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Industrial Clays

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 AUTHOR QUERY FORM Book: 1974

Chapter title: The geological setting for industrial mineral resources

Author: Eric Pirard
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 EMU Notes in Mineralogy, Vol. 9 (2011), Chapter 9, 341– 414

Industrial clays

GEORGE E. CHRISTIDIS
Technical University of Crete, Department of Mineral Resources
Engineering, 73100 Chania, Greece
e-mail: christid@mred.tuc.gr

Clays have been used by man since prehistoric times. Initially they were used almost
entirely in the fabrication of ceramics, nowadays they find numerous industrial and
technological applications including the production of materials with large added
value such as nanocomposites, cosmetics or pharmaceuticals. The term clay should
not be considered as a synonym for clay mineral, because clays consist of more than
one mineral. The versatile nature of clays is attributed to the presence of clay minerals,
which impart significant physical properties to the raw materials, such as particle size and
shape, ion exchange, hydration and swelling, plasticity, rheological properties, colour
properties and reactions with organic and inorganic compounds. Four types of industrial
clay raw materials are examined in this contribution, kaolins, bentonites, fibrous clays
(palygorskite and sepiolite) and common clays and shales. The latter are used in the
production of structural ceramics, bricks tiles and pipes. The industrial clay deposits
are classified as primary, residual formed from in situ alteration of various precursors
or hydrothermal) and secondary, formed from deposition of clastic clay materials
which were transported from their sources. Assessment of industrial clay deposits
comprises determination of physical properties and direct comparison with international
or regional standards, which include industrial specifications for particular applications.
These specifications are often dictated by the end industrial users. Exploitation of the
clay deposits is usually by means of traditional open-cast methods and processing
can involve anything from simple crushing, screening and tempering, to elaborate
mineral beneficiation techniques such as alkali or acid activation, delamination,
magnetic separation, selective flocculation, flotation and leaching. The method used
and the extent of beneficiation are dictated by the final industrial application of the clay.

1. Introduction – terminology for clays and clay minerals

Clays have been used by man as raw materials since the Upper Palaeolithic era. The use
of clays in pottery ceramics was the most widespread application of clays for mankind in
ancient times and the oldest ceramic article found is the Dolni Vestonice Venus, a female
figurine dated at 26– 30 ka (Vandiver et al., 1989). The oldest pottery ceramic articles
found so far, which have also been dated as Upper Palaeolithic (12– 14 ka), were found
in Japan and belong to the Jomon Period (Kaner, 2003). Since then, the use of clays in
the fabrication of pottery ceramics has increased significantly and additional appli-
cations have been invented. Robertson (1986) made a thorough review of non-pottery
applications of smectitic clays including bentonites and kaolins throughout history.
Among these applications, medical uses of clays are of particular interest (Gomes &
Silva, 2006).

# Copyright 2011 the European Mineralogical Union and the Mineralogical Society of Great Britain & Ireland
DOI: 10.1180/EMU-notes.9.1
342 G. E. Christidis

At present, clays are important industrial rocks with numerous industrial applications
due to their outstanding physical and chemical properties. The most important clays uti-
lized by industry are kaolins, bentonites, sepiolite/palygorskite clays and common
clays and shales. The important properties stem from the presence of fine-grained
clay minerals, usually ,2 mm in size, the composition of the non-clay minerals, the
presence or absence of organic matter, the type and amount of exchangeable ions and
soluble salts and the clay texture (Grim, 1968; Bennett & Hulbert, 1986). Clays are
used by industry either in bulk form without significant beneficiation or after application
of various processing techniques (Pruett & Pickering, 2006; Murray, 2007). Processing,
which may involve various physical or chemical treatments such as wet grinding, mag-
netic separation, selective flocculation, flotation, and treatment with inorganic or
organic compounds (Pickering & Murray, 1994; Christidis et al., 1997; Pruett & Pick-
ering, 2006; Lagaly et al., 2006), removes impurities, modifies the properties and
increases the clay mineral content and hence increases the added value. Major end
users of clays are agriculture, process industries (ceramics, paper, plastic, rubber and
food industries), the environmental protection or/and remediation sector, engineering,
construction, the pharmaceutical industry, etc. Recently, much attention has been
paid to the application of clays in the formulation of nanocomposites (LeBaron et al.,
1999; Xue & Pinnavaia, 2007).
According to Guggenheim & Martin (1995), the definition of clay was formalized by
Georgio Agricola in 1546. The accepted definition for clay, given by the Joint Nomen-
clature Committee (JNC) of AIPEA and The Clay Minerals Society, is a naturally occur-
ring material composed primarily of fine-grained minerals, which is generally plastic at
appropriate water contents and will harden when dried or fired (Guggenheim & Martin,
1995). Except for phyllosilicates, it may contain other materials which impart plasticity
and harden when dried or fired. Note that the JNC avoids using the term rock for clay.
By definition the ‘synthetic clays and clay-like materials’ (e.g. laponite, layer double
hydroxides) are not considered clays, although they may be fine-grained and plastic
and they may harden upon drying and firing (Bergaya & Lagaly, 2006). However, the
natural analogue of layer double hydroxides (LDH), called hydrotalcite, may be con-
sidered as clay. Note that particle size is not taken into account in the aforementioned
definition. This is because different scientific disciplines consider different particle
sizes for clays. In geology and soil science, for instance, the size used is ,2 mm; in sedi-
mentology it is 4 mm; and in colloid science, 1 mm (Guggenheim & Martin, 1995;
Moore & Reynolds, 1997).
According to the JNC, clay minerals are phyllosilicates and minerals that impart plas-
ticity to clay and which harden upon drying or firing (Guggenheim & Martin, 1995).
This definition includes synthetic minerals (Bergaya & Lagaly, 2006) and is not
restricted to phyllosilicates as was the case in previous definitions (Bailey, 1980). In
this sense it may well include LDHs. According to Guggenheim & Martin (1995), if
a non-silicate mineral, such as an oxyhydroxide, imparts plasticity to clay and
hardens upon drying or firing then it can be considered a clay mineral. In contrast, min-
erals that do not impart plasticity to clays or do not harden upon drying or firing are
associated minerals or phases. In several industrial clays such as bentonites, associated
 Industrial clays 343

phases may also include X-ray amorphous matter such as volcanic glass or gels. The
term clay should never be used as a mineral term and hence it should not be considered
as a synonym for clay mineral, because clays consist of more than one mineral
(Guggenheim & Martin, 1995). This notation will be adopted in the present work.
Industrial clay resources have been classified into four categories (Harvey & Murray,
1997; 2006)
(1) Clays of category 1 are high-quality, high-technology clays, which require major
investment for large-tonnage production to supply both local and international
markets. Typical examples of this category are the sedimentary kaolins of
Georgia and the SE USA, the hydrothermal kaolins of Cornwall in the UK and
the bentonites from Milos, Greece.
(2) Clays of category 2 are specialty clays, which require advanced technologies for
small-tonnage specialty markets, locally and internationally. Examples of this
category include the halloysite deposits of New Zealand and the hectorite
deposit at Hector, California.
(3) Clays of category 3 include low-technology clays of moderate quality, which
mainly supply local markets. Typical examples are the kaolin deposits of
central and eastern Europe in the Czech Republic, Ukraine and Germany and
the bentonite deposits at Wyoming, USA.
(4) Clays of category 4 justify little or no processing and may be suitable for large-
tonnage local markets. These clays may be of moderate to high quality but are
considered to be uneconomic due to isolation from markets, politically or econ-
omically unstable locations, or unfavourable legislative environments.
The different clay types are presented below.

2. Structure of clay minerals

As structural details are beyond the purpose of this contribution, only a general overview
of the structure of clay minerals is presented here. Detailed descriptions of clay mineral
structures can be found elsewhere (Bailey, 1984; Velde, 1985; Moore & Reynolds,
1997; Meunier, 2005; Brigatti et al., 2006) and the present overview has been based
on these references. Clay minerals are layer silicates or phyllosilicates, which can be
considered to consist of two modular units: a sheet of corner-linked tetrahedra and a
sheet of edge-linked octahedra. The two-dimensional tetrahedral sheets are continuous
with composition T2O5, whereby T is a tetrahedral cation, namely Si, Al or Fe (Fig. 1a).
Silicon is located in the centre of each tetrahedron and oxygens form the four corners.
The corner-linked tetrahedra sharing three corners each via the so-called basal oxygens,
form a hexagonal mesh pattern (Fig. 1a). The fourth tetrahedral corner, known as apical
oxygen, points in a direction normal to the sheet and at the same time is part of the adja-
cent octahedral sheet consisting of edge-linked octahedral (Fig. 1b). The smallest struc-
tural unit consists of three octahedra. If all three octahedra are occupied by bivalent
cations (Mg, Fe2þ) the sheet is ‘trioctahedral’, whereas if two of the three octahedra
are occupied by trivalent cations (Al3þ, Fe3þ) the sheet is ‘dioctahedral’.
344 G. E. Christidis

a Tetrahedral sheet

Ob
Ob [TO4]
Ob
Ob
Ob b
Oa Ob

Ob a

b Octahedral sheet

cis-configuration
Oa
OH M Oa
cis
OH Oa
Oa b

trans-configuration
OH
Oa Mtrans
Oa
a
Oa Oa
OH

Fig. 1. (a) Schematic presentation of tetrahedron and tetrahedral sheet. (b) octahedral sheet with cis and trans
configuration. Oa refer to the apical oxygens and Ob refer to the basal oxygens. The shaded octahedra depict
trans-vacant configuration. Modified from Brigatti et al. (2006).

The plane of junction between the tetrahedral and octahedral sheets comprises the
apical oxygens shared by the tetrahedra and the octahedra and unshared OH groups
which lie at the centre of each tetrahedral six-fold ring at the same level as the apical
oxygens (Fig. 1b). The assemblage of tetrahedral and octahedral sheets yields the
‘layer’. Two types of layers have been recognized: the ‘1:1 layer’ or T-O layer in
which a tetrahedral sheet lies over an octahedral sheet and the ‘2:1 layer’ or T-O-T
layer in which an octahedral sheet is linked with two tetrahedral sheets (Fig. 2). The
1:1 structure is typical of the kaolinite and the serpentine group and has an unshared
plane of OH ions in the octahedral sheet. The 2:1 structure is typical of most clay min-
erals such as micas, talc and smectites. The space between two successive 1:1 and 2:1
layers is the ‘interlayer’, which is empty if the layers are electrostatically neutral. If the
layer bears an excess charge, known as ‘layer charge’ then it is neutralized by various
interlayer materials such as cations (Ca, Na, Mg, K), hydrated cations and hydroxide
octahedral groups. The hydroxide interlayer often forms an additional octahedral
sheet yielding a 2:1:1 or T-O-T-O layer. This structure is typical of chlorites.
 Industrial clays 345

1:1 layer

2:1 layer
O

Fig. 2. Three-dimensional assembly of tetrahedral and octahedral sheets and formation of the 1:1
layer and the 2:1 layer. Large white circles are oxygens, large grey circles are hydroxyls, small white
circles are octahedral ions and small black circles are tetrahedral ions. Modified from Moore & Reynolds
(1997).

The sum of a layer and an interlayer is the ‘structural unit’ (Fig. 2), with thickness
of 7– 18 Å, depending on the type of layer and the interlayer content of each clay
species, which corresponds to a specific chemical ‘formula unit’. According to the
alignment of layers various stacking sequences may form which are known as
‘polytypes’ (Figs 3, 5). The basal spacing between two successive units is 7 Å for the
1:1 layer, 10 Å for the 2:1 layer and 14 Å for the 2:1:1 layer. Piling of many structural
units yields a ‘clay particle’ usually smaller than a few micrometres, which is the
most common component of the industrial clays and the ‘clay fractions’ of most sedi-
mentary rocks.
Classification of layer silicates is based on the layer type (i.e. 1:1, 2:1 or 2:1:1 layer).
Within each group ‘layer charge’ is used as a criterion for classification. The layer
charge stems from substitutions in the tetrahedral or/and the octahedral sheet. Within
these groups, further subdivision is made into subgroups according to the dioctaheral
or trioctahedral character of the layer silicate. A schematic classification of layer sili-
cates is given in Table 1. The table includes clay minerals, brittle micas, serpentine,
talc and palygorskite as well as fibrous layer silicates, namely sepiolite and palygorskite.
This set of criteria may not work well with mixed-layer clay minerals (see below),
and for the transitional boundaries between micas and illites, illites and vermiculites,
vermiculites and chlorites and vermiculites and smectites (Moore & Reynolds, 1997).
Mixed-layer clay minerals are not included in Table 1 because the components that
form them are listed in Table 1. Some of the layer silicates listed in Table 1 such as
serpentine talc and pyrophyllite, are not considered clay minerals sensu stricto and
will not be presented in detail.
346 G. E. Christidis

Table 1. Classification of layer silicates, with emphasis on clay minerals.

Layer Group Subgroup Species
type
1:1 Serpentine-kaolin (X  0) Serpentine (Tr) Chrysolite, antigorite, lizardite,
berthierine, odinite
Kaolin (Di) Kaolinite dickite, nacrite,
halloysite
2:1 Talc-pyrophyllite Talc (Tr) Talc
(X  0) Pyrophyllite (Di) Pyrophyllite
Smectite Tr-smectite Saponite, hectorite, stevensite,
sauconite
(X  0.2– 0.6) Di-smectite Montmorillonite, beidellite,
nontronite
Vermiculite Tr-vermiculite Trioctahedral vermiculite
(X  0.6– 0.9) Di-vermiculite Dioctahedral vermiculite
Mica Tr-illite
(X  0) Di-illite Illite, glauconite
Brittle micas Tr-micas Biotite, phlogopite, lepidolite
(X  0) Di-micas Muscovite, paragonite
Tr-brittle micas Clintonite
Di-brittle micas Margarite
2:1:1 Chlorite Tr-chlorite Clinoclore, nimite, chamosite
(X variable) Di-chlorite Donbassite
Di-Tri-chlorite Cookeite, sudoite
2:1 Sepiolite-palygorskite Sepiolite Sepiolite
inverted ribbons (fibrous clays) (X variable) Palygorskite
Palygorskite
X ¼ layer charge per half unit cell, Tr ¼ trioctahedral, Di ¼ dioctahedral

It has been proposed that illite has a distinct layer charge, which is slightly lower than the rest of micas i.e. 0.9 (Srodon
et al., 1992).

2.1. The 1:1 layer silicates

2.1.1. Kaolin-serpentine group
The kaolin-serpentine group comprises 1:1 dioctahedral and trioctahedral minerals
(Table 1). The 1:1 dioctahedral minerals are important constituents of industrial clays.
The ‘kaolin group’ includes the 1:1 dioctahedral minerals kaolinite (with various
degrees of structural disorder), dickite, nacrite and halloysite (7 Å or 10 Å). The minerals
of the kaolin group do not possess layer charge and have the general structural formula
Al4Si4O10(OH)8. Substitutions within the structure have not been observed. The negative
charge observed in some kaolinites has been attributed to the presence of fine-grained
impurities such as vermiculite mica or smectite. Kaolinite is a triclinic mineral with
well formed pseudohexagonal crystals (Fig. 4) and layer stacking characterized by inter-
layer shift along X  a/3, which leads to a 1M stacking pattern (Fig. 5). Dickite and
nacrite are polytypes of kaolinite, i.e. they form from different stacking of successive
layers and usually display thicker crystals than kaolinite. In dickite, alternating B and
C site occupancy yields a 2M pattern. In nacrite the 1:1 layers are rotated by 1808 and
the unit cells are displaced along the b axis by b/3 (Fig. 3) (Moore & Reynolds, 1997).
Halloysite is a hydrated form of kaolinite. Halloysite exhibits a variety of mor-
phologies namely tubes, spheres, plates, oblate spheroids, stubby cylinders and irregular
 Industrial clays 347

shapes (Giese, 1988; Pruett & Murray,

1993) (Fig. 4). The 10 Å halloysite
variety contains a layer of water 2.9 Å
thick between the silicate layers (10 Å
halloysite), which can be removed by
gentle heating or in vacuum (7 Å halloy-
site). It is generally accepted that there a/3
is a continuum between the two extr-
emes (Giese, 1988). The presence of
water has been attributed to structural
disorder (Brindley, 1984), to structural Kaolinite Dickite
characteristics of halloysite (Giese,
1988) or to the presence of exchange-
able cations, which balance layer
charge due to substitution of tetrahedral
Si by Al3þ (Bailey, 1993).
The ‘serpentine group’ [general struc- b/3
tural formula M6(Si4 – xAlx)O10(OH)8, Nacrite
where M ¼ Mg, Fe, Ni, Al, Mn] com-
prises 1:1 trioctahedral layer silicate Fig. 3. Stacking sequence in minerals of the kaolin
group. Black circles indicate occupied octahedral
minerals with differences in the chemi- sites and white squares correspond to vacant
cal composition and the extent of substi- octahedral sites. Modified from Moore & Reynolds
tutions in their structure. For instance, (1997).
antigorite, chrysotile, greenalite and
amesite do not display substitutions in their structure, whereas cronstedtite, berthierine
and odinite display both tetrahedral and octahedral substitutions. However, when there
is substitution in one sheet of the 1:1 silicate, there is almost always a compensating
substitution in the other sheet which maintains neutrality. The most common minerals
of this group are lizardite, antigorite and chrysotile with the general structural
formula Mg6Si4O10(OH)8. Lizardite has small amounts of Al in octahedral and tetrahe-
dral sheets. Serpentine minerals are not constituents of industrial clays and will not be
considered further.

