Bergaya F. Handbook of Clay Science

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correlated polarizations of the ionic clouds give rise to an ionic van der Waals type

force, much in the same way as correlations between ﬂuctuating electronic dipoles

give rise to the London dispersion force. These attractive forces are not taken into

account in the DLVO theory of colloid stability.

Ionic correlation forces increase as the surface charge density of the layers, and

the charge of the interlayer ions, increase. Using Monte Carlo simulations, Pellenq

and co-workers (Pellenq et al., 1997a, 1997b; Delville and Pellenq, 2000) carried out

a systematic study of correlation forces in water between surfaces with a broad range

of charge densities, within the framework of the so-called ‘primitive model’ (smooth

surfaces separated by ions in a dielectric continuum). As expected, the attractive ion

correlation forces are particularly strong when the CSH layers are fully ionized and

separated by hydrated and mobile calcium ions. Attractive pressures as strong as

60 atm are obtained at an interlayer distance of 0.7 nm.

A different kind of attractive electrostatic interaction may appear when the ionic

mobility vanishes. Just as the distribution of cations and anions in a NaCl crystal

gives rise to a strong Madelung cohesive energy, the distribution of adsorbed,

translationally immobile, charge-balancing ions in a narrow interlayer space may

lead to a strong cohesive energy. A covalent contribution may also appear at small

−

Fig. 13.3.4. Two uniformly charged walls separated by a dielectric continuum in which hy-

drated ions are free to move. In the overwhelming majority of instantaneous conﬁgurations,

the distribution is neither symmetrical with respect to the mid-plane nor homogeneous along

the vertical axis. An excess density of ions in some place leads to a deficit in other places. This

generates strongly attractive electrostatic forces that may overcome the osmotic repulsion and

are not taken into account in traditional double-layer theories.

Chapter 13.3: Cement Hydrates1122

counterion-surface distances. A correct modelling of this requires atomic level sim-

ulations, using either potential energy minimization or direct quantum methods.

The case of tobermorite-like CSH was recently considered (Gmira, 2003; Gmira

et al., 2004). The ab initio quantum approach indicates that the attractive contri -

bution of electronic dispersion interactions (van der Waals forces) is about 20% of

the cohesion energy. Since the partial electronic charge on the atoms is independent

of interlayer distance, the nature of the bonds between and within the layers remains

unchanged. The ionic character of the bonds in the layers is close to 60%. The charge

on the interlayer calcium ions is +1.38, which is slightly higher than for ions within

the layers (+1.29), but still much lower than the value expected for a purely ionic

bond. If the ions were totally hydrated and free to diffuse in an electric double layer,

they would have a charge of +2. Furthermore, the distance between the interlayer

calcium ions and the oxygen atoms of the layers is very close to the Ca–O distance

within the layers (Fig. 13.3.5). This indicates that the interlayer calciu m ions are

linked to the layers by a strong ionic bond, with a significant covalent character.

The potential energy (at 0 K), the free energy (at 300 K) and the derived interlayer

pressure of the structure were also calculated as a function of the interlayer distan ce.

A strongly cohesive behaviour is obtained with attractive pressures even higher than

in the primitive model.

13.3.6. SUMMARY AND CONCLUSIONS

The results summarized here show that there is probably a conceptual continuity

between the limited swelling of Ca-smectites and the strongly cohesive behaviour of

hardened cement with particular reference to calcium silicate hydrate.

For cohesive forces at very short length scales, the hydration state of the interlayer

cations and the possibl e speciﬁc interaction of these cations with surface sites are the

key parameters. In equilibrium with humid air, the calcium ions in the interlayer

space of smectites have a complete hydration shell and do not interact specifically

with some surface sites of the clay mineral. In CSH, on the other hand, the interlayer

calcium ions are partially dehydrated and interact specifically with Si–O

–

groups of

the short silica chains. This confers a partially covalent character to the bonding

between the surface and the interlayer cation. In addition, it allows for a much

stronger ‘surface–cation–surface’ ionic–covalent interaction.

