemical and mineralogical equilibrium, be it metastable or not. Nevertheless, the

clay minerals along a soil proﬁle often show great diversity because of the complex

transformation and reaction processes that operate with increasing depth.

In the weathering environment, most clays are formed by incongruent dissolution

of unstable silicates. Reaction of the unstable mineral with less soluble compounds

can produce new clay minerals. Direct precipitation from solution can also give rise

to neoformed clay minerals with compositions that are very different from those of

the (dissolved) parent minerals. Eventually, a clay mineral can become unstable

because of intense drainage, forming oxides and liberating most ions.

Hydrolysis generally involves the exchan ge of protons from the solution for sol-

uble mono- and divalent cations in the minerals. Thus, the ﬂow rate and acidity of

water play a key role in the formation of clay minerals. In a soil proﬁle developed

under high rainfall, kaolinite is found in the upper part, while montmorillonite

commonly occurs in the lower part of the weathering proﬁle, and closer to the rock

where the accumulation of leached ions can produce more chemically complex clay

minerals. Leaching is stronger when the topography is smooth or plain, and kaolinite

can predominate in soils of slight slope, while smectites are dominant in soils formed

in depressions (Kantor and Schwertmann, 1974).

In the classic example of granite weathering, a large part of the soils consists of

metastable muscovite, biotite, and chlorite. These minerals alter progressively into

clay minerals. Muscovite is degraded into illite, and both minerals can transform into

regular or random illite–smectite (I/S) mixed-layer minerals. Biotite and chlorite

transform into smectite through the formation of a regular mixed-layer phase

containing biotite and smectite, or chlori te and smectite. Corrensite, a 1:1 chlo-

rite–saponite mixed-layer mineral, is very common in soils. In many cases all these

primary minerals (muscovite, biotite, chlorite) can be transform ed into vermiculite.

When rainfall is high, muscovite and biotite eventually transform into kaolinite in

the upper part of the proﬁle. Like the micas, plagioclase and orthoclase are also

destabilized and transform into kaolinite and some illite. Kaolinite decreases to-

wards the lower part of the alteration proﬁle, while illite, I/S mixed-layer minerals,

vermiculite, and neoformed clay minerals (smectites) tend to increase with depth.

This evolution pathway also typiﬁes the weathering of gneisses, and other types of

aluminous pellitic rocks.

In basic rocks, smectites (Fe-beidellite, nontronite) and vermiculite can form from

Ca-plagioclases, amphiboles, and pyroxenes. Depending on the Al content, minor

amounts of kaolinite may be formed. On ultrabasic rocks, saponite, serpentine,

chlorite, and talc are the most frequent phy llosilicates derived from pyroxenes and

14.1. Geological Environments for Clay Formation 1133

olivine. Nevertheless, under a tropical climate, kaolinite can be the major clay min-

eral formed from basalt (Siefferman and Millot, 1969), anorthosites (Gomes et al.,

1994), or gabbros (Gala

n et al., 1996) by hydrothermal or weathering alteration.

The weathering of quartz–chlorite schists can produce a complex sequence of clay

minerals starting from chlorite and ﬁnishing with kaolinite. In the host rock of

uranium ore deposit at Koongarra, Australia, Murakami et al. (1996) observed the

following weathering sequence: (i) chlorite; (ii) chlorite–vermiculite intergrade; (iii)

interstratiﬁed chlorite and vermiculite; (iv) vermiculite; and (v) kaolinite.

A specia l case of weathering takes place when eruptive rocks react with sea water

in a deep ocean environment. Basalt, the most frequent rocks found in ocean ﬂoors,

are hydrolyzed to form clays on the ﬂoor surface. These clays contain much Fe, and

become progressively enriched in alkaline elements at lower temperatures. Clay

mineral formation begins below 30 1C, with saponite forming at high, and Fe-rich

beidellite and celadonites at intermediate temperatures.

The genesis of the various clay minerals found in speciﬁc weathering proﬁles

remains to be resolved in detail. The multitude of factors involved in this environ-

ment are responsible for many variations in weathering sequence an d the crystallo-

chemical properties of the intermediate and end member products. However, some

generalization is possible.