2.2. The 2:1 layer silicates

2.2.1. Pyrophyllite-talc group
‘Pyrophyllite’ [Al2Si4O10(OH)2] and ‘talc’ [Mg3Si4O10(OH)2] are 2:1 layer silicates
ideally without layer charge. Nevertheless, natural materials often have limited substi-
tutions. The primary atomic forces holding the 2:1 layers together are weak van der
Waals bonds, which account for the soapy and soft nature of talc and pyrophyllite
(Evans & Guggenheim, 1988). Although they are not considered constituents of indus-
trial clays, they serve as structural prototypes of several clay minerals with layer charge,
which are main constituents of industrial clays. The disordered variety of talc is ‘kero-
lite’ (Brindley, 1984). The existence of a small number of octahedral vacancies in talc
348 G. E. Christidis

Fig. 4. SEM images of clay minerals: (a) pseudohexagonal crystals of kaolinite; (b) tubular crystals of
halloysite; (c) spheroidal crystals of halloysite; (d) wavy subhedral montmorillonite crystals (from
Fesharaki et al., 2007); (e) flaky illite crystals; and ( f) fibrous illite. Images courtesy of The Clay Minerals
Society and the Clay Minerals Group of the Mineralogical Society (Images of Clay Gallery, available at
www.minersoc.org/pages/gallery/claypix/index.html).

yields ‘stevensite’, a smectite with small layer charge, which stems from these vacancies
(see below).

2.2.2. Mica group

‘Micas’ are 2:1 layer silicates characterized by a layer charge of 1 per half unit cell
(p.h.u.c.) (Table 1). This charge stems (1) from substitutions of trivalent ions (R 3þ)
for Si4þ in the tetrahedral sheet, (2) from substitution of univalent (R þ) for bivalent
 Industrial clays 349

(R 2þ) cations, or substitution of bivalent a Z

(R 2þ) for trivalent (R 3þ) cations in the
octahedral sheet, and (3) from vacancies
in the octahedral sheet. The layer charge
is balanced by large univalent interlayer
cations (bivalent in the case of brittle –X
micas) which are fixed i.e. not exchange-
b –X c –X
able. The most common interlayer
þ
cation in true micas is K , whereas in
brittle micas is Ca2þ. Depending on the
stacking of layers, different polytypes
may form. The most common polytypes
observed in micas are 1M, 2M1 and 2M2, Y
whereas 3T and 2O polytypes are less
common. The micas can be dioctahedral d –X
(muscovite, paragonite, celadonite) or
trioctahedral (biotite, phlogopite lepido-
lite, annite). Similarly, brittle micas are
dioctahedral (margarite) or trioctahedral
(clintonite). Y
‘Illite’ is a clay mineral, and a member
of true mica group, which forms platy, Fig. 5. (a) Translation of the unit cell in the X-Y
fibrous or lath-like crystals (Fig. 4). It plane in micas caused by layer rotation. (b) The 1M
polytype. (c) The 2M1 polytype. (d) The 3T
differs from muscovite because (1) it con- polytype (modified from Moore & Reynolds, 1997).
tains more Si, Mg and H2O, (2) it contains
less tetrahedral and interlayer K, (3) it often contains a small amount of expandable
layers, which do not exceed 10%. Illite has a smaller layer charge than muscovite
which is constant for all illites (0.89 p.h.u.c. according to Środoń et al., 1992), although
this has not become widely accepted yet (Moore & Reynolds, 1997). Illite may be
present as 1M or 2M1 polytypes (Fig. 5). Other than K, illite may contain interlayer
Na or NHþ þ
4 . Na-illite is a rare mineral and is known as bramallite. When NH4 is the pre-
þ
dominant interlayer cation, the NH4 -mica is called tobelite (Higashi, 1982). The pres-
ence of tobelite layers have been increasingly recognized in sediments over recent years
(e.g. Drits et al., 2002, 2007).

2.2.3. Smectite group

Smectites are 2:1 layer silicates with layer charge varying from 0.2 to 0.6 equivalents
p.h.u.c. due to substitutions in the tetrahedral or the octahedral sheet, or to the existence
of vacancies in the octahedral sheet. The layer charge is balanced by the interlayer
cations which are exchangeable, leading to an important property, the cation exchange
capacity (CEC). Smectites form small crystals. usually ,0.5 mm in size, and their
crystal shape varies from subhedral lamellae with irregular outlines (Fig. 4), to euhedral
lamellae with rhombic outline and lath and ribbon crystallites. Smectite layers display
350 G. E. Christidis

turbostratic stacking without ordering, i.e. they are randomly stacked one on top of the
other, like a pile of playing cards.
Although smectites are characterized as low- and high-charge, so far there has been
no acceptable classification scheme. Christidis et al. (2006), based on the XRD charac-
teristics of K-saturated ethylene glycol-solvated smectite, proposed a classification of
dioctahedral smectites according to layer charge (Fig. 6). Low-charge smectites have
layer charge of ,0.42 equivalents p.h.u.c. and high-charge smectites have layer
charge .0.47 equivalents p.h.u.c. Smectites with layer charge between 0.42 and 0.47
equivalents p.h.u.c. are intermediate-charge smectites. Due to their low layer charge
the interlayer cations are fully hydrated, yielding the remarkable property of swelling
when hydrated. Swelling may also take place in the presence of certain polar organic
compounds such as ethylene glycol or glycerol. Their ability to expand when exposed
to organic vapours is a criterion for recognizing smectite from other clay minerals.
They are the main constituents of bentonites, which are important industrial clays.
Smectites are either dioctahedral or trioctahedral. In order to consider the compo-
sition of the various smectites it is useful to use as a reference the pyrophyllite
[Al2Si4O10(OH)2] or the talc [Mg3Si4O10(OH)2]. In the dioctahedral smectites, substi-
tution of octahedral Al by Mg in the composition of pyrophyllite yields montmorillonite,
the most common smectite, whereas substitution of tetrahedral Si by Al yields beidellite
and substitution of Si by Fe3þ yields nontronite. Hence the layer charge is in the octa-
hedral sheet in montmorillonite and in the tetrahedral sheet in beidellite and nontronite.
Another major chemical element present in the octahedral sheet is Fe3þ, which does not,
however, contribute to layer charge. Smectites containing .0.3 Fe3þ atoms p.h.u.c. are
known as Fe-rich smectites, either Fe-rich montmorillonite or Fe-rich beidellite (Güven,
1988). Finally, volkonskoite is a rare dioctahedral smectite with Cr3þ as the main octa-
hedral cation and with the layer charge located mainly in tetrahedral sites.
Most natural dioctahedral smectites have compositions between those of montmoril-
lonite and beidellite because pure end members are extremely rare. In natural smectites,

Dioctahedral smectites

d001 > 16.6 Å 16.6 Å < d001 < 16 Å d001 < 15.0 Å
Rational order of higher- Irrational higher-order 003 reflection at
order reflections basal reflections 4.60 – 4.75 Å

Characterization Characterization Characterization

Low-charge smectites Intermediate-charge High-charge smectites
Proportion of 17 Å smectites Proportion of 17 Å
layers > 0.75 Proportion of 17 Å layers < 0.3
layers 0.4 – 0.70

Fig. 6. Classification scheme of dioctahedral smectites by XRD, according to their layer-charge

characteristics.
 Industrial clays 351

beidellites have .50% of layer charge AlMg

in tetrahedral sites and montmorillonites
.50% of layer charge in octahedral
sites. Because of their octahedral compo-
sition and layer charge, montmorillonites

.3
=0
can be distinguished into Otay, Cham-

Fe
OT
bers, Tatatilla and Wyoming types Fe-M
(Schultz, 1969, Newman & Brown, CH

1987) (Fig. 7). The Otay, Chambers and TA WY

Tatatilla montmorillonites are often des- Fe-Bi
Bi
cribed as Cheto-type montmorillonites.
However, in the Glossary Terms of The
Clay Minerals Society, use of the term AlAl AlFe
Al-rich dioctahedral montmorillonite is Fig. 7. Classification scheme of dioctahedral smec-
suggested (CMS, 2009). It is suggested tites according to Güven (1988). TA ¼ Tatatilla-type
here that the term Al-rich dioctahedral montmorillonite, OT ¼ Otay-type montmorilonite,
smectite with predominantly octahedral CH ¼ Chambers-type montmorilonite, WY ¼
charge be used. In trioctahedral smectite, Wyoming-type montmorillonite, Bi ¼ beidellite,
Fe-M ¼ Fe-rich montmorillonite, Fe-Bi ¼ Fe-rich
hectorite is derived from talc com- beidellite.
þ
position by substitution of Li for octa-
2þ
hedral Mg . In saponite the layer charge stems from substitution of tetrahedral Si by
Al3þ. Saponite is different from the other smectites as part of the negative tetrahedral
charge is balanced by substitution of octahedral Mg2þ by trivalent cations, Al3þ or Fe3þ,
i.e. the octahedral sheet bears a positive charge. Stevensite is a different trioctahedral smec-
tite in the sense that the layer charge stems from a small deficiency in octahedral cations, not
from substitutions in the structure (Brindley, 1984). Stevensite can be considered as a phase
with a variable number of octahedral vacancies which yield smectite-like and talc-like
domains (Christidis & Mitsis, 2006). Finally, sauconite is a rare trioctahedral smectite
with Zn2þ as the main octahedral cation, and layer charge located mainly in tetrahedral
sites. A classification scheme of smectites is shown in Table 2.
Apart from the chemical variations, dioctahedral smectites also display variability
in octahedral-site occupancy. An octahedral site may exist either on the mirror

Table 2. Classification scheme of natural smectites (adapted from Güven, 1988).

Dioctahedral smectites Trioctahedral smectites
Ratio between Predominant Smectite species Predominant Smectite
tetrahedral (zt) and octahedral octahedral species
octahedral (zo) charge cation(s) cation(s)
zo . zt Al3þ(R 2þ) Montmorillonite Mg Stevensite
Mg(Li) Hectorite
3þ
zo , zt Al Beidellite Mg Saponite
Fe3þ Nontronite Fe2þ Iron saponite
Cr3þ Volkonskoite Zn Sauconite
V3þ Vanadium smectite

Octahedral substitutions.
352 G. E. Christidis

plane (site M1) or to the left and right of the mirror plane (site M2) (Fig. 1). The
hydroxyl configuration around the M1 and M2 sites is different. M1 sites have a trans
configuration because the hydroxyls are located across the site, whereas M2 sites
have a cis configuration i.e. the hydroxyls are adjacent to one side of the site (Fig. 1).
This difference seems to control the dehydroxylation temperature of smectites,
because cis-vacant smectites have greater dehydroxylation temperatures (650 – 7008C)
compared to their trans-vacant counterparts (,6008C) (Drits et al., 1998), and dictates
the performance of smectites in foundry applications which are characterized by high
temperatures.

2.2.4. Vermiculite group

Vermiculites are 2:1 minerals with layer charge of 0.6 – 0.9 equivalents p.h.u.c.,
which arises mainly from substitution of Al3þ for Si in tetrahedral sites. Similar to
smectites, layer charge is balanced by the interlayer cations which are hydrated and
can be exchanged. Clay-sized vermiculites may be dioctahedral or trioctahedral
and result from pedogenesis or diagenesis, whereas most macroscopic varieties
are trioctahedral, forming platy crystals like the micas and result from alteration
of mica or chlorite. A typical structural formula of trioctahedral vermiculite is
Mg3(Si3Al)O10(OH)2Mg(exch)0.5(H2O)4, where Mg(exch) is exchangeable Mg. However,
there are several types of substitution in the octahedral and the tetrahedral sheet. Hence,
octahedral Mg is replaced by Fe3þ, Fe2þ, Al3þ and Ti4þ, and Si is replaced by Fe3þ
except for Al3þ.

2.2.5. Chlorite group

Chlorites are considered to be either a negatively charged 2:1 layer having the general
2
formula [(R 2þ, R 3þ)3(Si4 – xR 3þ
x )O10(OH)2] and a positively charged interlayer octa-
hedral sheet with general formula: ([R , R 3þ)3(OH)6]þ, or as a 2:1:1 or 2:2 layer sili-
2þ

cate (older literature) although the latter approach has been rejected by AIPEA (Moore
& Reynolds, 1997). The most common octahedral cations are Mg2þ, Al3þ, Fe2þ and
Fe3þ. The vast majority of chlorites are trioctahedral in both octahedral sheets (tri-
chlorites in Table 1). The chlorite which is dioctahedral in both octahedral sheets
(di-chlorite) is a rare mineral called donbassite. Although in the past there have been
several classification schemes for chlorites with several different species, a simplified
nomenclature based on the main divalent octahedral cations Mg, Fe, Ni and Mn has
been accepted. The end members are clinochlore, chamosite, nimite and pennantite.
Chlorites are characterized by different polytypes produced by different translations
of 2:1 layers relative to the hydroxide sheets and by the orientation of the hydroxide
octahedral sheet compared to the octahedral sheet in the silicate layers. Two polytypes
have been recognized, I and II. Translations of 2:1 layers in the X-Y plane is consistent
with the geometric constraints of hydrogen bonding between O and OH surfaces and two
possibilities namely a (the hydroxide sheet cations are located directly over the tetrahe-
dral Si positions) and b (the hydroxide sheet cations are located directly over the octa-
hedral 2:1 sheet cations). Position a is not favoured energetically.
 Industrial clays 353

2.2.6. Sepiolite-palygorskite group

Sepiolite [Mg8Si12O30(OH)4(OH2)4 . n(R 2þ(H2O)8] and palygorskite [MgAl3Si8O20
(OH)3(OH2)4 . n(R 2þ(H2O)4] are 2:1 layer silicates which differ from other clay min-
erals because the octahedral sheets are continuous in one dimension only in ribbons
and because the tetrahedral sheets are also divided into ribbons by inversion of every
two or three rows of tetrahedra (Fig. 8). These structural features create channels
between the ribbons, which are larger in sepiolite (4 Å  9.5 Å) compared to palygors-
kite (4 Å  6 Å). Palygorskite is dioctahedral and sepiolite trioctahedral (Newman &
Brown, 1987). Both minerals have elongated habits often forming bundles of lath-like
or fibrous crystals. They contain two types of water, structural water coordinated to
the octahedral cations and zeolitic water, which is loosely bound in the channels. Due
to the existence of channels, palygorskite and sepiolite have large microporous volumes.

2.2.7. Mixed layer clay minerals

The term ‘mixed-layer clay minerals’ refers to clay minerals which are formed by at
least two types of intergrown layers, not physical mixtures, and is synonymous with
‘interlayered’ or ‘interstratified’ clay minerals (Moore & Reynolds, 1997). The different
types of layers are stacked perpendicular to (001) and stacking may either be in a
random manner (random interstratification) in partially regular or in regular manner
(regular or ordered interstratification). The degree of ordering refers to the probability
of one type of layer X to be followed by another type of layer Y. In random

Palygorskite
b = 17.9 Å

4 Å × 6Å

Sepiolite
b = 26.95 Å

4 Å × 9.5Å

Fig. 8. Schematic structural diagrams of palygorskite and sepiolite. Due to the inversion points, the octahedral
sheets do not maintain continuity and open cages form in which water molecules reside.
354 G. E. Christidis

interstratification the probability of an X layer being followed by a Y layer is the same as

the probability of an X layer being followed by an X layer. In perfectly regular interstra-
tification, when the weight fraction of layer Y is  0.5, the probability of forming YY
pairs is zero. Mixed-layer clay minerals are very common phases in many types of
clays, including bentonites (K-bentonites), kaolins and common clays and shales. The
most common mixed-layer clay minerals are illite-smectite (I-S), chlorite-smectite
(C-S) and kaolinite-smectite.