As far as cohesive forces at somewhat larger (but still sub-nanometer) length

scales are concerned, the ion correlation forces are the best candidates, both in clay

minerals and in cement. However, they require mobile (i.e., exchangeable) cations.

Calcium may be involved in partially covalent bonds. Even in CaO the bonding scheme is not 100%

ionic. Pauling’s rule predicts that the Ca–O bond has 21% covalent character, which is not far from our

estimate (30%). The electron density, calculated by Gmira, is too high for a purely ionic bond because of

the inner-sphere character of the conﬁguration. All inner-sphere complexes have a partially covalent

character.

13.3.6. Summary and Conclusions 1123

This requirement is clearly met by smectites but not by CSH. To be significant, ion

correlation forces also require a high surface charge density. From that point of view,

CSH is a priori favoured with respect to smectites. However, the extensive screening

of surface charge by the specifically adsorbed cations may well reduce the effective

surface charge of CSH to values that are not much larger than those of smectites.

A surprising feature is that, in spite of strong differences in cohesion, especially

as far as short range cohesive forces are concerned, smect ites and synthetic CSH

(Fig. 13.3.2) form particles not vastly different from each other in thickness (typi cally

from a few layers to a few tens). The lack of explanation for this observation in terms

of surface forces (Sposit o, 1992) shows that we are still largely disarmed by some-

thing as widespread and as apparently simple as the self-associating properties of

layered miner als in water.

The relationship between local order and local cohesion, on the one hand, and

macroscopic strength, on the other, is an open question. The lateral extension of

Interlayer

space

Interlayer

space

C-S-H

layer

Fig. 13.3.5. A relaxed conﬁguration of tobermorite-like CSH (Hamid’s structure) after po-

tential energy minimization, at Ca=Si ¼ 0:83 and four water molecules per unit cell in the

interlayer space. The double plane of calcium ions in the central CSH layer is indicated by

horizontal arrows. The hydrogen atoms of the water molecules, on each side of this layer, are

in white. The interlayer calcium ions are in light grey and indicated by tilted arrows. Note

the short distance between the interlayer calcium ions and the closest oxygen atom from the

neighbouring layer, showing that the calcium ions form an inner sphere complex with the

surface.

Chapter 13.3: Cement Hydrates1124

particles is definitely much larger in natural smectites than in the CSH of hardened

cement. The connectivity between locally ordered regions is also much better in

smectites. This is due to the large size and ﬂexibility of individual smectite layers,

allowing a single layer to participate in several particles. Yet, in spite of poorer order

and poorer single layer connectivity, hardened cement has a tensile and compressive

strength higher than that of clay minerals (at compara ble porosity). What this ob-

servation suggests is that a quantitative description of defects is probably as im-

portant as the knowledge of order. This is not only valid for the comparison between

cement and smectites, but also applies to comparing smectites exchanged with dif-

ferent cations (Zabat et al., 1997).

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Kjellander, R., Marcelja, S., Quirk, J.P., 1988b. Attractive double-layer interactions between

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Handbook of Clay Science

Edited by F. Bergaya, B.K.G. Theng and G. Lagaly

Developments in Clay Science, Vol. 1

r 2006 Published by Elsevier Ltd.

1129

Chapter 14

GENESIS OF CLAY MINERALS

E. GALA

Departamento de Cristalografı

a, Mineralogı

a y Quı

mica Agrı

cola, Facultad de

Quı

mica, Universidad de Sevilla, ES-41071 Sevilla, Spain

The origin of clay minerals is an important topic for applied clay scientists to

consider. Clay minerals formed in diverse environments unde r varying temperature

and pressure conditions may exhibit different properties. The crystallochemical

make up, the physical and physico-chemical properties, and the technical behaviour

of a kaolin formed in one environment, for example, may be considerably different

from a kaolin formed in another. The special properties of clays are origin depend-

ent. Knowledge of clay genesis can also contribute to the exploration and selective

mining of the clay materials most appropriate for speciﬁc industrial requirements.