Illite and other micaceous minerals are inherited from parent rocks or other

materials, where they form under different P–T conditions from those existing at the

earth’s surface (Fanning et al., 1989). The formation of illite by a pedogenic process

is possible but it is very rare (Wilson, 1999). The formation of illite from micaceous

minerals involves cleavage of the relatively thick, micro-size mica particles with a full

content of K into very thin particles depleted in K (illite) (Roberts et al., 1991;

Romero et al., 1992). The majority of micaceous soil clays are dioctahedral and

aluminous, although some rare trioctahedral illites, derived from biotite, do exist.

Smectite formation in soils takes place when the following factors come together:

base-rich parent rocks, poor drainage, low-lying topography, high pH , high silica

activity, and abundance of basic cations (Table 14.1)(Keller, 1970; Borchardt, 1989).

These conditions can occur under very different climates (temperate, cold, or even

tropical) when leaching is limited. Soil smectites consist principally of dioctahedral

montmorillonite and beidellite that are somewhat enriched in Fe as compared with

their ideal compositions (Wilson, 1987a). They are also somewhat different from the

smectites in ben tonite deposit s. This might be because the chemical variability is

higher in the weathering environment than during bentonization since the smectites

in most soils can be inh erited, transformed, and ne oformed (Table 14.2). Smectites of

different composition may crystallize locally in microenvironments from different

parent materials (Aoudjit et al., 1995). The illite-to-smectite transformation seems to

be a simple degradation pr ocess but illitization of smectite (‘aggradation’) in wea-

thering environments is not fully understood.

Vermiculite is very common in soils and weathering proﬁles developed on acidic

and basic rocks. Vermiculite seems to be formed by the weathering of micas, mainly

Chapter 14: Genesis of Clay Minerals1134

Table 14.1. Geological environment of primary-stage montmorillonitization (Keller, 1970)

A. Retention of Mg

,Ca

,Fe

,Na

Evaporation exceeds precipitation, as semi-arid climate

Ineffective leaching

Stagnant water and water logging

Ash in lakes and marine basins

Low effective permeability of rocks

Alkalinity

not combined with O

2–

or S

2–

Silicates characterized by

High speciﬁc surface, as for volcanic ash

High susceptibility to hydrolysis

B. Retention of silica

Flocculated by Ca

,Mg

, and other cations

Rock rich in Ca

,Mg

, and Fe

Ineffective leaching

Clay-size cristobalite

C. Retention of parent texture and mineralogy

Shards or ﬂow

Igneous-type minerals

Table 14.2. Abundance of smectites in soils as assessed from the literature (Wilson, 1999)

Soils Inheritance Neoformation Transformation

Entisols +++ +++ +

Aridisols +++ +++ 

Inceptisols +++ +++ +++

Vertisols +++ +++ +

Mollisols + ++ ++

Alﬁsols + +++ ++

Spodosols  + +++

Ultisols +++ + 

Oxisols  

+++ major importance, ++ moderate importance, + minor importance,  no importance.

14.1. Geological Environments for Clay Formation 1135

from biotite (trioctahedral vermiculite), involving the release of interlayer K

, and

the oxidation of structural Fe

(Vicente-Herna

ndez et al., 1983). Soil vermiculite

may also derive from muscovite (dioctahedral vermiculite) but this transformation

was not well understood until the application of high-resolution transmission elec-

tron microscopy (HRTEM). Aoudjit et al. (1996) presented HTREM images of the

vermiculitization of muscovite in an acidic soil, showing particles with a muscovite

core and a vermiculite rim consisting of about 15 layers. Very ﬁne-grained muscovite

particles of high speciﬁc surface area and reactivity seem necessary for transforma-

tion into vermiculite in soil. The degradation of chlorite by decomposition of the

interlayer hydroxide sheet is another means of obtaining vermiculite (Makumbi and

Herbillon, 1972; Ross and Kodama, 1976).