3. Properties of clays
3.1. Particle size and shape
Our knowledge about the particle sizes and shapes of clay minerals has been enriched
through detailed studies using transmission and scanning electron microscopy (TEM/
SEM). The size and shape of particles shape of clay minerals and aggregate character-
istics affect the physical properties of industrial clays. Clay minerals are, in principle,
,2 mm in size. Nevertheless, minerals of the kaolin group often form larger crystals,
which, in the case of dickite, may reach up to 20 mm (Beaufort et al., 1998), although
disordered minerals of the group are considerably ,2 mm. Smectites tend to form
smaller crystals and the average smectite crystal size in bentonites is ,0.5 mm (Grim
& Güven, 1978; Christidis, 1995). Smectites exceeding 2 mm in size are uncommon.
Palygorskite and sepiolite form fibres which often exceed 2 mm in size and have stria-
tions (Martin Vivaldi & Robertson, 1971; Jones & Galán, 1988).
The small crystal size yields a large specific surface area for most clay minerals and
this can be both external and internal. This is especially true for smectite, vermiculite,
sepiolite and palygorskite, which have large fractions of internal surface area. The chan-
nels in the structure of sepiolite and palygorskite contribute to the internal surface area.
The existence of large specific surface area creates a great degree of surface reactivity,
which can be increased further by treatment with inorganic acids, a process known as
‘acid activation’ (Kaviratna & Pinnavaia, 1994; Christidis et al., 1997; Myriam et al.,
1998; Balcı, 1999; Nguetnkam et al., 2005). Acid activation is applied to smectite,
sepiolite and palygorskite. The specific surface areas of various clay minerals are
listed in Table 3.

Table 3. Cation exchange capacity (CEC) and specific surface area of clay minerals.
Clay mineral CEC (meq/100 g) Specific surface area (m2/g)
Kaolinite 1 – 15 10– 20
Illite 10 – 40 50 –100
Chlorite 10 – 40 10– 20
Smectite 70– 150 10 – 800
Vermiculite 130– 210 10 – 800
Sepiolite/palygorskite 10 – 45/5 – 30 150 – 900

Depending on the fraction of internal specific surface area.
 Industrial clays 355

Crystal shape varies in different clay minerals even within the same group. Kaolin-
group minerals may form euhedral to subhedral pseudo-hexagonal crystallites in kaoli-
nite (e.g. Martin Vivaldi & Robertson, 1971; Pruett & Murray, 1993; Beaufort et al.,
1998; Psyrillos et al., 1999), to blocky pseudo-hexagonal or trapezoidal crystals in
dickite (Beaufort et al., 1998), and tubular, spherical or platy in halloysite (Giese
1988; Pruett & Murray, 1993) (Figs 4, 9). Smectite crystal morphology may vary
from rhombic to pseudo-hexagonal, lamellar to lath-shaped and Fe-rich smectites (non-
tronite) can be even fibrous (Grim & Güven, 1978; Güven, 1988, Christidis et al.,1995). Q1
Crystals may also vary from euhedral to subhedral, with the latter being more common
(Figs 4, 9). Sepiolite and palygorskite always form fibrous crystals (Fig. 9). Illite may
form euhedral to subhedral pseudo-hexagonal to lath-like crystals. Diagenetic hairy
illite is also common in sedimentary rocks (Fig. 4). Finally, chlorite forms pseudo-
hexagonal crystals usually in well defined aggregates (Fig. 9).
Clay minerals tend to form aggregates in clay deposits. Typical examples of well
known aggregates are the kaolin booklets forming vermiform aggregates, which are
common in most kaolins (Fig. 9) or the smectite aggregates which are characterized
by a honeycomb texture (Fig. 9). In the latter case, even when dispersed in water, smec-
tites tend to form aggregates rather than single crystals. Vermiform aggregates are
common for other clay minerals like chlorites (Fig. 9). Aggregates of sepiolite and paly-
gorskite usually form bundles (Fig. 9). The existence of aggregates has an adverse affect
on the physical and chemical properties of industrial clay minerals. Hence, vermiform
kaolin aggregates require delamination before application in the paper industry (Pruett
& Pickering, 2006; Murray 2007). The specific surface area of smectite aggregates is
also lower than the actual specific surface area of isolated smectite crystallites. In
order to improve rheological properties of smectite, particles have to be disaggregated
by a strong shearing force (Odom, 1984). Differences in the particle size of smectite
aggregates may explain variation in the ion-exchange properties of smectites (Neal &
Cooper, 1983).

3.2. Ion-exchange properties

Cation exchange capacity is a characteristic property of all clay minerals and is of par-
ticular importance for smectites and vermiculites and, to a lesser degree, of sepiolite and
palygorskite (Table 3). It is normally measured at pH 7. It is related to substitutions in
the tetrahedral and/or the octahedral sheet which create a charge deficit known as the
layer charge, and to adsorption or dissociation of protons at crystal edges, often referred
to as interrupted bonds, at the edges of the crystals. The charge which results from struc-
tural substitutions is the ‘permanent charge’. The charge from interrupted bonds
between the structural cations and the oxygens or the OH groups of the tetrahedral
and the octahedral sheet at the edges of the crystals is known as the non-permanent
or pH-dependent charge. The layer charge is balanced by the interlayer cations (Na,
Ca, K, Mg, H and Li), which are exchangeable. The role of non-permanent charge is
more important for kaolinite than for smectite and vermiculite because of the limited
edge surface of the latter. In the case of kaolinite, the edge surface may contribute up
356 G. E. Christidis

Fig. 9. SEM images of clay-mineral aggregates: (a) smectite with honeycomb texture (after Fesharaki et al.,
2007); (b) sepiolite fibres forming bundle-like aggregates; (c) pseudo-hexagonal crystals of Fe-rich chlorite;
(d) vermiform kaolinite booklets from Eyre Peninsula, South Australia; (e) vermiform chlorite aggregates
from Strzegom pegmatite (Poland). Images courtesy of The Clay Minerals Society and the Clay Minerals
Group of the Mineralogical Society (Images of Clay Gallery, available at www.minersoc.org/pages/
gallery/claypix/index.html).

to 15% of the total surface (James & Williams, 1982), whereas in smectites it is only
1% (Sondi et al., 1997; Benna et al., 1999). The CEC due to the non-permanent charge
varies accordingly, although in smectite it may reach 14% (Anderson & Sposito, 1991). Q2
Thermodynamic study of ion exchange is carried out via exchange isotherms, but this is
beyond the scope of the present work.
Although the minerals of the kaolin group have small CEC values, the ‘anion
exchange capacity’ (AEC) can be significant. The small size of clay minerals (usually
 Industrial clays 357

,5 mm) compared to most minerals, increases the contribution of edge surfaces, com-
pared to the overall surface. Anion exchange is attributed to the charge of the edges of
the crystals, which becomes positive by adsorption of Hþ at acidic pH, forming a water
molecule. This molecule is weakly bonded and thus can be displaced easily and
exchanged by other anionic groups. Alternatively, anionic groups can replace OH
groups located at crystal edges directly (cf. Lagaly, 2006). Due to their increasing
edge surface area mentioned previously, minerals of the kaolin group display significant
AEC for phosphates (Dixon, 1989). Illite and chlorite may also develop significant AEC
(Bain & Smith, 1987). In contrast, smectite and vermiculite display limited AEC (Bain
& Smith, 1987; Borchardt 1989).
Ion exchange is important for many industrial and environmental applications of
clays with CEC such as bentonites. Although Na-smectites are more suitable for
most industrial applications, natural bentonites usually contain Ca-smectites. It is
common industrial practice to produce Na-smectites by ion exchange through
‘alkali-activation’ (Christidis & Scott, 1993; Inglethorpe et al., 1993). However,
alkali activation often yields unpredictable results (Odom, 1984) because of the het-
erogeneity of layer charge and charge distribution of smectites in bentonites (Chris-
tidis, 2008a). The layer charge and charge heterogeneity of smectites affects ion
exchange because it controls selectivity for various cations during ion exchange
(Maes & Cremers 1977; Maes et al., 1985). Different selectivity for the various
cations is the main reason for deviation from ideality during the thermodynamic
study of ion exchange (Loudelout 1987; McBride, 1994). Therefore, determination
of layer-charge distribution is a significant challenge during assessment of bentonite
deposits (see below).

3.3. Hydration and swelling

Clay minerals tend to adsorb water molecules on their external surfaces, mainly in inter-
rupted bonds (see section 3.2), or in their interlayer spaces, associated with the interlayer
cations or with the internal surface. With increasing water activity, the following modes
of clay hydration are distinguished (Güven, 1992a): (1) ‘interlamellar hydration’ via
adsorption of limited amounts of water molecules on the internal surfaces of clay par-
ticles; (2) ‘continuous (osmotic) hydration’ via unlimited adsorption of water on the
internal and external surfaces of clay particles; and (3) ‘capillary condensation’ of
free water molecules in the interaggegate and intraaggegate micropores (Fig. 10a). Inter-
lamellar hydration affects mainly smectite and vermiculite and depends (McEwan &
Wilson 1984; Güven, 1992a): (a) on the hydration energy of the interlayer cations
and the polarization energy of the water molecules by the interlayer cations; (b) on
the variation of electrostatic potential on the clay surface determined by the magnitude
and distribution of the layer charge on the silicate layers; (c) on the activity of water; and
(d) the size and morphology of clay particles and the clay fabric. At low water activity,
interlamellar hydration is controlled by the interactions of the interlayer cation with the
water molecules forming water complexes and to a lesser degree by interactions of the
water molecules with the clay surface. Interlamellar hydration leads to interlamellar or
358 G. E. Christidis

a b
Clay
Clay particle aggregate

Intra-aggregate
micropore

Fig. 10. (a) Schematic sketch showing a clay aggregate consisting of several clay particles and the
development of intra-aggregate porosity. (b) Schematic sketch showing the development of interaggregate
and intra-aggregate porosity.

crystalline swelling and may lead to intercalation of 0 to 4 discrete layers of water

molecules between individual 2:1 layers (Laird, 2006).
Continuous hydration is observed in certain smectites when the water activity
increases, principally when they are immersed in water. In this case additional water
molecules are bonded to the hydrated complexes of the interlayer cations, which are
subsequently destabilized (Güven, 1992a). This unlimited hydration is called
‘osmotic hydration’. The magnitude and localization of layer charge affects the electric
field of clay surfaces (Bleam, 1990) and this, in turn, along with the nature of the inter-
layer cation, affects the dynamics and the structure of water in suspension. In smectite,
the octahedral charge creates weak, delocalized, electric field at the surface comparable
to that of pyrophyllite (Bleam, 1990). If Mg2þ and Ca2þ are the main exchangeable
cations, the H bonds to the surface are maintained yielding small stacks of 4 –7 smectite
layers oriented with parallel c axes and randomly oriented a and b axes, known as ‘qua-
sicrystals’ (Aylmore & Quirk 1971). Smaller quasicrystals may form in suspensions of
smectites with Kþ and Csþ interlayer cations. When Naþ or Liþ smectite quasicrystals
with octahedral charge are immersed in water they may dissociate in individual layers
and they may be separated by distances ranging from one to tens of nanometres and form
“tactoids” (Güven, 1992a).
Capillary condensation of water may occur in the various pores formed in the micro-
structure of clays and occurs at high relative water pressures (Güven, 1992a; Cases
et al., 1997; Laird, 1999). Interaggregate pores may range from 0.2 to several micro-
metres in size and are often lenticular (Fig. 10b). Intraaggregate pores are smaller
(0.01– 0.2 mm) and are delineated by the boundaries of the primary particles within
the aggregates (Fig. 10b). In smectite, the water content increases exponentially at rela-
tive water pressure greater than 0.96 (Tardy & Touret, 1987). Hydration and swelling
controls many important applications of clays such as geotechnical and environmental
applications. Moreover it affects other important properties of clays such as plasticity
(see below).
 Industrial clays 359

3.4. Rheological properties

Clays are often used in suspensions. Suspensions of kaolins are used in the slip casting
processes during the manufacture of ceramics and in coating paper, whereas smectite
and attapulgite are used in drilling fluids. In these applications, the flow properties of
the suspensions are of primary importance. The science of the deformation and flow
of matter is known as ‘rheology’ (Hiemenz & Rajagopalan, 1997). The viscosity of a
liquid is a measure of the internal resistance offered to the relative motion of different
parts of a liquid. Clay suspensions can display Newtonian, Bingham plastic, shear
thickening (pseudoplastic) or shear thinning (dilatant) behaviour and may develop
yield stresss (Fig. 11a) (Lyckham & Rossi, 1999). The equations which describe the
rheological behaviour of various types of suspensions are shown in Table 4. Moreover
they can develop time-dependent phenomena such as thixotropic or rheopectic (antith-
ixotropic) behaviour (Brandenburg & Lagaly, 1988; Lagaly, 1989; Lyckham & Rossi,
1999) (Fig. 11b). The term thixotropy refers to the ability of a suspension to form a
gel upon standing and to become fluid when subjected to shear stress i.e. under stirring
or agitation.
Suspensions of kaolin develop different rheological characteristics compared to their
smectite or attapulgite counterparts. The viscosity of kaolin suspensions increases
sharply above 65% solid content. Above 20% solid content, flow is plastic and thixo-
tropy appears; above 30% kaolin, the yield stress increases considerably (Lagaly,
1989). The pH also reduces the viscosity and yield stress. Kaolin coating suspensions
may contain up to 70% kaolin minerals (Jepson, 1984). Montmorillonite, hectorite
and some saponites develop high viscosity and thixotropy when added in concentrations
of 5– 6% in water, although in low concentrations they form suspensions with Newto-
nian properties (Rand et al., 1980; Kasperski et al., 1986; Brandenburg and Lagaly,
1988; Chen et al., 1990). In contrast, Ca-Mg smectites do not yield high viscosity
and do not display thixotropic behaviour even at ahigh percentage of solids. Other
than its concentration, the viscosity of a bentonite suspensions depends on the pH,
the particle size and shape of the smectites, the type and concentration of electrolytes

a Dilatant b behavio
ur
otropic
y Antithix
Bulkle
plast
ic chel
ham Hers
Bing aviour
opic beh
Thixotr
Shear stress

Shear stress

stic
dopla
Pseu y
trop
thixo
without
curve
an Flow
toni
New

Shear rate Shear rate

Fig. 11. (a) Different types of flow curves observed in clay suspensions. (b) Typical flow curves for
concentrated thixotropic clay suspensions.
360 G. E. Christidis

Table 4. Rheological models and equations which describe the behaviour of clay suspensions.
Type of suspension Equation Type of flow
Newtonian t ¼ hg Newtonian
Bingham Plastic t ¼ tb þ hplg Plastic
Power low t ¼ Kgh Shear thickening or shear thinning
Herschel-Bulkley t ¼ tg þ Kgh Shear thinning
Where h ¼ viscosity, t ¼ shear stress, g ¼ shear rate, tb ¼ Bingham yield stress and tpl ¼ plastic viscosity. In the Herschel
Bulkley model, ty ¼ yield stress and K ¼ measure of the consistency of the fluid.

and the magnitude and localization of the layer charge of the smectites (Brandenburg &
Lagaly, 1988; Lagaly, 1989; Benna et al., 1999; Duran et al., 2000; Tombacz &
Szekeres, 2004; Christidis et al., 2006, see also Güven, 1992b, and Lagaly, 2006 for
a review). In general, low-charge smectites (see section 2.2.3) develop high viscosity
and high-charge smectites and beidellites tend to develop low viscosity (Christidis
et al., 2006).
Palygorskite and sepiolite also have significant rheological properties. When dis-
persed in water, the bundle-like aggregates break up and the fibrous crystals form a
random structure that entraps water and increases viscosity (Jones & Galán, 1988).
Their suspensions are less affected by electrolytes compared to bentonite suspensions
(Galán, 1996). The viscosity and yield point of palygorskite suspensions are affected
by the geometric characteristics of the crystals increasing with the length/width ratio
of palygorskite fibres (Neaman & Singer, 2000). Similar to smectite, and other than
the particle size and shape of the palygorskite, the rheological properties of palygorskite
suspensions are affected by the suspension concentration, the pH, and the type and con-
centration of electrolytes, (Neaman & Singer, 2000; 2004). Sepiolite is the only clay
mineral which forms stable suspensions at high temperatures (Galán, 1996).