Clays occur in many different geological environments: weathering crusts and

soils, continental and marine sediments, volcanic deposits, geothermal ﬁelds, wall-

rock alteration produced by intrusion of plutonic rocks and hydrothermal ﬂuids, and

very low-grade metamorphic rocks. Clay minerals mostly form from pre-existing

minerals, primarily from rock-forming silicates by transformation, and/or neofor-

mation, where rocks are in contact with water, air, or steam.

As Kon ta (1992) stated ‘everything starts in weathering crust’. While many com-

mon igneous an d metamorphic rock-forming silicates are destroyed by weathering

processes, phyllosilicates, quartz, and some heavy minerals are highly resistant, and

in this environment the crystallization and growth of clay minerals are favourable.

Erosion, transportation, and deposition of these weathering crusts in lakes, seas, and

oceans, lead to sedimentary sequences where clay-sized material, composed mainly

of phyllosilicates and Fe-hydr(oxides), makes up about 50% of the volume of sedi-

mentary rocks (Taylor, 1952). The phyllosilicates are mostly composed of clastic

(inherited) muscovite, biotite, and chlorite, with small amounts of kaolinite,

smectites, and vermiculite produced in soils and weathering crusts, or after depo-

sition in the new environment.

The clay minerals formed in the weathering crust should be considered stable

under supergene conditions. In many cases, however, they are metastable needing

more time to reach equilibrium with ambient conditions; in other cases, their trans-

formation into stable forms is not thermodynamically and/or kinetically possible.

DOI: 10.1016/S1572-4352(05)01042-1

If these new clay minerals can be reasonably considered stable in the supergene

environment, and erosion, transportation, and post-depositional processes have not

basically modiﬁed them, then sedimentary records can be used for paleoclimatic

interpretations (Millot, 1964; Singer, 1979–1980, 1984; Gala

n, 1986; Lo

pez Aguayo,

1990).

However, transportation, deposition, burial, and diagenesis of the weathered

materials usually ch ange the physico-chemical properties of clay minerals, and even

inﬂuence their transformation into other clay minerals.

Clay formation and evolution are distinctive events in the geological history of a

basin. Source area, transport, depositional environment of clays, and the transfor-

mation and diagenetic changes are fundamental for analysis of sedimentary basins

(Chamley, 1989; Merriman, 2002, 2003). Clay minerals are particularly sensitive to

pressure and temperature variations and to the chemical environment. This sensivity

is expressed in terms of their chemistry and mineralogy. The study of clay minerals

can therefore provide information about the processes that occur during the evo-

lution of the seri es of which they form a part. Following Millot (1964) clay minerals

in a sedimentary basin can be neoformed (i.e., by direct precipitation or via the

ageing of X-ray amorphous materials), inherited from the source area and trans-

ported with little modiﬁcation, or transformed (i.e., derived from inherited clay

minerals in response to environmental reactions).

During transformation, the essential silicate structure of the clay mineral is largely

maintained although the interlayer space can be greatly mod iﬁed due to changing

composition of the tetrahedral and octahedral sheets. Millot (1964) distinguished

between ‘degradation’ and ‘aggradation’ in the transformation process. The con-

version of illite to smectite (illite-vermiculite-smectite) is an exampl e of degra-

dation since it involves a depletion of K and other elements, while the reverse

transformation is one of aggradation because it involves the addition of various

elements. Degradation is characteristic of rocks and soils weathering, while aggra-

dation is very rare in such environments although it probably occurs in burial dia-

genesis.