Chlorite is an inherited clay mineral in soils and weathering crusts. This layer

silicate is easily weatherable and transformable into regularly interstratiﬁed ch lo-

rite–vermiculite and vermiculite (Proust, 1982; Buurman et al., 1988). Chlorite can

also be formed pedogenically by the intercalation of Al (hydr)oxides into pre-existing

smectite, vermiculite, or interstratiﬁed expansible minerals. The source of the inter-

layer Al is the degradation of both the tetrahedral and octahedral sheets of layer

silicates, weathering of Al-minerals (i.e., feldspars), or decomposition of organic

matter containing adsorbed Al. Like smectite and illite, chlorite is very sensitive to

acidiﬁcation. In soils affected by ‘acid rain’ or acidiﬁed surface water, trioctahedral

chlorite is destroyed in the E horizons, illite transforms into vermiculite via regularly

interstratiﬁed illite–vermiculite, interlayer hydroxy-Al sheets are liberated (particu-

larly in B horizons), and poorly ordered material (probably allophane) is preci-

pitated in B horizons (Wilson, 1987b). Gala

n et al. (1999) shown that acid mine

drainage is responsible for the progres sive degradation of chlorite in the basin of the

Tinto River (Spain).

Kaolinite can be a neoformed, transformed , or inherited mineral in soils. As

already menti oned, high rainfall and a temperate climate can transform muscovite

and biotite into kaolinite together with some illite. Al-(hydr)oxides liberated by

extensive leaching can react with silica to produce kaolinite under slightly acidic

conditions (pH E5), and when silica activity is moderate and the concentration of

basic cations is low. Keller (1970) summarize d the general conditions for the weath-

ering formation of kaolinite (Table 14.3). Like kaolinite that derives from weath-

ering, soil kaolinite is usually highly disordered and may contain Fe

from

isomorphous substitution. In some cases, it is also interstratiﬁed with smectites

(Linares and Huertas, 1971; La Iglesia and Van Oosterwyck-Gastuche, 1978;

Herbillon et al., 1981 ; Dixon, 1989).

Halloysite occurs frequently as a neoformed mineral in weathering crusts devel-

oped on acidic volcanic rocks (Siefferman and Millot, 1969). Under supergene con-

ditions halloysite can transform into kaolinite according to the sequence: disordered

halloysite-ordered halloysite-metahalloy site-disordered kaolinite-ordered ka-

olinite (Ponder and Keller, 1960; Chem, 1969). As La Iglesia and Gala

n (1975)

shown, this process is one of recrystallization, requiring the presence of organic

Chapter 14: Genesis of Clay Minerals1136

acids. Kaolinite can also transform into halloysite (Singh and Mackinnon, 1996;

Bobos et al., 2001).

Palygorskite and sepiolite, the ﬁbrous clay minerals in soils and paleosols, are

generally neoformed (Singer, 1979). The instability of these minerals in wet climates

favours their presence in soils under dry or semi-dry climates. The occurrence of

palygorskite in soils is associated with one of the following conditions: (i) in modern

soils that, at present or in the past, were affected by rising ground water of pH 7–8

and abundant salinity; (ii) in soils with distinct and sharp textural transitions because

these minerals accumulate in the coarse fraction (this group includes many paleo-

sols); and (iii) in calcretes, caliches (Singer, 1984). In all these cases, palygors-

kite (and very rarely sepiolite, when Al is absent or immobilized (Jones and Weir,

1983)) is precipitated by evaporation of the vadose water. Cemented soils (duri-

crusts) can be caused by palygorskite (palycretes), or by silica (silcretes) and calcite

(calcretes) (Rodas et al., 1994). The transformation of smectite into palygorskite was

described in terms of a dissolution–precipitation process (Jones and Gala

n, 1988), or

Table 14.3. Geological environment for primary kaolinitization (Keller, 1970)

A. Removal of Ca

,Mg

,Fe

,Na

Precipitation exceeds evaporation

Permeable rocks

Percolating water

Oxidation of Fe

to Fe

or FeS

B. Addition of H

Fresh water

Acids

Sulphur compounds

Carbonic, air and soil atmosphere

Organic, living and dead organisms

C. High Al:Si ratio

Removal of silica

Na and K silicates

Organic complexes

High concentration of Al

in acid solution

Al–OH polymers

14.1. Geological Environments for Clay Formation 1137

transformation through kerolite–smectite mixed-layer minerals (Eberl et al., 1982).

The transformation of smectite and illite to palygorskite was veriﬁed by Suarez et al.

(1995), Lo

pez Galindo et al. (1996), and Sa

nchez and Gala

n (1995). Palygorskite is

more easily formed in weathering environments than sepiolite, or else sepiolite is less

stable under supergene conditions, and hence is rarer. Traces of palygorskite found

in some arid soils are considered as inherited (Shadfan and Dixon, 1 984; Mackenzie

et al., 1984).