3.5. Colour properties

Colour is an important property of clays which are utilized as fillers such as kaolins and
white bentonites (Christidis & Scott, 1997; Pruett & Pickering, 2006; Murray, 2007). In
most applications, fillers have to be white. Colour is affected by differences in particle
size and shape and by mineralogical and chemical characteristics; in general, the finer
the particle size, the whiter the colour (Grimshaw, 1971; Scott, 1990a; Christidis &
Scott, 1997; Christidis et al., 2004). The sensation of colour results from stimulus
which is transferred from the retina via the optic nerve to the brain (Billmeier &
Saltzman, 1981). The human eye responds to wavelengths in the visible spectrum
(from 400 nm for the violet to 700 nm for the red colour). Colour can be described
from its three attributes, namely hue (the dominant colour which corresponds to a wave-
length), the saturation known also as purity or chroma (the density of colour) and the
brightness known also as lightness or value or whiteness (the light intensity or visual
lightness of a colour).
The colour of industrial clay fillers is measured according to the CIE system
(Billmeier & Saltzman, 1981; Wyszecki & Stiles, 1982). The CIE system uses the three
 Industrial clays 361

primary colours (red, green and blue) to match

natural colours. However, some colours can
be matched only by negative values of red or
blue. To circumvent this problem three imagin-
ary tristimulus colours X, Y and Z, correspond-
ing approximately to red, green and blue,
respectively, which are not the pure spectral
colours, however. The most modern system of
colour measurement is the CIELab system. It
includes the tristimulus values, from which
hue, saturation and brightness (Y tristimulus
value) can be determined. Moreover parameters Fig. 12. Schematic representation of the
  
L , a and b are calculated from the X, Y and Z L a b colour system.
tristimulus values (Fig. 12). L , a and b rep-
resent lightness on a scale of zero (black) to 100 (white), redness (positive value) –
greenness (negative value) and yellowness (positive value) – blueness (negative
value), respectively (Billmeyer & Saltzman, 1981). Another useful parameter used to
describe colour is DE ab, which determines the colour difference between the sample
and perfect white, which has a value of L ¼ 100, a ¼ 0 and b ¼ 0. The common
industrial practice is simpler and brightness is often measured according to the
TAPPI system in the United States and the ISO system in Europe (Pickering &
Murray, 1994). These systems record the reflectance of light at a specific wavelength.
White clay-mineral fillers have different characteristics from coarser mineral fillers
(e.g. calcite, wollastonite, feldspars, nepheline, etc.) (Christidis et al., 2004). Clay-
mineral grains consist of crystallites of variable width, having random orientation,
loosely bound together. These crystallites are the diffuse reflection units of light. By
contrast, in coarser minerals, grains comprise a single crystal or a limited number of
crystals with well defined boundaries. Each calcite grain is a diffuse reflection unit.
During comminution, a clay grains break up into smaller parts and simultaneously a
fraction of crystallites may display incipient delamination. Delamination increases the
number of diffuse reflection units because the
actual crystallite size is not affected (Fig. 13).
In coarser minerals, the number of reflecting a
units increases considerably with comminution, S

because the grains break up continuously to

smaller units creating a smoother surface
thereby enhancing light reflectance (Christidis b
S
et al., 2004) (Fig. 13). The geological history
of the materials controls the crystallite-size dis-
tribution (single vs. multi-modal distributions) Fig. 13. Schematic sketch showing the
surface roughness of clay and coarse-
and thus delamination. These inherent charac-
grained filler. (a) Clay-mineral fillers; (b)
teristics are not modified by other beneficiation Coarse-grained fillers (calcite, wollastonite,
processes, which mainly remove colouring feldspar, etc). S ¼ filler surface (from
admixtures such as Fe- and Ti-oxides. Christidis et al., 2004).
362 G. E. Christidis

3.6. Plasticity
Plasticity is the property of a material to be deformed under stress and to retain the new
shape after the stress is removed. It is a characteristic property of clays because other
minerals which may be of clay size are not plastic. The nature of plasticity is related
to the water molecules which are adsorbed on the clay mineral surfaces forming a
rigid film with certain order, which links together clay particles (Grimshaw, 1971).
The clay particles form coherent networks, which can deform and retain their shape
after the stress is removed. If the clay aggregates form band-like textures, the
network can be deformed by rotation of particles (Lagaly, 2006). Alternatively, if the
particles form card-like textures (i.e. the particles are linked with an edge-to-face
mode forming T-type contacts) they can be shifted without loosing their coherency, pro-
vided that the water content does not exceed a certain limit (Lagaly, 2006).
The plasticity of clays is affected by the type of clay mineral, the amount of water
present, the particle size, shape and size distribution of clay aggregates, the specific
surface area of the clay particles, the orientation of particles in the aggregates, the
nature of non-clay minerals and the previous history of the clay (Grimshaw, 1971).
Clay minerals impart plasticity, whereas the non-clay minerals effectively reduce the
plasticity of clays. In minerals with CEC values such as that of smectite, plasticity is
affected by the type of the exchangeable cation; Na-smectites develop greater plasticity
than their Ca-counterparts (Bain, 1971). Sepiolite and palygorskite contain water in their
channels, which does not contribute to the development of plasticity. Illite-rich clays in
general also have small degrees of plasticity.
The presence of water is critical for the development of plasticity. At water contents
of less than a lower limit, the clay particles come in contact and plasticity is lost,
whereas at water contents greater than a higher limit, some water molecules may not
be held rigidly on the clay surface and plasticity is reduced (Grimshaw, 1971). The
minimum amount of water necessary to make the clay plastic is the ‘plastic limit’
(PL) and the amount of water beyond which the clay deforms under its own weight is
the ‘liquid limit’ (LL). The plastic and liquid limits often known also as Atterberg
limits, refer to the optimum interval for water content, within which the clay is
plastic and is referred to as ‘plasticity index’ (PI). By definition, it follows that:
PI ¼ LL 2 PL. Clay minerals can be classified according to their liquid limit and the
plasticity index (Fig. 14) (Bain, 1971). Plasticity measurements are important for assess-
ment of structural or pottery clays (clays used for bricks or pottery) (Bain & Highley,
1978), for clays used in engineering applications (Christaras, 1991), in foundry clays,
and for clays used as animal litters (Scott, 1990b).

3.7. Organic reactions

Clays are known to interact with various organic compounds to form various complexes,
via several types of chemical bonds between the oxygen surfaces and the organic mol-
ecules (Mortland, 1970). This reaction, which converts clay surfaces from hydrophilic to
hydrophobic, yields organoclays (Lagaly et al., 2006; de Paiva et al., 2008). The inter-
calation of certain organic molecules is used in the identification of clay minerals. Hence
 Industrial clays 363

180
Sepiolite and
160 palygorskite
Trace of
Casagrande
A line
140

120
Plastic Limit (PL)

Ca-smectite
100

80 Halloysite

60
Kaolinite
40
Illite
Na-smectite
20
Plastic
Kaolins
0
10 20 50 100 200 500 900
Plasticity Index (PI)

Fig. 14. Chart for identification of clays using plastic limit and plasticity index (modified from Bain, 1971).

ethylene glycol or glycerol is used to help identify swelling clay minerals such as smec-
tite and vermiculite (MacEwan & Wilson, 1984; Wilson, 1987) and formamide is used
to differentiate between kaolinite from halloysite (Churchman et al., 1984). Intercala-
tion of amines has also been used to determine layer charge and charge heterogeneity
of smectites and vermiculites (Lagaly, 1981, 1994), although the validity of this appli-
cation has been questioned, at least for smectites (Laird, 1994; Christidis, 2008b).
Organic molecules may be taken up by clay via ion exchange, whereby alkylammonium
cations replace interlayer cations (Lagaly, 1981, 1994), via replacement of interlayer
water molecules by polar organic molecules such as ethylene-glycol (MacEwan &
Wilson, 1984) or via grafting reactions, whereby covalent bonds form between reactive
surface groups and organic molecules (Lagaly et al., 2006; de Paiva et al., 2008).
Organoclays find important application in the formulation of polymer nanocom-
posites, because they improve mechanical, physical and chemical properties of the
polymer matrix and may reduce cost. Clay particles have to be exfoliated successfully
364 G. E. Christidis

– – –
– –
+ + +
+ + + + + +
– – – –
Monolayer d001 = 13.6 Å Bilayer d001 = 17.6 Å
– – – – –
– – – + + + + +
+ + +
+ + + + + + + +
– – – – – – – –
Pseudotrilayer d001 = 22 Å Paraffin type

Fig. 15. Schematic configuration of the alkylammonium ions in the interlayer space of smectite.

in order to obtain proper organophilization (de Paiva et al., 2008). Although interaction
with organic compounds has been studied thoroughly, mainly for 2:1 clay minerals,
grafting of organic compounds has been studied for kaolinite (Gardolinski & Lagaly,
2005) and halloysite (Breen et al., 2002). Other than for formulation of polymer
nanocomposites, organoclays have been used as adsorbents, rheological control
agents, paints, cosmetics, personal-care products, oil-well drilling fluids etc. (de Paiva
et al., 2008).
Alkylammonium ions in expandable clay minerals may form monolayers, bilayers,
pseudotrimolecular layers or paraffin-type structures (Fig. 15), depending on the layer Q2
charge of the clay mineral and the chain length of the organic ion (Lagaly, 1981,
1994). Due to their hydrophobic surfaces the organoclays display increasing selectivity
for organic molecules compared to their hydrophilic counterparts, although natural clays
also adsorb various organic molecules from vapours or aqueous solutions (Laird et al.,
1992; Sawhney, 1996). The organoclays may behave either as organophilic clays or as
adsorptive clays, depending on the chain length of the intercalated alkylammonium ions
(Boyd & Jaynes, 1994). The organic phases of the organophilic clays behave as partition
media, whereas the adsorptive clays behave as typical solid adsorptives (Boyd & Jaynes,
1994). The importance of this behaviour is that organoclays can be used for environ-
mental applications such as removal of organic contaminants from soils, natural
waters etc., or as components of composite landfill liners for retention of organic
contaminants (Czurda, 1993).

4. Kaolins

4.1. Introduction
Kaolin is a clay consisting of substantially pure kaolinite or related clay minerals
(halloysite, dickite, nacrite), which is naturally or can be beneficiated to be white or
nearly white, will fire white or nearly white and it is amenable to beneficiation by
 Industrial clays 365

known methods to be suitable for use in whiteware, paper, rubber, paint and similar uses
(Murray, 1976). The term kaolin is used both as a clay rock and as a name of a mineral
group, the latter comprising kaolinite, halloysite, dickite and nacrite. Kaolin-group min-
erals are less reactive than smectite or palygorskite. There are several terms which
describe kaolin-rich clays. The term ‘china clay’ is used mainly in Europe, interchange-
ably with kaolin. It usually refers to clays with well ordered kaolinite, whereas kaolin is
more generic and comprises all types of kaolinitic clays. ‘Tonsteins’ are non-marine,
generally kaolinitic layers derived from in situ alteration of air-fall volcanic ash,
which are usually associated with coal deposits (Bohor & Triplehorn, 1993). ‘Soft
kaolins’ contain coarse-grained kaolinite often forming booklets, with a high Hinckley
Index. ‘Hard kaolins’ contain fine-grained kaolinite with a low Hinckley Index. ‘Ball
clays’ are highly plastic sedimentary kaolinitic clays deposited mainly in fresh water,
which are often associated with lignite strata. ‘Fireclays’ are non-marine sedimentary
clays associated with higher-rank coal strata (usually sub-bituminous coals), which
have a fusion point above PCE 15 (14248C or 25958F). In the UK, these are called
‘seatearths’ and in the USA ‘underclays’ because they often underlie coal seams.
‘Flint clays’ are smooth, tough, flint-like kaolins with conchoidal fracture, which fre-
quently contain Al-oxyhydroxides. The terms ‘grog’, ‘chamotte’ and ‘molochite’ are
synonymous and refer to previously calcined kaolinitic clays, which form a rigid skel-
eton for refractory bricks. The term molochite is used only for calcined china clays.

4.2. Genesis – geological characteristics of kaolin deposits

Kaolin deposits are classified either as primary or secondary (Kuzvart, 1984; Bristow,
1987). Primary kaolins develop in situ by the alteration of aluminosilicates such as
feldspars to minerals of the kaolin group. A schematic reaction for the alteration of
K-feldspar is shown below:

4KAlSi3 O8 þ 4Hþ þ 2H2 O ! Al4 Si4 O10 (OH)8 þ 8SiO2 þ 4Kþ

K-feldspar kaolinite

The alteration results from surface weathering, circulation of groundwater below the
surface or hydrothermal activity or a combination of these processes (Murray & Keller,
1993). Secondary kaolins were eroded, transported and deposited as sediments in beds
or lenses and are associated with other sedimentary rocks. In most sedimentary kaolins,
kaolinite formed in the source area and was transported and deposited to the present site
(Murray & Keller, 1993).

4.2.1. Primary kaolins

Depending on the mode of origin, primary kaolins can be classified into ‘weathering or
residual’ and ‘hydrothermal’. Weathering of aluminosilicate minerals in various rock
types to kaolinite and halloysite is a common process and kaolin minerals are very
common constituents of subtropical and tropical soil profiles, which may be tens of
metres thick (Gilkes et al., 1980; Keller et al., 1980; Dixon, 1989; Wilson, 1999;
Hart et al., 2003; Arslan et al., 2006; Sousa et al. 2007). Kaolin minerals can form
366 G. E. Christidis

from feldspathic rock types with compositions ranging from acidic (granites-rhyolites
and related rocks) to basic (dolerites, basalts and related rocks) (Dixon, 1989; Joussein
et al., 2005; Sousa et al., 2007). Kaolin group minerals form via weathering of feldspars,
ferromagnesian minerals and micas (Dong et al., 1998; Jeong, 2000; Joussein et al.,
2005) by a process similar to lateritization. Precipitation of kaolin minerals requires
acid pH with moderate silica activity and small amounts of base cations (Dixon,
1989). Both soil kaolinites and halloysites contain structural Fe, which contributes to
greater disorder (Wilson, 1999). Soil halloysite may coexist with or transform to kaoli-
nite (Dong et al., 1998; Papoulis et al., 2004a; Joussein et al., 2005).
Although kaolin minerals form readily in soils, deposits of primary kaolins which
arise through weathering processes are less common because they contain abundant
contaminating Fe oxides and other pigments, which render these soils unsuitable for
large-added-value industrial-kaolin products (Pickering & Murray, 1994). Nevertheless
kaolin-rich soil profiles often contain a zone with white pure kaolin, which forms from
in situ deferrugination under reducing acidic conditions facilitated by organic activity
(Sousa et al., 2007). In these cases, important residual kaolin deposits may form,
which may consist of kaolinite and/or halloysite. Halloysite is favoured in the water-
saturated zone close to the bed rock but converts to kaolinite when exposed to alternat-
ing wetting and drying conditions. The transition of halloysite to kaolinite is gradual and
may occur in distinct steps (Papoulis et al., 2004a). Residual kaolin deposits may gradu-
ally convert to bauxites if leaching of silica is extensive; in this case they contain
Al-oxyhydroxides, namely boehmite or diaspore.
Hydrothermal alteration of feldspathic rocks is an important mechanism for the for-
mation of primary kaolins in Europe, Asia Minor, South America and Oceania (Kuzvart,
1984; Bristow, 1987, 1993; Harvey & Murray, 1997; Dill et al., 1997; Ece & Schroeder,
2007). The mineralogical composition of hydrothermal kaolins depends on the type of
the parent rock, the mechanism of alteration (single or multiple alteration steps invol-
ving different fluids, low or high temperature) and the composition of the hydrothermal
fluids. Alteration is controlled by structural features such as faults or joints, which may
reflect regional stress fields or volume loss due to contraction during cooling and can be
pervasive over large masses of the host rock or located close to the structural feature
(Fig. 16). The heat source which sustains the hydrothermal circuit also varies from mag-
matic, associated with the cooling of magmatic rocks, to radiogenic from the radioactive
decay of uranium (Bristow, 1993). Kaolin morphology seems to be controlled by the
precursor mineral. Fine-grained kaolin forms from dissolution of feldspar, whereas ver-
miform kaolin forms from expansion of pre-existing micas (Psyrillos et al., 1999).
A special case of hydrothermal kaolins is the so-called ‘solfatara-type’ kaolins, which
frequently occurs in areas of recent volcanism (Bristow, 1987). Solfatara-type kaolins
have been reported in the Mediterranean (Bristow, 1987; Christidis & Marcopoulos,
1995; Ece & Schroeder, 2007) and in South America (Dill et al., 1997). These deposits
contain abundant opal-CT, which is difficult to separate from kaolin minerals due to
their fine-grained size, and sulphates such as alunite. The main use of these kaolins is
the manufacture of white cements. Most of the large hydrothermal kaolin deposits
have been affected by weathering after kaolin formation, which has modified both
mineralogy (e.g. conversion of halloysite to kaolinite, oxidation of pyrite) and texture.
 Industrial clays 367

Aureole rock

Kaolinized granite

Fresh granite
Fresh granite

Fig. 16. Schematic cross section of the kaolinized granites in Cornwall, UK, which yielded the so-called
‘china clay’ deposits (modified after Bristow, 1993).