Hydrothermal alterati on of some volcanic and plutonic rocks can also give rise to

massive accumulation of clays. The intrusion of plutonic rocks, on the other hand,

leads to wall-rock alteration and transformation of some rock-formi ng minerals into

clayey material, along with the neoformation of many other minerals, including

phyllosilicates. As the transformed and neoformed minerals occur sequentially from

the contact towards the coldest part of the host rock in different assemblages, they

can be used as geological thermometers. Their geochemistry can be a good indicator

of the composition of the intrusive material and associated ores.

In summary, clays can be used in interpreting and understanding geological

processes, such as environmental conditions of sedimentary facies, zonation, stra-

tigraphic correlations, source area of sediments, weathering and hydrothermal al-

terations, diagenesis, and low-grade metamorphism (i.e., Weaver, 1960; Keller, 1970;

Singer, 1979–1980; Gala

n, 1982, 1986; Henley, 1985; Ortega-Huertas et al., 1991;

Chapter 14: Genesis of Clay Minerals1130

Konta, 1992; Chamley, 1993; Essene and Peacor, 1995; Hillier, 1995; Pletsch et al.,

1996; Gutierrez -Mas et al., 1997; S

rodon

, 1999; Carretero et al., 2002; Merriman,

2002).

14.1. GEOLOGICAL ENVIRONMENTS FOR CLAY FORMATION

14.1.1. Weathering

The weathering environment is usually subaerial. It involves physical disaggregation

and chemical decomposition, leading to the transformation of original minerals into

clay minerals. The factors controlling rock weathering include: rock type, climate

(rainfall: a chemical factor, and temperature: a physical factor), time (the age of the

weathering proﬁle), topography (affecting the ratio of water to rock through drain-

age), and the presence of organisms and organic matter. The chemical composition

of the weathered rock and the water/rock ratio are the principal factors determining

the type of clay minerals formed (Fig. 14.1), while temperature and time inﬂuence the

rate of chemical processes (Velde, 1992; Foley, 1999).

Weathering of non-sedimentary rocks commonly produces either of two c lay

minerals: kaolinite and dioctahedral smectite. These species are thermodynamically

stable with respect to the chemical composition of near-surface pore waters. The

chemical composition of this water is determined by the interplay between rainfall,

drainage rate, chemical composition, and dissolution rate of the parent rock. These

clay minerals and the degradation products of the micaceous rock-forming minerals

(i.e., illite) constitute the bulk of the weathering products by volume (Velde, 1992).

On a global scale the clay distribution depends on the prevailing climatic con-

ditions, because these determine the weathering intensity (Millot, 1979). In cold

(polar) climates illite and chlorite are favoured, while in mild climates vermiculite

and mixed-layer structures are the most frequently formed clay minerals. In tropical

zones and Mediterranean climates with seasonal contrast, smectites are favoured,

and ﬁnally under wet tropical and equatorial climatic conditions kaolinite and Al–Fe

(hydr)oxides are the main clay mineral components of soils. This simple an d at-

tractive zonation scheme based on the present-day clay mineral distribution is not

true in detail, because the other factors cited above have an inﬂuence that often

disturb this generalized distribution of clay minerals.

In fact, stable isotope studies on kaolinites from weathering proﬁles in Australia,

India, and South America shown that kaolinite is formed under cold climates.

Chivas and Bird (1995) suggested that Permian kaolins of Australia are formed from

glacial melt waters. Thus, the presence of a kaolinite weathering crust by itself can no

longer be interpreted as being indicative of tropical weathering.

Velde (2001) shown that the clay minerals formed from different starting mineral

assemblages (ﬁne-grained phyllosilicates) in the uppermost horizons of agri-

cultural soils, under similar geomorphologic and climatic conditions, consist of two

14.1. Geological Environments for Clay Formation 1131

Fig. 14.1. The effect of mean annual precipitation (1 inch ¼ 25:4 mm) on the frequency dis-

tribution of clay minerals and gibbsite in the surface layer of residuals soils from acid (A) and

basic (B) igneous rocks in California. After Barshad (1966).

Chapter 14: Genesis of Clay Minerals1132