In summary, we can say that under weathering conditions most common rock-

forming silicates (except quartz) are destroyed, and practically all clay minerals can

form. Illite derives from micaceous minerals and feldspars, chlorite can be inherited

or formed from biotite. Kaolinite can be neoformed, or derived from muscovite and

feldspars, and more rarely from smectite or vermiculite. Vermiculite comes from

muscovite/biotite and chlorite through interstratiﬁed structures. Smectite can form

directly by alteration of basic silicates (pyroxenes, amphiboles), or from illite and/or

vermiculite. Palygorskite can be directly precipitated under certain climatic condi-

tions, or formed from smectite and illite. Sepiolite seldom forms in soils. The com-

plexity of the systems is very large and not all the possibilities were yet demonstrated.

Some of the most frequent pathways are illustrated in Fig. 14.2.

14.1.2. Sedimentation

Ancient sedimentary rocks are composed of about 70% mudstones (containing

about 50% clay-size particles) and shales (Blatt et al., 1980). Presently, sedimentary

environments that contain muds cover about 60% of marine continental shelves and

40% of deep ocean basins (Hillier, 1995). Continental aquatic environments such as

lakes, rivers, estuaries, and deltas also contain high proportions of clays. The

amounts and varieties of clay minerals found in sediments and sedimentary rocks are

a direct result of source area differences, changes occurring during transport and

deposition, and changes taking place after sedimentation (diagenesis).

The erosion of soils is governed by the proﬁle structure, the humidity, the clay

content, the slope, and the climatic conditions. A well-developed soil is more sen-

sitive to erosion, because it shows a thick upper zone where clays predominate and

which can be easily removed. The ﬁnest particles of the proﬁle (mostly clay minerals)

are most easily transportable. On the other hand, clays can contain signiﬁcant

amounts of adsorbed water, providing protection against ero sion by sudden abun-

dant rainfall, which is the principal erosion factor.

Clay transportation mainly occurs by water ﬂow (stream, river), but the timespan

of this transport is very short and has practically no inﬂuence on clay alteration. The

deposition zone spreads from the mouth of rivers to lagoons, lakes, seas, and oceans,

producing a particle-size fractionation from the mouth to the interior, together with

different mineral settings. According to Parham (1966), the order of the ﬁrst ap-

pearance of major clay mineral groups from shoreward to basinward areas is as

follows: kaolinite, illite, chlorite, palygorskite, and sepiolite. Montmorillonite and

Chapter 14: Genesis of Clay Minerals1138

I/S mixed layered minerals are present over the entire range. This approach con-

tinues to be true in the most cases, except for the ﬁbrous clay minerals (specially

sepiolite). On the other hand, salinity changes can also produce coagulation proc-

esses affecting some clay minerals. A typical clay mineral distribution found from the

coastline to the open sea is: kaolinite-illite- smectite.

After deposition the inter stitial water solution in the sedim ent has a higher ionic

concentration than that in the weathering proﬁle, leading to a change in chemical

equilibrium. The chemical potential of the elements in solution is different from that

for clay formation in the continent, and clays react with the environment solution.

The time of connection between the sediment pore water and water reservoir can be

very long or very short depending on sediment accumulation rates. Typical reactions

are clay–clay, and clay-(hydr)oxide transformations with chemical contribution from

the sedimentary solution. The new phases formed depend on the redox potenti al of

the ambient solution which is partly controlled by organic matter, and partly by

living organ isms at the water/sediment interface. For example, microbial action can

change the oxidation state of Fe, or stabilize sulphides, and hence control the

Transformation

Vermiculite

Smectite

Palygorskite

Smectite

(Kaolinite)

V-S

B-V

Inheritance

Chlorite (**) Smectite (***)

Ch-V

Ch-S

Ke-S

B-S

Neoformation Kaolinite Smectite

Transformation

Vermiculite Smectite

Smectite

Kaolinite

(Palygorskite)

AlV-S

Inheritance

I-S

I-V

Acidic

Rocks

Basic

Rocks

(Kaolinite)

Biotite

Palygorskite

(Sepiolite)

(Kaolinite)