4.2.2. Secondary kaolins

Deposits of secondary kaolins form by sedimentary processes. There are three main
types of secondary kaolins, sedimentary kaolins sensu stricto (s.s.), kaolinitic sands
and specific kaolinitic clays (ball clays, fireclays and flint clays) (Bristow, 1987). For-
mation of secondary kaolins involve mainly genesis of kaolin minerals elsewhere and
deposition in a sedimentary basin as clay or post depositional alteration of clastic
grains of feldspar in arkosic sands. The latter deposits include mainly kaolinitic
sands. Finally, sedimentary kaolin deposits may form by dissolution of argillaceous car-
bonates, which leaves a kaolinitic insoluble residue (Pickering & Murray, 1994).
Sedimentary kaolins s.s. are often rich in kaolinite, which is usually transported from
weathering profiles. Parent rocks are usually feldspathic rocks, similar to the residual
kaolins. The next step involves weathering of the soil profiles and transportation to
the sedimentary basin. The nature of the parent rock and the depositional environment
determine the particle-size characteristics of the kaolins. In the case of the Georgia
kaolins the soft kaolins formed from weathering of granites and gneisses and the detritus
(kaolinite, quartz and metahalloysite) were deposited in fresh water, which facilitated
edge-to-face particle association (Dombrowski, 1993). The hard kaolins originate
from weathering of finer-grained rocks such as phyllites and deposition of the detrital
material in saline to brackish waters, which caused face-to-face particle associations.
The different kaolin particle associations have also affected post-depositional alter-
ation. Face-to-face flocculation constrained groundwater movement and restricted
post-depositional alteration. In contrast, edge-to-face flocculation facilitated diagenetic
alteration including the growth of large vermicular kaolin booklets (Dombrowski,
1993).
368 G. E. Christidis

‘Kaolinitic sands’ are a main source of kaolin in many parts of the world. These
deposits contain up to 20% kaolinite and may form either in situ, via alteration of detrital
feldspar or mica to kaolin minerals, or, by transportation and deposition of detrital kao-
linite and sand grains. Kaolin minerals may undergo diagenetic transformations after
burial. Significant kaolinitic sand deposits are common in Europe (Bristow, 1987).
Ball clays are plastic non-marine sedimentary kaolins with organic matter which
often are linked with lignite beds. The higher grades of ball clays fire to white or
nearly white colour in an oxidizing atmosphere. Ball clays are usually interbedded
with sands, silts and silty clays and the source of kaolinite is the deep weathering of
feldspathic igneous or metamorphic rocks (Patterson & Murray, 1984; Bristow,
1987). They consist of variable amounts of kaolinite (which usually has a large
number of structural defects), illite and quartz, with variable amounts of organic
matter. In some deposits kaolinite is well ordered (e.g. Bovey basin, UK). Fireclays
are non-marine kaolins which contain disordered kaolinite, smaller amounts of mica,
quartz, ironstone nodules and carbonaceous matter. They are found in coal-bearing
strata, often underlying coal seams with higher coal rank than lignite. Fireclays are con-
sidered to form by diagenesis of ball clays, because the latter have many similarities
with fireclays and are not older than the Tertiary (Bristow, 1987). In some cases fireclays
have been considered as palaeosoils and represent periods when water was sufficiently
shallow to allow colonization by land plants. Flint clays are hard, Al-rich, Fe-poor, non-
slaking rocks with conchoidal fracture, which frequently contain boehmite or diaspore
and therefore are used mainly in refractories. They contain well ordered kaolinite and
form either by intensive leaching and recrystallization of kaolin precursor rocks or by
low-grade regional metamorphism of ball clays at P-T conditions which do not allow
the formation of pyrophyllite (Bristow, 1987). Flint clays are generally considered to
be allochthonous (Loughnan, 1978).

4.3. Assessment of kaolin deposits

4.3.1. Preliminary assessment
The preliminary assessment of kaolins involves recognition in the field, determination of
mineralogical composition and colour assessment. Recognition of the kaolin in the field
is based on the white-off white colour of the raw material and the sticky feeling to the
tongue. Residual kaolins, especially those in tropical and sub-tropical areas may contain
abundant Fe oxides which impart red-pink colourations and are unsuitable for large-
added-value applications (Bloodworth et al., 1993). These coloured impurities are
fine-grained and usually cannot be removed during beneficiation. The original field rec-
ognition of the kaolins is confirmed by means of X-ray diffraction (XRD) on whole-rock
random powder samples or certain size fractions. X-ray diffraction is also used to deter-
mine the mineralogical composition of the grit content (usually coarser than 325 mesh or
44 mm according to Murray, 2007), to distinguish between halloysite and kaolinite
(Churchman et al., 1984; Joussein et al. 2005) and to recognize dickite and nacrite
(Brindley, 1984; Moore & Reynolds, 1997). Particle-size distribution is determined
using wet sieving for the fractions coarser than 63 mm and by laser diffraction or
with X-ray sedigraph for the subsieve residue.
 Industrial clays 369

Common industrial practice is to use various indices to determine the degree of

‘crystal disorder’ of kaolinite crystals (Pruett & Pickering, 2006), previously expressed
erroneously as “crystallinity” (Guggenheim et al., 2002). These indices are determined
on randomly ordered powders by XRD by examination of the diffraction maxima attrib-
uted to prismatic reflections in the interval 19– 2782u (Fig. 17). The most commonly
used index is the Hinckley Index (HI) (Hinckley, 1963). Alternatives include the
Aparicio-Galán-Ferrell Index (AGFI) (Galán et al., 2006) and the Lietard test
(Lietard, 1977) (Fig. 17). All indices are affected by the presence of mineral impurities
such as smectite, quartz, feldspar, halloysite and amorphous gels (Galán et al., 2006).
Lower values for the indices indicate a smaller degree of ordering. The degree of order-
ing of kaolinite may affect the potential end-use of the kaolin. This is because generally,
but not always, well ordered kaolinites are coarser grained with better crystal mor-
phology than more poorly ordered varieties (Bloodworth et al., 1993).
The ‘kaolin grade’ refers to the amount of kaolin minerals present in the kaolin. Grade
can be determined either directly by quantitative determination of kaolin minerals by
XRD on randomly ordered samples (Bish & Ploetze, 2011, this volume) or indirectly
by thermogravimetric (TG) analysis. The latter method is based on the fact that pure
kaolin minerals loose 14 wt.% due to dehydroxylation when heated at 500 –6008C.
The validity of this method is restricted by phases with structural hydroxyls, which
dehydroxylate at the same temperature range, such as illite or trans-vacant smectite,
which are often present in kaolins.
Preliminary colour assessment is carried out using the Munsell Soil Colour Chart on
unprocessed dry samples (Munsell Color Company, 1954). The colour charts give sep-
arate notations for hue, value and chroma. Determination of colour parameters of crude
kaolin samples using the CIELab system (Fig. 12) is usually more useful for the end
products after processing. However it can also be used for the crude samples as a
rough estimate of the colour properties of the raw materials (Murray, 2007).

4.3.2. Next stage of kaolin assessment

The aforementioned tests, though important for primary characterization of the raw
kaolin materials, provide no information about the suitability of the material in question

Aparicio-Galan-Ferrel Lietard test R2

Hincley Index (HI)
Index (AGFI)
A+B 002 IA + IB 002 131
HI = 110 AGFI =
2IC
131
At 1 / 2(K1 + K 2 )– K
110 R2 =
020 111 [1 / 3(K1 + K 2 + K )]
020 A 111 003
B Κ2
021 Κ1
At Q CA B Κ

20 22 24 26 20 22 24 26 37.5 39.5
°2θ (CuKα) °2θ (CuKα) °2θ (CuKα)

Fig. 17. Indices used to determine the degree of crystal order in kaolinite.
370 G. E. Christidis

for specific industrial applications. In the case of kaolins which are assessed as fillers
and coatings, the next step for assessment usually involves a series of beneficiation
tests, which are followed by characterization of the end products after each test. The
methods used for beneficiation are roughly the same in the various areas, although
the sequence in which they are used may vary. Typical tests include dispersion and
degritting of the kaolin in aqueous suspensions, fractionation into different size frac-
tions using different hydrocyclones and delamination via vigorous forms of agitation
(Bloodworth et al., 1993; Murray, 2007). Colouring impurities such as Fe oxides, Fe
oxyhydroxides and TiO2 may be removed by high-density magnetic separation
(Iannicelli, 1976), chemical bleaching using reducing agents such as Na dithionate
(Bloodworth et al., 1993; Murray et al., 2007) and froth flotation or selective floccula-
tion (Murray, 2007).
Kaolins for ceramics are assessed by a series of physical properties such as plasticity
(determination of Atterberg limits), green strength, dry strength, volume and linear
drying shrinkage and modulus of rapture. Chemical composition is also important for
ceramic grades. Physical properties are determined on wet and dry test specimens
with specific dimensions which are prepared from the raw materials using extruders.
Specimens with 80% relative humidity usually have a smaller modulus of rapture
than dried specimens (Bloodworth et al., 1993).
The end products from each beneficiation or firing test are characterized with a series
of analytical techniques. Assessment involves determination of mineralogical, physical
and chemical properties which are related to the potential applications of the clay. End
products with large proportions of particles of ,2 mm may suggest further investigation
for use in the paper industry, whereas ceramic applications may require products domi-
nated by coarser particles (Bloodworth et al., 1993). The main technique for determin-
ing the mineralogical composition of the end products is XRD, while examination by
binocular microscope will help to determine the heavy minerals present. The mineralo-
gical composition is usually determined in the various fractions obtained from hydrocy-
clone separation and confirms the efficiency of the beneficiation method. It may also
provide an indication of the presence of abrasive mineral impurities such as quartz.
The abundance of kaolin minerals can be confirmed either by quantitative XRD analysis
or by TG analysis. Other bulk properties, which are determined at this stage, are particle-
size distribution using laser diffraction and specific surface area using nitrogen adsorp-
tion, according to the BET method. The particle-size distribution determines the
suitability of kaolin slurries as coatings or fillers (Fig. 18).
The white colour is an important prerequisite for most kaolin applications. Beneficia-
tion usually improves lightness (L ) and reduces redness (a ) and yellowness (b ),
because it removes impurities which impart colour to kaolins. An increase of L fol-
lowed by decrease of a and b compared to the crude kaolin indicates efficient
beneficiation. In fired products the colour properties are of vital importance. The
colour properties of sedimentary kaolins containing fine-grained organic matter such
as ball clays, usually improve after firing due to oxidation of organic matter. In contrast,
other than some mineralogical transformations (e.g. conversion of goethite to hematite)
the influence of Fe oxides on colour is not affected significantly by firing. For most
 Industrial clays 371

100

% finer by weight
80
Fillers
60 Plates

325 mesh
Coatings
40

20 Stacks

0
100 50 20 10 5 2 1 0.5 0.2 0.1
Equivalent spherical diameter (µm)
Fig. 18. Particle-size distribution curves of paper filler and coating grades of kaolin (modified from Pickering
& Murray, 1994).

kaolin products, assessment of colour includes determination of brightness at a specific

wavelength (usually at 457 nm). In general, a minimum brightness of 80% is required
for filler kaolins and 85% for coating kaolins (Pickering & Murray, 1994).
The rheological properties of kaolin slurries are important for the assessment for
large-added-value applications such as paper making and slip-casted ceramic bodies.
The viscosity of kaolin suspensions depends on the particle size and shape, specific
surface area and surface charge and impurities and is determined using Brookfield
and Hercules viscometers (Pickering & Murray, 1994; Murray, 2007). The parameters
which are assessed are ‘flowability’, which is the solids content at which the clay slurry
begins to flow, ‘deflocculant demand’, which is the minimum amount of deflocculant
required to obtain a slurry with minimum viscosity and the ‘viscosity concentration’,
which is the solids concentration of a fully deflocculated clay with a viscosity of
5 poise at 228C (Bloodworth et al., 1993). The flow properties of kaolin suspensions
are dictated by certain specifications. For paper coating applications 80% of the
kaolin particles are usually ,2 mm, whereas for paper-filler applications, the kaolins
are coarser, ranging between 40 and 60% of which are ,2 mm.
Assessment of kaolins for ceramics is carried out by determination of physical prop-
erties of fired kaolins via a series of tests on ceramic specimens prepared in extruders.
The specimens are fired in the temperature range 900 – 13008C and the properties deter-
mined are fired shrinkage, porosity, specific gravity, firing colour and modulus of
rapture. The fired properties are presented as vitrification curves in which the evolution
of each property is shown as a function of temperature.

4.4. Mining and processing of kaolins

The mining methods used depend on the type of the kaolin deposit. Primary hydrother-
mal kaolins such as the china clay deposits of Cornwall are mined with high-pressure
water monitors, which break down the friable altered granite and release kaolinite
into suspension. The resulting clay slurry flows to the lowest part of the mines, the
372 G. E. Christidis

coarser particles settle, and the slurry is pumped out to the processing plant. Secondary
kaolin deposits are mined by opencast methods after removal of the overburden
using bulldozers, scrapers and excavators. The kaolin may be carried to the
processing plant either in the raw state by haul tracks or is dispersed in slurries, the
coarser particles are allowed to settle and the slurry is pumped to the processing units
(Pruett & Pickering, 2006).
Processing of kaolin concentrates is carried out either in the dry or the wet state
(Pickering & Murray, 1994; Pruett & Pickering, 2006). Dry processing is performed
in refractory and ceramic applications, in which purity handling and appearance are
less stringent. The dry process involves mainly drying, grinding and air floating,
which removes most of the grit. The wet processing includes a variety of different ben-
eficiation techniques such as delamination, magnetic separation, selective flocculation,
flotation and leaching, which remove impurities and decrease kaolin content. Delamina-
tion breaks up kaolinite books and renders the material suitable for paper coating
(Fig. 18). Magnetic separation removes magnetic colouring minerals such as hematite,
Fe-rich anatase, ilmenite, magnetite and biotite. Froth flotation is used for production of
coating kaolin with high brightness and removes mainly Fe-rich anatase. The processing
flow sheet may differ in the various areas because of the different properties of the raw
material. Figure 19 shows a typical processing flow sheet followed during the processing
of Georgia kaolins. The processing of Brazilian kaolins does not include flotation
and surface treatments (Murray et al., 2007), whilst the processing of Cornish china
clays does not include magnetic separation and leaching with reducing compounds
(Bloodworth et al., 1993).

4.5. Distribution of kaolin deposits

Important deposits of primary kaolin occur in Europe, South America, Asia and New
Zealand. The most famous hydrothermal kaolin deposits in the world occur in Cornwall
and Devon, UK. The deposits formed from intense kaolinization of feldspar and
mica present in granites (Bristow, 1993; Psyrillos et al., 1998, 1999). The precursor
mineral dictated the morphology of the kaolin crystals (Psyrillos et al., 1999). The
kaolin bodies have funnel or trough-like forms and their grade ranges between 10 and
20% kaolinite (Fig 16). Hydrothermal alteration involved several steps and kaoliniza-
tion is considered to have taken place at a later stage from low-temperature meteoric
waters and high-salinity brines heated by radioactive decay of uranium (Bristow,
1993; Psyrillos et al., 1998). The hydrothermal flow occurred along pre-existing hydro-
thermal veins and fractures, formed during a previous alteration stage, which caused
softening of the parent rock (Psyrillos et al., 1998). After their formation, the kaolin
deposits were affected by deep weathering. Kaolinization is considered to still be
active today.
In New Zealand, a Plio-Pleiostocene, ceramic-grade halloysite deposit occurs in the
North Island. The deposit was formed by hydrothermal alteration at the expense of a
rhyolitic flow and has been affected by subsequent weathering (Harvey & Murray,
1993). It contains 50% halloysite and 50% silica polymorphs (quartz and opal-CT).
 Industrial clays 373

DRILLING

STRIPPING

MINING

DEGRITTING

STORAGE AND BLENDING

FRACTIONATION AND PARTICLE-SIZE SEPARATION

SELECTIVE MAGNETIC DELAMINATION FLOTATION

FLOCCULATION SEPARATION

LEACHING SURFACE
TREATMENT
DEWATERING

DEFLOCCULATION
SLURRY
DRYING (HIGH SOLIDS)
SPRAY AND/OR
DRUM DRYING
PARTICLE SEPARATION

CALCINATION
BAGGING AND LOADING

Fig. 19. Schematic flowsheet of wet processing of kaolin (modified after Murray, 2007).

In China there are many primary hydrothermal kaolins, the most important of which are
the hydrothermal deposits at Suzhou, west of Sanghai, eastern China, and Fujian and
Zhanjiang in southeastern China (Wilson, 2004; Murray, 2007). The kaolin deposits
formed from hydrothermal alteration of Jurassic volcanics and the kaolinized zones
resemble those of the St Austell granite, Cornwall in places (Wilson, 2004). The depos-
its consist of kaolinite, halloysite and quartz, with minor smectite, sericite, alunite and
pyrite. Finally, in western Turkey, a large number of small hydrothermal kaolins with
kaolinite and halloysite, are ‘solfatara-type’ deposits and contain alunite and silica poly-
morphs as minor phases (Ece & Schroeder, 2007). The deposits formed from alteration
of acidic volcanics. Similar deposits occur along the south Aegean volcanic arc, Greece,
in the islands of Milos, and Kos (Christidis & Marcopoulos, 1995; Papoulis et al.,
2004b).
Important resources of residual primary kaolins occur worldwide. The most important
example is Brazil, which in 2005 accounted for 10% of the world’s kaolin production
(Wilson et al., 2006). The most important residual deposits occur in the Amazon area
in northern Brazil close to Jari River, a tributary of the Amazon and the Capim
River. The kaolins derived from weathered granitic rocks of the Guyana Shield and
were deposited in a deltaic environment. The material deposited was leached intensively
374 G. E. Christidis

after deposition. The Rio Capim kaolin is more complex and consists of six kaolin facies
(Sousa et al., 2007). Lateritization originally produced Fe-rich soft kaolin with well
ordered kaolinite, but subsequent diagenesis caused ferrugination and yielded flint
kaolins on the top of the kaolin sequence. In the Ukraine, important residual kaolin
deposits occur at Prosyanovski, north of the sea of Azof and at Glhovetski, 200 km
southwest of Kiev (Pickering & Murray, 1994). The deposits formed from deep weath-
ering of granites and gneisses mainly along the main fracture zones. In the Belitung and
Pangka Islands, Indonesia residual kaolins formed by weathering alteration of porphyri-
tic biotite granites to kaolinite, halloysite and smectite (Murray et al., 1978). Finally, in
China, major residual kaolin deposits derived from weathering of granites and acidic
volcanics occur in Fujian and Guangdong Provinces in the southwesten part of the
country (Wilson, 2004).
The most important secondary sedimentary s.s. kaolins are the deposits in Georgia
and South Carolina, USA. The deposits are Late Cretaceous – Early Tertiary in age
and contain both soft and hard kaolins. Soft kaolins are coarser grained and often
contain stacks of vermicular kaolinite, whereas hard kaolins are fine-grained with
face-to-face particle contacts (Murray, 2007). Both kaolin types formed from weather-
ing of basement rocks to kaolinite and/or halloysite and transportation of kaolin detritus
to the sedimentary basins. The differences in particle size and morphology of the soft
and the hard kaolins are attributed to genetic, depositional and post-deposition con-
straints. The soft kaolins formed via weathering of feldspars and micas of orthogneisses
and granites, transportation of the detritus via fluvial processes and deposition in deltaic
and estuarine environments. i.e. in fresh waters (Dombrowski, 1993). Post-depositional
leaching, oxidation and diagenesis modified the original mineralogical features. In con-
trast, the hard kaolins formed via weathering of phyllites and fine-grained schists and
were deposited in brackish to saline water. This environment facilitated face-to-face
flocculation, caused tight packing of kaolin particles and inhibited groundwater circula-
tion, thereby restricting post-depositional diagenetic modifications (Dombrowski,
1993). Deposits of sedimentary kaolin are mined in India, China and Australia.
The most characteristic ball clay deposits occur in the UK, the USA, Germany and the
Ukraine. The English ball clay deposits are Paleocene in age and occur in the Petrock-
stow Basin and the Bovey basin in north and south Devon respectively, and the
Wareham Basin at Dorset, UK (Bristow et al., 2002). The clays were deposited under
freshwater lacustrine, overbank or fluviatile conditions and are associated with sands,
silty clays and lignites. They have been derived from rocks which underwent intense
chemical weathering in the early Paleocene and subsequently eroded and deposited in
the basins. Post-depositional weathering and diagenesis modified the original sediments.
In the USA, important ball clay deposits of Eocene age occur mainly in Kentucky and
Tennessee, in the form of lenticular units interstratified with sand, silt and lignite
(Patterson & Murray, 1984). Their origin and depositional is similar to their English
counterparts, but the nature of their parent rocks is unknown. Both ball clays contain
fine-grained disordered kaolinite which imparts high plasticity to the clays.
Sedimentary refractory kaolinitic clays are also widespread throughout the world.
They include the fireclays and flint clays. Significant deposits of fireclays occur
 Industrial clays 375

within the Carboniferous coal measures in western Europe and the eastern USA, in the
Appalachian region and in parts of the Mississippi Valley (Bristow, 1987; Patterson &
Murray, 1984; Pickering & Murray, 1994). Early Cretaceous fireclays occur in Colorado
and Tertiary fireclays in California, Oregon and Washington (Patterson & Murray,
1984). In some areas (Kentucky, Missouri and Colorado, USA) fireclays coexist with
flint clays and semiflint clays (Patterson & Murray, 1984), suggesting that at least
some of the flint clays are underclays which have undergone intensive leaching and
recrystallization, so that kaolinite has grown to form large crystals (Bristow, 1987).
Hence the underclays in Kentucky differ in terms of plasticity and degree of ordering
of kaolinite; fireclays are plastic and contain poorly ordered kaolinite whereas filnt
clays are hard with well ordered kaolinite. In Missouri, however, the fire clays are
not clearly associated with overlying coal beds. In Colorado, plastic fireclays contain
large kaolinite crystals (Patterson & Murray, 1984). In the UK, fireclays are Carbonifer-
ous and are associated with the Coal Measures. They occur in north and central England,
Wales, Scotland and Ireland. Today they are not used as refractory clays but as brick
clays (see below). Refractory clays are also mined in Europe (Germany, France,
Italy, Czech Republic, Poland, Hungary and Russia), Asia (China, India, Japan) Argen-
tina, Mexico and Australia.

4.6. Applications of kaolins

Kaolins are very important industrial clay minerals, the majority of which consist of
kaolinite. They have many applications because of the physical and chemical properties
of kaolinite (Table 5). In summary, kaolinite has a small layer charge and small CEC and
specific surface area. These properties vary according to the degree of structural order of
the kaolinite, which has characteristic pseudo-hexagonal particle shape (Fig. 4). It also
develops low viscosity when dispersed in water and it has a white colour. Of the remain-
ing minerals of the kaolin group, only halloysite forms deposits of economic signifi-
cance. They are used in the natural state, partially calcined (heated at 650 – 7508C) or
fully calcined (soaked at 1000– 10508C) (Bundy, 1993). The most important appli-
cations of kaolin are summarized in the following paragraphs.
The paper industry consumes large amounts of kaolin every year for both filling and
coating purposes. As a filler, the kaolin is mixed with the cellulose fibres in wood pulp to

Table 5. Important properties of kaolin (modified after Murray, 2007).

White or nearly white colour
Chemically inert for pH 4 – 9
Fine particle size
Soft and non-abrasive
Platy, with large basal surface dimensions relative to lateral dimensions
Hydrophilic with high dispersibility in water
Good covering and hiding power when used as pigment or extender in coatings
Plastic, refractory, which fires to a white or nearly white colour
Low heat and electrical conductivity
Very low layer charge and specific surface area compared with other clay minerals
Low viscosity. Some kaolins flow readily at 70% solids
376 G. E. Christidis

reduce cost and improve printing characteristics. Important properties of filler kaolins
are particle size, white colour and low abrasiveness. When used as a coating pigment,
it enhances the surface properties of the paper, such as brightness, smoothness,
opacity and gloss and improves the printability by increasing ink receptivity. As well
as particle size and white colour, the rheology is of particular importance for coating
applications (Bloodworth et al. 1993; Bundy, 1993; Murray, 2007). In general,
coating kaolins have finer particle size (Fig. 18), better brighness and less yellowness
than filler kaolins. Lightweight, coated papers may contain up to 40% of kaolin, as
both filler and coating. Major competitors of kaolin fillers are calcium carbonate
(with the move of the paper industry to use of acid-free paper) and talc.
Kaolins are used as extender pigments in interior water-based and in exterior oil-
based paints. The addition of kaolin reduces the cost and enhances the opacity of the
paint. Moreover, it contributes to the suspension viscosity and functions as a suspension
aid, which prevents pigment settling. Both calcined and delaminated kaolins are used,
with the former imparting greater opacity and toughness to water-based paints
(Bundy, 1993). Delaminated kaolins give a smooth surface to paint films and a
greater sheen. Important parameters for use in paints are particle size and brightness.
Kaolin was used almost exclusively in the production of ceramics until the end of the
19th century. Kolins forms an important constituent of several body formulations
(Table 6). In these formulations ball clay is added to increase plasticity and green and
dry compression strength because china clays have low plasticity and low dry com-
pression strength. This is due to the larger particle size of the china clays; a finer particle
size induces greater dry compression strength and plasticity. In general, the dry com-
pression strength of china clays and ball clays is greater than their green compression
strength. After firing at high temperatures, important properties are modulus of
rapture, firing colour, shrinkage, porosity and bullk density. The modulus of rapture
depends on the particle size of the kaolin and the firing colour depends on the colouring
admixtures (Fe oxides, Mn oxides) because kaolinite is white. Finally, for the manufac-
ture of sanitary ware, the rheological properties of the casting slip (viscosity, defloccu-
lant demand) are critical. Reviews of the requirements for kaolins in ceramics were
given by Jepson (1984) and Murray (2007).
Kaolins are used as pigment fillers and extenders in rubber and plastic (vinyls and
polyesters) to reduce the cost and to reinforce the structure. In general, the finer the par-
ticle size of kaolin the better the reinforcement of physical properties of the end pro-
ducts. The hard kaolins are so-called because of the greater stiffness they impart to

Table 6. Typical body formulations for whiteware ceramics (from Jepson, 1984).
Product China clay (%) Ball clay (%) Flux1 (%) Quartz2 (%) Others (%)
Hard porcelain 50 –55 0 15 – 25 20 – 30 –
Soft porcelain 40 10 20 – 30 20 – 30 –
Bone china 25 0 25 0 50 (bone ash)
Vitreous sanitaryware 28 24 18 30 0 – 3 talc
Earthenware 25 25 10 – 20 30 – 40 –
Lime wall tiles 25 25 0 40 10 (limestone)
1
Usually K-feldspar or nepheline syenite
2
Silica sand, calcined sand or flint
 Industrial clays 377

the rubber products, compared to the soft kaolins. Kaolin is the main inorganic pigment
used in ink. It improves ink holdout, and extends both coloured and white pigments
(Bundy, 1993). The particle size of kaolin extenders in ink ranges between 0.2 and
0.5 mm.
Kaolin is used for the synthesis of industrial minerals used as ‘catalyst carriers’. An
important application is the manufacture of ceramic monoliths of the ‘catalytic conver-
ters’ in the exhaust systems of automobiles. The catalytic converters consist of cordierite
which is produced synthetically from calcined kaolin, calcined talc, alumina and hydrous
kaolin (Murray, 2007). Moreover, kaolin is used for the synthesis of FAU-type zeolites
(zeolite-X and Y) and zeolite-A, by the hydrogel process. Zeolite A is used as a water soft-
ener and deflocculant in detergents, while zeolite Y is used as a catalyst carrier in the
‘catalytic cracking’ of petroleum. Zeolite-X is used as a molecular sieve and adsorbent.
Finally, kaolins of lower quality, such as the ‘solfatara-type’ kaolins, are used in the man-
ufacture of ‘white cements’. These kaolins contain abundant – SO24 – , usually in the form
of alunite and can then replace gypsum, which is used as a fast-setting retardant. Minor
amounts of kaolins are consumed in numerous other applications such as adhesives, sea-
lants and caulks, in cosmetics and pharmaceuticals, in crayons and chalk, in enamels,
fertilizers, pencil leads, polishing compounds, soaps and detergents, etc.

5. Bentonites

5.1. Introduction
Bentonites are clays which form at the expense of volcanic glass and which consist pre-
dominantly of smectite regardless of origin or occurrence, the physical and chemical
properties of which are dictated by this mineral (Grim & Güven, 1978; Fischer &
Schmincke, 1984). The term bentonite is not used in the same manner throughout the
world, however, and different terms are used according to the type of smectite
present. Table 7 lists the nomenclature used to describe bentonites worldwide. Note
that the term ‘fuller’s earth’ in the UK is used to describe all Ca-bentonites, whereas

Table 7. Nomenclature used to describe smectite-rich materials (modified after Anonymous, 1978).
Principal mineral Synonymous terms Regional terms
Na-smectite Na-bentonite Wyoming bentonite (USA)
Synthetic bentonite Western bentonite (USA)
Swelling bentonite Bentonite (UK)
Na-activated bentonite
Na-exchanged bentonite
Ca-smectite Ca-bentonite Southern bentonite (USA)
Sub-bentonite Texas bentonite (USA)
Non-swelling bentonite Fuller’s Earth (UK)
Mg-smectite Saponite
Stevensite
Armagosite
K-smectite Metabentonite
K-bentonite
Li-Mg-smectite Hectorite
378 G. E. Christidis

in the US it is used for clays with large absorbent capacity (Murray, 2007). Moreover,
the term ‘metabentonite’ or ‘K-bentonite’ is used for bentonite beds in which smectite
has been converted to mixed-layer I-S. The important properties of smectite clay min-
erals include crystal structure and chemical composition, small crystal size and hence
large specific surface area, type of exchangeable cations and ion exchange, hydration
and swelling, colloidal properties, dehydration and reactions with organic and inorganic
reagents (Odom, 1984). Due to these properties bentonites find a large variety of indus-
trial applications (drilling industry, foundries, iron-ore pelletization, civil engineering,
adsorbents, filtering, decolorizing and clarifying, etc.). Most recently developed appli-
cation fields include the formulation of nanocomposites (de Paiva et al., 2008; Schoon-
heydt & Bergaya, 2011).
The vast majority of bentonite deposits form from alteration of volcanic glass in sub-
aqueous, mainly seawater, environments. Large bentonite deposits are found in many
countries around the world (Grim & Güven, 1978; Elzea & Murray 1994; Murray,
2007), suggesting that bentonite formation via alteration of volcanic glass is a rather
common geological process. The volcanic nature of bentonite precursors is verified
by the presence of primary igneous minerals (b-quartz, biotite, sanidine, zircon,
apatite, ilmenite, magnetite), the presence of fresh, partially altered or pseudomorphi-
cally replaced glass shards by smectite in the bentonite and the distribution pattern of
characteristic trace elements such as the REE. However, although the role of water is
essential for bentonite formation, alteration of volcanic glass is not always observed
in water-dominated environments such as seawater even in older sedimentary strata
(Hein & Scholl, 1978; Weaver, 1989). Therefore, fresh or slightly altered glass shards
are frequently observed in marine sediments. Moreover, the presence of unaltered
volcanic glass shards is common in commercial bentonites.