Smectite/

Chlorite/

Vermiculite

Talc

Ca-Plag/Amp/Py

Illite + Kaolinite

Feldspars

Muscovite (Biotite)

Illite*

Fig. 14.2. Pathways for the formation of clay minerals in weathering environment from

different starting materials. Minerals in bold type indicate the most probable end products. In

brackets are minerals that do not form easily according to that pathway, or are rare. Clay

minerals derived from other rock-forming silicates could be formed by transformation and/or

neoformation. (



) Illite can be considered inherited in soils, and as the ﬁrst degradation

product of muscovite, (



) Chlorite may be inherited or transformed from biotite, (



)

Smectite may be inherited in soils. Also some palygorskite may be inherited in soils.

14.1. Geological Environments for Clay Formation 1139

availability of certain elements that clays can ﬁx. This early post-deposition changes

are usually termed ‘early diagenesis’.

In general, clay minerals of sedimentary sequences (in areas not subject to tectonic

activity) mainly reﬂects the climate, relief, and lithology of source areas. However, in

zones of intense tectonic activity (ﬂysch) the clay mineral assemblages in sediments

can reﬂect the composition of the parent rocks because tect onic activity prevents the

formation of soils in the continent. The rapid transport of mate rial by turbidity

currents with high solid concentrations does not produce well-sorted sediments;

rather, the sediments usually have a grain-size distribution similar to that of the

parent suspensions (Ruiz Cruz, 1999). The mineralogy of recent sediments in coastal

shelves and estuaries is generally consistent with terrigenous origin, that is, from soils

and weathered rocks of the hinterland (Chamley, 1989 ). Sometimes a particular

mineral can be an indicator of continental origin. For example, sediments from the

Galician Coast of Spain contain signiﬁcant amounts of gibbsite, a common con-

stituent of soils found in weathering proﬁles of this region (Belzunce-Segarra et al.,

2002).

In continental-shelf and deep-sea sediments the bulk of the clay fraction is detrital

(Kastner, 1981; Chamley, 1993) but some of the smectites, I/S mixed-layer minerals,

glauconite, and berthierine are neoformed by alteration of volcanic material, and by

low-temperature interaction between X-ray amorphous hydr(oxides) and biogenic

silica.

Certain clay minerals formed in marine environments can serve as paleoenviron-

mental indicators. For example, both glauconite and berthierine are typical marine

clay minerals formed from pre-existing clays (kaolinite, smectite, and iron oxides, at

temperature around 4 1C) by diffusion of ions in the solid state and then recrys-

tallization. Berthierine does not incorporate K

as does glauconite, but takes up

large amounts of Fe

and Fe

. The geological environment of both minerals

seems to be different. Glauconite occurs in a calmer ambient situation (shallow

shores of oceanic platforms) than berthierine.

Glauconization is analogous to illitization. It involves a gradual alteration of

neoformed Fe-rich smectite into Fe-illite, similar to celadonite, through intermediate

mixed-layer structures. Glauconitization is complete when the layers no longer ex-

pand and the K

O content increases to 9%, which requires about 1 Ma (Odin, 1988).

Transformations of detrital clay minerals at the bottom of seas and oceans include

the formation of Al-smectites, showing that these minerals can be stable in this

environment. In addition, devitriﬁcation of volcanic materials in the deep ocean

environment yields new crystalline phases, such as smectites and zeolites.

Palygorskite of detrital origin can also be found in oceans, having come from the

continent, as in the case of deep-sea Atlantic palygorskite close to Morocco, which

was transported by SW winds from near-shore shallow African basins into the ocean

(Pletsch et al., 1996). Palygorskite can also be derived from smectite by diagenetic

transformation in the marine environment (Lo

pez Galindo, 1986), sometimes

associated with marine phosphorites (Chahi et al., 1999). Sepiolite can also be of

Chapter 14: Genesis of Clay Minerals1140

diagenetic origin through the transformation of magnesite at pH 10.5–11.5 in silica-

rich lake waters (Ece, 1998). However, most of the palygorskite and sepiolite in the

world are neoformed (precipitated) in lacustrine and peri-marine environments. Al-

though these minerals can be derived from other clay minerals, the most common

mechanism is crystallization from solution, promoted by the presence of amorphous

silica. Experiments by Birsoy (2002) indicated that low aqueous Al-activities favour

sepiolite and kerolite relative to palygorskite and saponite.