5.2. Genesis-geological features of bentonite deposits

Alteration of volcanic glass may take place through vapour-phase crystallization often
associated with welding of ignimbrites, burial diagenesis, contact metamorphism,
hydrothermal activity, percolating groundwater, and in alkaline lakes or the sea floor
in marine sediments. Vapour-phase crystallization is not important for bentonite for-
mation. Alteration takes place in the presence of a fluid phase; hence bentonites form
in aqueous environments, usually shallow marine or in inland lakes. However, although
water is essential for bentonite formation, alteration of volcanic glass is not always
observed even in older volcano-sedimentary strata (Weaver, 1989). Most commercial
bentonites contain fresh glass shards.
Leaching of alkalis and a high (Mg2þ)/(Hþ) activity ratio are necessary for the for-
mation of smectite and facilitates alteration of volcanic glass to bentonite instead of zeo-
lites (Senkayi et al., 1984; Christidis, 1998). Magnesium is often supplied by the fluid
phase, especially when the parent rock has acidic composition. When the parent rock has
dacitic-andesitic composition, then an external source of Mg is not necessary (Christidis
et al., 1995). Alteration of acidic precursors results in migration of excess Si, which
either migrates downwards, e.g. the Wyoming bentonites (Grim & Güven, 1978), or
 Industrial clays 379

precipitates within the bentonite in the form of distinct opal-CT beds or as distinct opal-
CT crystals within the smectite matrix (Christidis et al., 1995; Cravero et al., 2000).
According to mass-balance calculations, large water:rock ratios, i.e. an open system,
are necessary for the formation of smectites in bentonites regardless of the parent
rock (Christidis, 1998). In contrast, lower water:rock ratios or closed systems favour
the formation of zeolites (Hay & Sheppard, 2001). High water:rock ratios are facilitated
by high permeability of the parent rocks. In contrast zeolites form under lower water:
rock ratios, dictated by lower permeability. Smectite has been reported to form from
poorly crystalline precursors, usually gels, in various environments (Zhou & Fyfe,
1989; Christidis, 2001), although the existence of such a precursor phase is not
always necessary.
The most usual smectite present in bentonites is montmorillonite. However, in several
bentonites, montmorillonite coexists with beidellite or/and nontronite (Christidis &
Dunham, 1993, 1997; Christidis, 2006). Other authigenic phases present include zeolites
(mainly clinoptilolite and/or mordenite), Si-polymorphs (opal-CT, opal-C, fine-grained
quartz), K-feldspar and Ti-oxides (anatase and brookite). Although minerals of the
kaolin group (kaolinite and halloysite) are common in bentonites, they form from altera-
tion of smectites. Similarly, palygorskite is rarely present as a minor mineral in some
bentonites, replacing smectite (Christidis, 2006).
There are three main mechanisms of formation of bentonites with economic impor-
tance: (1) in situ diagenetic alteration of volcanic glass; (2) hydrothermal alteration of
volcanic glass; and (3) formation of Mg-smectite-rich sediments in inland, saline-
alkaline lakes and sabkha environments. The latter mechanism may not include volca-
niclastic rocks (Galán & Castillo, 1984) and may explain the uncertain origin of some of
the bentonites. However, these bentonites consist mainly of trioctahedral smectites (sapo-
nite and stevensite or mixed-layer stevensite-kerolite) in small amounts, i.e. they are
usually of inferior quality and are often associated with sepiolite or/and palygorskite.
In fact they are considered to be sepiolite and/or palygorskite deposits rather than ben-
tonites. Hydrothermal alteration includes the so-called deuteric alteration caused by
gases and vapours after emplacement of the volcaniclastic rocks (Grim & Güven, 1978).
Diagenetic alteration of volcanic glass shards deposited in aqueous environments
yields beds a few cm to a few metres thick, which often form deposits extended over
large areas (Elzea & Murray, 1994). For this reason, bentonite beds formed in previous
epochs are useful for stratigraphic-correlation purposes (Kolata et al., 1996; Huff et al.,
1999; Bertog et al., 2007). Parent rocks are usually volcanic ash fall deposits derived
from sub-plinian or plinian eruptions, which are deposited mainly in seawater or
inland waters. Bed thickness and particle size of glass shards and igneous minerals
depend on the mechanism of volcanic eruption. Thicker deposits usually form near
the volcanic vent. Diagenetic alteration of volcanic glass to bentonite may begin soon
after deposition of the volcanic ash (Berry, 1999) and is facilitated by fluid flow,
which is sustained by hydraulic gradients and is controlled by permeability of the
parent rock. Temperature gradients are not considered important to support fluid flow
because thermal equilibrium is established shortly after deposition of volcanic frag-
ments. This is because the volcanic ejecta travel long distances and cool down before
380 G. E. Christidis

deposition. Typical examples of diagenetic bentonites are found in many locations

worldwide, the most characteristic ones being the Wyoming bentonites (Elzea &
Murray, 1990; Bertog et al., 2007).
Hydrothermal alteration is another important mechanism for bentonite formation.
Two types of alteration are considered, hydrothermal s.s. in which the energy source
is located at depth, associated with a magmatic body and one in which the energy
source is the cooling pyroclastic rock itself. In the former case, alteration of volcanic
rocks occurs via fluid flow around fracture zones or faults and the bentonite bodies
are elongated or elliptical, often displaying zonation parallel to the axis of the structural
feature (Ddani et al., 2005; Yildiz & Kuşcu, 2007). In the latter case, cooling of the
parent rock sustains a hydrothermal s.l. system and causes alteration. Such a hydrother-
mal system is active as long as the temperature difference between the cooling rock and
the aqueous phase (seawater or fresh water) can sustain fluid flow. In other words, high
temperatures during emplacement favour such alteration. A high emplacement tempera-
ture is common for pyroclastic flows (300 to 8008C according to Cas & Wright, 1988),
which are the parent rocks of most bentonite deposits in Milos Island, Greece. These
bentonites, which form world-class deposits, have some important differences from
their diagenetic and hydrothermal s.s. counterparts. They are stratiform like the diage-
netic bentonites but are much thicker and are spread over limited distances, dictated by
the dimensions of the pyroclastic flow (Christidis, 2008a). The total thickness of these
deposits may exceed 60 m if successive pyroclastic flows are emplaced and sub-
sequently altered (Fig. 20).

WNW ESE

Palaeo-sea floor

20 m

Green bentonite Unaltered volcanics (overburden)

Yellow bentonite
Fault
Red marl

Fig. 20. Schematic cross section of a bentonite deposit, Milos, Greece, showing the distribution of the various
bentonite horizons.
 Industrial clays 381

The formation of bentonites in inland, saline-alkaline lacustrine environments is

closely related to the formation of palygorskite and/or sepiolite. Two types of smectites
are usually recognized in these deposits: detrital dioctahedral smectites, the composition
of which depends on the source material and authigenic Mg-rich trioctahedral smectites,
mainly saponite and/or stevensite. Authigenic saponite may form from alteration of det-
rital clay minerals, such as kaolinite, smectite, chlorite or illite (Jones & Galán, 1988) or
from carbonates (Hay et al., 1986; Chahi et al., 1999) whereas stevensite forms from
direct precipitation from pore fluids along with sepiolite (Meunier, 2005). The deposits
are indicative of arid climatic conditions and usually display zonation. Detrital minerals
occur in the margin of the lake, whereas authigenic phases form in the most distal zone
of the lake-shore areas (Galán & Castillo, 1984; Velde & Meunier, 1987; Kastritis et al.,
2003). In fact the ratio authigenic/detrital clay minerals is an indication of the distance
from the source.

5.3. Assessment of bentonite deposits

5.3.1. Mineralogical-chemical investigation
Assessment begins with recognition of the bentonite in the field. A simple guide for rec-
ognition is the desiccation texture known as ‘popcorn’ texture which forms from alter-
nating wet and dry seasons (Christidis & Scott, 1993). High-quality, i.e. smectite-rich
bentonites, are compact, soft and soapy to the feel and become slippery when wet. Ben-
tonites have variable colours which depend on the nature of the parent rock, the history
of alteration and the exposure of the clay to oxidation. Bentonites derived from acidic
rocks have brighter colours in general, whereas those derived from intermediate–
basic rocks are dark green-grey when buried, becoming yellowish-green to yellow
upon exposure to weathering and near the surface. Weathering of pyrite often yields
goethite (Komadel, 2008). The colour of bentonites may also be modified by hydrother-
mal alteration resulting in the precipitation of sulphates such as alunite, gypsum or
jarosite (Christidis & Scott, 1993).
Macroscopic observation must be confirmed by a series of mineralogical techniques,
the most important of which is XRD. Analysis by XRD is carried out both on the bulk
rock (random samples) and on the clay fraction, i.e. the ,2 mm fraction (oriented
mounts). Analysis of bulk rock using random powders will provide information about
the impurities present and the abundance of the various phases (Bish & Ploetze,
2011). Moreover, additional information can be extracted about the dioctahedral or
trioctahedral nature of the smectite present. Analysis of the clay fraction can elucidate
the distribution of layer charge (octahedral or tetrahedral) after application of suitable
tests such as the Green-Kelly test (Brindley, 1984).
X-ray diffraction is the only effective way to recognize bentonites consisting of
mixed-layer clay minerals with random interstratification such as mixed layer-illite-
smectite (R0 I-S), which often resembles discrete smectite in XRD traces. Pure smectite
can be distinguished from R0 I-S by the lack of migration of the higher-order basal dif-
fraction maxima (Moore & Reynolds, 1997). Mixed-layer I-S is the dominant clay
mineral present in K-bentonites (Huff et al., 1999). K-bentonites have inferior properties
382 G. E. Christidis

and they cannot be used in most industrial applications of bentonites. However, some
of these clays can be activated with acids and are subsequently used for purification-
decolorization of edible oils (Christidis et al., 1997).
X-ray diffraction results are complemented by other techniques such as thermal
analysis and infrared (IR) spectroscopy. The use of Differential Thermal Analysis
(DTA) or Differential Scanning Calorimetry (DSC) coupled with Thermogravimetry
(TG) and Evolved Gas Analysis (EGA) provide information about the dehydroxylation
and heat decomposition of the smectite present in the bentonite (Emmerich et al., 2011),
which is useful in assessing the likely behaviour of the clay in foundry sands and in iron-
ore pelletizing. Trans-vacant smectites, in general, have dehydroxylation temperatures
of ,6008C (Drits et al., 1998); hence, they are unsuitable for foundry applications.
Infrared spectroscopy may be used to recognize the type of smectite present in the ben-
tonite (Wyoming or Cheto montmorillonite, beidellite, nontronite, saponite etc.). Ben-
tonites containing Cheto-type montmorillonite are generally richer in Mg and poorer
in Fe than their Wyoming counterparts, and this can be observed in their IR spectra.
Electron beam techniques such as scanning and transmission electron microscopy and
electron microprobe analysis (Pownceby & MacRae, 2011) and image analysis (Pirard
& Sardini, 2011, this volume) provide additional information about the textural relation-
ships between the smectite and associated mineral impurities. In this respect, it is impor-
tant to determine the textural relationships between smectite and opal-CT commonly
present in bentonites. Opal-CT is usually ,2 mm in size and forms intimate inter-
growths with smectite flakes, affecting the rheological and binding properties of the ben-
tonite (Christidis & Scott, 1996). Finally, chemical analysis which is important in the
evaluation of many industrial minerals is not particularly useful in the assessment of
bentonites, because it does not give guidance in assessing the technical properties and
it does not give much indication of the nature of the minerals present (Christidis &
Scott, 1993). An important chemical restriction is the necessity for limited amounts
of S for foundry applications.

5.3.2. Quality and grade of bentonites

In the assessment of many industrial mineral and ore deposits the terms grade and
quality are synonymous and express the concentration of the useful chemical element
or mineral in the raw material. However, the performance of bentonites in the various
applications does not depend only on the smectite content. Although a bentonite rich
in smectite i.e. a high-quality bentonite, is usually a good raw material for many
uses, other bentonites containing lesser amounts of smectite can perform equally well
in some applications (Christidis & Scott, 1993). In this context, the term quality
refers to the expected performance of bentonites in the different applications, whereas
the term grade refers to the smectite content of the bentonite (Inglethorpe et al.,
1993; Christidis & Scott, 1993).
The quality of a bentonite can be assessed in the laboratory by the Na-carbonate
exchange and subsequent determination either of the free swelling index or the liquid
limit. The Na-carbonate exchange (also known as Na- or alkali-activation) refers to
the carbonate content added to Ca and/or Mg bentonites required to obtain the
optimum property. In general Na-exchanged bentonites have inferior properties for
 Industrial clays 383

most applications compared to natural Na-bentonites (Christidis & Scott, 1993). Na-
exchange also often forms materials with unpredictable behaviour (Odom, 1984). The
free swelling index is the volume of the gel of a predetermined amount of clay
(usually 2 g or 10 g) in distilled water. The liquid limit is the water content required
for flow of bentonite due to its weight. It is a measure of the bonding efficiency of
the clay and it is a routine test required in the foundry industry. The liquid limit can
be measured either by the Casagrande method or the cone penetrometer method
(BS 1377:1990: part 2). For Ca- and/or Mg-bentonites the swelling test and liquid
limit are determined by adding different Na-carbonate contents (1 – 6%) to assess the
optimum properties (Fig. 21). In the case of bentonites the Casagrande method or the
cone penetrometer methods give comparable liquid limit values (Fig. 22).
The assessment of bentonite grade can be carried out either directly by quantitative
XRD methods (Bish & Ploetze, 2011, this volume) or indirectly by measurement of
the CEC or the surface area by ethylene glycol monoethylether (EGME) (Inglethorpe
et al., 1993; Quirk & Murray, 1999) or polyvinylpyrrolidone (PVP) (Blum & Eberl,
2004). Measurement of the CEC can be achieved by several methods including
exchange with protons, organic cations such as methylene blue, ammonium ions,
alkali or alkaline earth cations transition metal ions and organo-metal complexes such
as silver thiourea cations, [Cu(en)2]2þ or [Cu(trien)3]2þ cations (Bergaya et al.,
2006). Methods based on the adsorption of methylene blue and organo-metal complexes
are not affected by the presence of zeolites, due to the large size of the molecules, which
cannot enter the zeolite channels. In contrast, ion exchange with alkalis or alkaline
earths index cations are affected by the presence of zeolites. Although adsorption of
methylene blue cations is routinely used by the industry, determination of CEC is pro-
blematic (Ammann et al., 2005). According to Bergaya et al. (2006), the use of

300

1
250
2
3
ml gel/ 10 g clay

200 4
5
150

100

50

0
0 1 2 3 4 5 6 7
Na 2CO3 added (% wt)

Fig. 21. Determination of the swelling index (ml gel/10 g clay) for a series of bentonites from Milos, Greece.
384 G. E. Christidis

600

LL Cone Penetrometer method

550
500
450
400
350
300
250
200
200 300 400 500 600
LL Casagrande method

Fig. 22. Projection of the Liquid Limit (LL) determined in a series of bentonites from Milos, Greece, using the
Casagrande method and the Cone Penetrometer method.

[Cu(en)2]2þ or [Cu(trien)3]2þ complexes is currently the most versatile method for

determination of CEC of clay minerals.
An alternative method for calculation of the CEC is from the structural formula and
the layer charge of smectite determined either from chemical analysis of the clay frac-
tion or from microprobe analyses (Lagaly, 1994). The CEC determined by this method
refers to a totally anhydrous basis and is not comparable to the CEC determined by the
wet methods described previously (Christidis, 2008b).

5.3.3. Factors affecting quality and grade

The swelling of a bentonite is controlled by both quality and grade. Grade is affected by
the amount of mineral impurities, which do not swell, or by geological processes, which
may postdate bentonite formation, such as hydrothermal alteration. During hydrother-
mal alteration, smectite may be altered to mixed-layer I-S, or kaolinite. Moreover, pre-
cipitation of secondary phases such as carbonates and sulphates effectively decreases
smectite content, thereby decreasing grade. Quality depends on the crystal-chemical
characteristics of the smectite present, mainly the layer charge. Distribution of layer
charge in a bentonite deposit has to be taken into account during assessment of a ben-
tonite deposit, because it assists in understanding unpredictable behaviour of bentonites
in routine industrial processes like Na-activation. Quality can also be affected by sec-
ondary processes such as hydrothermal alteration, which may result in a reduction of
the pH of the bentonite gel, the formation of smectite with different layer charge and
the formation of smectite with different lattice dimensions caused either by different
substitutions in the octahedral sheet and/or by oxidation or reduction of octahedral
Fe. Reduction of the pH of a bentonite gel due to circulation of acidic hydrothermal
water through the bentonite may affect Na-activation with Na2CO3, because a
 Industrial clays 385

significant proportion of the carbonate is consumed to neutralize the acidic clay

(Christidis & Scott, 1993).