In some shallow restricted basins with evaporitic conditions ﬁbrous clay minerals

and Mg

-rich smectites (saponite, stevensite, kerolite) can form authigenic min-

erals, usually together with sulphate and carbonate minerals. In this environment

smectite can also be transformed into palygorskite (Sa

nchez and Gala

n, 1995). In a

typical playa-lake sequence, from the basin edge to the centre, detrital miner als

(illite, ka olinite, smectite) dominate at the lake edge, palygorskite with some

-smectite near the shore, while sepiolite and other Mg

-clay minerals abound

towards the centre.

In summary, we can say that sedimentary basins are repositories of clay miner als,

forming part of the clay cycle (weathering-deposit ion-diagenesis/metamor-

phism-magmagenesis). The clay minerals can be formed by neoformation (i.e.,

direct precipitation from solution or through ageing of amorphous materials, such as

glassy volcanic ash), inheritance (i.e., transported from the source to the basin with

little modiﬁcation), and transformation (i.e., when neoformed and inherited clays in

the basin react by diagenesis and low-grade metamorphism). Basins formed at active

plate tectonic margi ns tend to have a greater diversity of clay origins than at passive

margins or within-plate basins (Fig. 14.3).

14.1.3. Burial Diagenesis and Low-Grade Metamorphism

When sedimentary clays are buried, free or weakly bonded water is eliminated. As a

result, the porosity decreases from 80% to about 20% at the ﬁrst kilometer of burial.

As depth increases, sediments tend to reach a global chemical equilibrium and new

phases can form under the new environmental conditi ons. The unstable phases tend

to disappear and the number of species diminishes. The reactions are mostly of the

clay–clay and clay–solution types with time and temperature being the principal

factors governing the transformations. The new minerals are more highly ordered

and of large size than their original minerals.

After the disappearance of most inherited soil clay minerals, new clay minerals are

formed: I/S- and chlorite–smectite mixed-layer minerals, chlorite, kaolinite together

with zeolite, some quartz, and feldspars. In general, illite is the most common dia-

genetic product, and I/S mixed-layer minerals are the main indicator of diagenetic

evolution of a sedimentary basin (Dudek and S

rodon

, 2003).

The last stage of diagenesis is the beginning of metamorphism (anchimetamor-

phism) when most of the interstratiﬁed clay minerals and kaolinite are destroyed,

resulting in an illite–chlorite facies with pyrophyllite, biotite, and paragonite.

14.1. Geological Environments for Clay Formation 1141

Typical low-grade metamorphic rocks derived from the clay sediments are pelitic

rocks (slates), mainly composed of mica (or illite), chlorite, alkali feldspar and

quartz, with some pyrophyllite, and/or paragonite. Shales are gradually transformed

into phyllites, and with increasing metamorphism, into schists. Phyllosilicate par-

ticles become too large to be considered clay minerals, although they can be similar

crystallochemically.

Perhaps the most signiﬁcant mineral change on burial diagenesis is the gradual

loss of expandable miner als, mainly the transformation of Al

-smectite into illite

through I/S mixed-layered structures. The change occurs step-by-step. The ﬁrst one

is the random interstratiﬁcation of smectite with illite (Reichweite, R ¼ 0), and

composition ranging from pure to 50% smectite. The 50% I/S mixed-layer structure

is a regular interstratiﬁcation of good stability. The last stage is the formation of illite

Transformed

Clays

Fore Arc Basin

Transformed

Clays

Inherited

Clays

Neoformed (volcanogenic)

and inherited clays

Inherited

Clays

Forelan

Basin

Fold-and-thrust

Belt

Remnant

Arc

Back-Arc

Basin

Active

Arc

Accretionary

Complex

Neoformed

Clays

Inherited + Neoformed

clays

Neoformed + Inherited

clays

Mid-ocean ridge

and ab

ssal

lain

Passive

Margin

Continental

Rift Basin

Continental

Sag Basin

Fig. 14.3. Characteristic modes of origin of clays in sedimentary basins at an active plate

margin (above), and a passive margin and within-plate setting (below). From Merriman

(2003).

Chapter 14: Genesis of Clay Minerals1142