5.3.4. Next stage of bentonite assessment

Determination of quality and grade are appropriate for preliminary evaluation of a ben-
tonite deposit. The aforementioned tests are not specific to any industrial application
because they provide data on the properties of the clay from which potential uses can
be established. Additional tests on the same samples, such as acid-activation exper-
iments for various times using different acid strengths followed by specific surface
area measurements and determination of bleaching capacity will provide information
about the potential of the clay for purification and clarification of edible oils (Christidis
et al., 1997).
Further tests on bentonite are unique to individual industrial users, e.g. viscosity and
fluid loss for drilling mud and strength of bonding with sand for foundry uses, or may
require trials in use. These tests usually require substantial amounts of bentonite and
should be representative of the available material in the mineral deposit (Christidis &
Scott, 1993). The tests are carried out by specialized equipment, which is not usually
available in laboratories and procedures are often defined in industry standards or
national and international norms. The standards describe specifications for acceptable
quality, which the bentonite designed for certain applications must meet. Special care
should be taken to avoid contamination or weathering of samples before testing.
A special case of bentonites are the so-called white bentonites. They have been classi-
fied as specialty clays (Russell, 1991), and because of their whiteness, their market value
is greater than the usual coloured bentonites. During assessment of white bentonites, as
well as the steps described before, two additional physical properties are also deter-
mined, colour and grit content. Determination of colour follows the CIELab system
using colourimeters (Billmeier & Saltzman, 1981). Good white bentonites have bright-
ness values in excess of 85%. The grit content is made up of non-clay minerals, which
are coarser than smectite. In most white bentonites, opal-CT is the major impurity
(Christidis & Scott, 1997; Allo & Murray, 2004), but is very fine-grained (usually of
clay size), and therefore is difficult to remove with processing. Opal-CT-bearing bento-
nites are often not exploited because of their abrasiveness, which renders them unsuita-
ble for large-added-value uses such as pharmaceuticals and cosmetics. Thus, it might be
more appropriate to concentrate on development of mineral separation techniques for
these materials rather than to search for new white bentonite deposits.

5.4. Mining and processing of bentonites

Bentonites are mined with opencast methods, using bulldozers, scrapers and excavators.
Explosives may be used during removal of the overburden but they are not necessary
during excavation of bentonites because they are soft materials. When mined, most ben-
tonites have moisture contents of 25– 35%. Processing includes drying, grinding, sizing
and Na-activation in the case of Ca-bentonites. Na-activation takes place before drying
and includes addition of suitable amounts of Na-carbonates. Activation may take place
386 G. E. Christidis

either by dry or wet methods. Dry methods are applied in the field where bentonites are
spread over large areas. Wet methods are applied in the processing plant in the wet clay
with or without addition of water. Excessive water is avoided because of higher drying
costs (Eisenhour & Reisch, 2006). The final products which result from drying have
7 – 12% moisture. Drying is carried out in the field in countries with dry climate such
as Greece and Cyprus and then in the processing plant, usually with rotary dryers.
Drying in the field is not complete but reduces moisture and thereby reduces the pro-
duction costs. In arid regions, field drying can reduce moisture to 12% (Eisenhour &
Reisch, 2006).
Production of large-added-value products (cosmetics and pharmaceuticals) is feasible
by washing with water. Bentonites form slurries with low solids and are subjected to wet
screening, hydrocyclones and centrifugation which remove coarse-grained impurities
(Eisenhour & Reisch, 2006). Na-activation is carried out before drying to produce
flake products or by spray to produce powdered products. Fine milling may reduce par-
ticle size further (Eisenhour & Reisch, 2006). Acid-activated bentonites are produced by
addition of sulphuric or hydrochloric acid to a bentonite slurry. The reaction lasts for
several hours. Residual acid is removed by pressure filtration. Alternatively, the bento-
nite is extruded forming noodle-like grains which are subsequently acid leached and
washed (Eisenhour & Reisch, 2006).

5.5. Distribution of bentonite deposits

Bentonite deposits are widespread throughout the world. The most important bentonites
from the economic point of view form from in situ digenetic alteration of volcanic
glass in an aqueous, mainly marine, environment, or from hydrothermal alteration of
volcanic/volcaniclastic rocks. Most commercial bentonites contain Ca or/and Mg
smectites, whereas bentonites with Na-smectites are rarer. In the following paragraphs,
important bentonite deposits from various places are presented.
The most significant Na-bentonites in occur in the western United States in Wyoming,
Montana and South Dakota (Fig. 23) and are known as Wyoming bentonites or Western
bentonites. The bentonite occurs in the form of beds of various thicknesses in the
Newcastle, Mowry and Belle Fourche Formations of Upper Cretaceous age. The most
important bentonite bed is the 0.5– 2 m thick Clay Spur bed (Grim & Güven, 1978). It
has been generally accepted that the Wyoming bentonites formed from alteration of
acidic volcanic ash erupted from volcanoes located west of the deposits (Murray,
2007), although recent geochemical work has shown that the volcanic ash had variable
composition, reflecting parent rocks with different geochemical affinities (Bertog et al.,
2007). Alteration is said to have taken place by diagenetic pore fluids that evolved from
Mowry seawater (Elzea & Murray, 1994).
Other commercial bentonites in the USA, which contain Ca/Mg-smectite occur in the
area of the Gulf coast and the southwestern states of California, Arizona and Nevada.
Moreover, most fuller’s earth deposits in the US present in Georgia, Illinois, Missouri,
Mississippi and Tennessee contain Ca-smectite as the main mineral and therefore
are classified as bentonites (Murray, 2007). The bentonites from the Gulf coast are
 Industrial clays 387

known as Southern bentonites

Edmonton
and crop out in Texas, Mississ-
ippi and Alabama. They occur
as thin beds formed also from Calgary ALBERTA
diagenetic alteration of acidic vol-
canic ash deposited in seawater,
which belong to the Eutaw, Ripley
and Vicksburg Formations with CANADA
USA
age ranging from Upper Cretac-
eous to Middle Tertiary (Grim & Glasgow
Güven, 1978). In the southwestern MONTANA NORTH
DAKOTA
states, significant bentonite depos-
its occur in Cheto (Arizona), in Billings
the Amargoza Valley (Nevada) SOUTH
and in Hector (California) and Sheridan DAKOTA
Greybull Belle Fourche
are younger in age (Murray, 2007). Newcastle
Black Hills
A common characteristic of these Big Horn Kaycee
Mountains
bentonites is that they contain
WYOMING Casper
Mg-rich smectites. The Cheto
bentonite, which occurs in the
Pliocene Bidahochi Formation
formed at the expense of an
0 100 200 Miles
acidic vitric ash and consists of
0 100 200 300 Km
Mg-rich montmorillonite. In the
Amargosa Valley, several ben- Fig. 23. Location of outcrop area of the Mowry bentonites
tonite deposits consisting of (modified from Murray, 2007).
trioctahedral smectites (mainly
saponite, stevensite and mixed-layer stevensite-kerolite) and Mg-montmorillonite are
present, with age ranging from Miocene to Pleistocene. Finally, the bentonite in
Hector contains hectorite, which formed by reaction of ascending Li-F-rich hydrother-
mal waters with travertine, i.e. hectorite did not form from alteration of volcanic glass.
As well as in the USA itself, other commercial bentonite deposits from the American
continent have been reported in Canada, Mexico, Argentina, Peru, Brazil and Chile
(Elzea & Murray, 1994; Eisenhour & Reisch, 2006; Murray 2007). The Canadian depos-
its in Saskatchewan and Alberta have the same age and geological characteristics as the
Wyoming bentonites but they contain Ca-rich smectites. In Argentina, several types of
bentonites including opal-CT-bearing white bentonites have been described in several
parts of the country (Lombardi et al., 2003; Allo & Murray, 2004).
Europe is a major bentonite producer with large deposits in Greece, Italy, the UK,
Spain, Germany, Turkey and Cyprus. The Greek bentonites occur in the Islands of
Milos, Kimolos and Chios. The most extensive deposits occur in Milos Island. They
are composite, up to 60 m thick and have formed by in situ low-temperature hydrother-
mal alteration mainly of successive pyroclastic flows and to a lesser degree of lavas and
tuffs of Pleiostocene age in marine conditions (Christidis et al., 1995; Christidis, 2008a).
388 G. E. Christidis

The volcaniclastic rocks of variable composition ranging from rhyolitic to andesitic,

which were erupted from three different volcanic centres, provided the energy for cir-
culation of alteration fluids. The deposits have subsequently been affected by posterior
hydrothermal alteration, which caused illitization of smectite and the formation of barite
deposits (Christidis, 1995). Most deposits also display evidence of alteration by acidic
waters rich in – SO24 – .
The Italian bentonites occur mainly in Sardinia. Two different types have been distin-
guished: type I includes in situ altered bodies within thick rhyolitic-rhyodacitic pyro-
clastic sequences formed by deuteric alteration and type II are distal fall-out lens
deposits interlayered with fluvial-deltaic and/or lacustrine Miocene sediments,
formed at 508C (Palomba et al., 2006). In Spain, important bentonite deposits of
Miocene age occur in Almerı́a. The deposits formed via low-temperature alteration of
pyroclastic rocks with composition ranging from rhyolitic to andesitic by hydrothermal
systems dominated by marine and/or meteoric waters (Caballero et al., 2005). Altera-
tion of the volcanic rocks by acidic solutions also took place, and produced sulphates
and kaolinite. In England, Ca-bentonites known as fuller’s earths of Middle Jurassic
(Bathonian) and Lower Cretaceous (Aptian) age have been described in Bath, Redhill
and Aspley Heath in southern England (Jeans et al., 1977; Jeans, 2006a,b). The bento-
nite beds are up to 5 m thick and were formed from submarine alteration of trachytic
volcanic ash, which was transported by rivers draining areas of freshly deposited
ash (Jeans et al., 1977). The origin of the ash has been attributed to volcanoes in
the North Sea. The Cyprus bentonites are of Upper Cretaceous age and have
formed by low-temperature hydrothermal alteration of andesitic to basaltic-andesitic
pyroclastic rocks related to the Troodos ophiolitic complex (Christidis, 2006). In
Turkey, numerous deposits of Neogene age occur in several parts of the country.
Most deposits have hydrothermal origin (e.g. Yildiz & Kuscu, 2007), but bentonites
derived from alteration of glass in lacustrine environments have also been reported in
SW Turkey (Kadir, 2007).

5.6. Applications of bentonites

Bentonites are used in numerous applications both in the Na-form and the Ca-form.
A list of the applications of bentonites is given in Table 8 (Murray, 2007). The most
important in terms of the amount of bentonite consumed and of the potential for future
applications are summarized in the following paragraphs.
Drilling fluids contain Na-smectites, either natural or Na-activated. The drilling muds
remove cuttings from the drill hole, lubricate and cool the drilling bit and build an
impermeable filter cake on the drill wall to keep the formation fluids from penetrating
into the drilling mud and vice versa. Generally a high viscosity and a low yield stress and
thixotropic behaviour are required from the bentonites. Rheological properties are opti-
mized by use of suitable additives. For instance, addition of activated lignite improves
the rheological properties (Kelessidis et al., 2007). The specifications for bentonites used
in drilling fluids are described in standards of various organizations such as the Amer-
ican Petroleum Institute (API) and the Oil Companies Materials Association (OCMA).
 Industrial clays 389

Table 8. Applications of smectite-rich clays (adapted from Murray, 2007).

Drilling muds Crayons Paint
Foundry sands De-inking newsprint Paper
Iron ore palletizing Desiccants Pencil leads
Cat litter Deodorizers Pharmaceuticals
Absorbents Detergents Pillared clays
Adhesives Emulsion stabilizers Plasticizers
Aerosols Fertilizer carrier Rubber filler
Animal feed bonds Food additive Sealants
Barrier clays Fulling wool Seed growth
Bleaching earths Herbicide carrier Soil stabilization
Catalysts Industrial oil absorbent Slurry trench stabilization
Cement Insecticide and pesticide carrier Suspension aids
Ceramics and refractories Medicines Tape joint compounds
Cosmetics Nanoclays Water clarification
Organoclays

The majority of foundry sands consist of silica sand and 5 – 10% bentonite. Bentonite
provides bonding strength and plasticity to the sand-clay mixture in order to maintain
the shape of the mould before and after pouring of the molten metal. Although according
to Inglethorpe et al. (1993) both the Na- and the Ca-bentonite are used in the foundry
industry, the former have superior performance due to the greater wet tensile strength
(Christidis & Scott, 1996), which increases their resistance to metal defects known as
scabs (Fig. 24) (BCIRA, 1985). In summary, mixtures of sand and Na-bentonites
have medium to low green compression strength and high dry compression strength,
whereas mixtures of sand and Ca-bentonites have medium green compression strength
and low dry compression strength. Recycling after use is an important parameter con-
sidered when a bentonite is selected for foundry applications. Trans-vacant smectites
have lower dehydroxylation temperatures than their cis-vacant counterparts. Therefore,
they display a lesser tendency for recycling and they are unsuitable as bonding agents in
foundry sands.
Large quantities of natural sodium or Na-activated bentonite is used for pelletizing of
fine-grained iron ores. Usually a proportion of 0.50% of bentonite is added to the ore

Fig. 24. Formation of scab defects in greensand moulds formed from penetration of molten metal in
weakness zones.
390 G. E. Christidis

to form pellets, which are sintered at 12508C. In some pellets, olivine is also added in
small amounts. The role of bentonite is to adsorb excess water from the iron fines, to
provide sufficient drop, green and dry compression strength to the pellet and to
improve the mechanical properties of the fired pellets reducing the quantity of the
fines. Bentonites are also used to pelletize animal feeds.
Na-bentonites are used in several civil engineering applications including grouting,
construction of diaphragm walls, and lubrication of caissons and piles. They are also
added in small proportions (usually 1 – 2%) to Portland cement in concrete and
cement slurries, to improve workability, to lessen aggregate segregation and increase
impermeability. The physical properties required for these applications include rheolo-
gical properties (viscosity and thixotropy) and impermeability. Natural Na-bentonites
and Na-activated bentonites are used in the construction of clay liners in engineered
landfill sites for solid and liquid waste (Czurda, 1993; Koch, 2002) and in the fabrication
of geosynthetic clay liners used for the same purpose (Browning, 1998). In multibarrier
systems, Ca-bentonites can also be used (Czurda, 1993). Extensive research has been
carried out on the use of Na-bentonite as backfill in high level radioactive waste repo-
sitories (Pusch, 2006). Important properties examined for this application are the low
hydraulic conductivity and the large CEC value of bentonites and their resistance to
alteration to illite when in contact with pore waters.
Large quantities of acid-activated bentonites, often referred to as ‘bleaching earths’
are used by the mineral oil and foodstuffs industry, in sulphur production, in forest
and water conservation, in environmental protection, by the beverages and sugar indus-
try, by the chemical industry, by the paper industry and in cleaning and detergent pro-
ducts. Their main application is the purification, decolorization, and stabilization of
vegetable oils. They remove phospholipids, soaps, trace metals, organic compounds
(carotenoids, xanthophylls, chlorophyll, pheophytin, tocopherols, and gossypol) and
their degradation products (Zschau, 2001). Acid activation is carried out on Ca-bento-
nites and increases their specific surface area significantly. Evaluation of bentonites
as bleaching earths includes construction of STS diagrams and evaluation of bleaching
capacity (Christidis et al., 1997).
Bentonites are used as granular adsorbents in most types of animal litters. In this
application there is direct competition from sepiolite, attapulgite and diatomite. The
advantage of bentonites is that they form clumps when wetted and can therefore be
removed easily. The specifications are not rigid and include water adsorption (Westing-
house test), freedom from dust and uniform grain size. Except for aqueous solutions
Ca-bentonites can adsorb large amounts of oil and grease. However, they tend to
slake, producing a slippery surface and therefore are easily replaced by sepiolite and
attapulgite (Inglethorpe et al., 1993). Many Ca-bentonites can adsorb up to 100% of
their dry weight of water and up to 80% of their weight of oil.
The tendency of bentonites to react with organic molecules yields ‘organoclays’,
which have gained increasing significance in the chemical industry (Schoonheydt &
Bergaya 2011). Usually Na-bentonites are used, due to the greater tendency of Na-smec-
tites to break quasicrystals and form single layers. The most frequently used organic
compounds are quaternary ammonium salts and to a lesser degree polymeric quaternary
 Industrial clays 391

alkylammonium salts and copolymers (de Paiva et al., 2008). Organoclays are used
mainly as polymer nanocomposites and adsorption materials and to a lesser degree as
gel formers, catalysts, soil remediation and as electrodes (Xu et al., 1997; Le Baron
et al., 1999; de Paiva et al., 2008) and their importance has increased significantly
over recent years.

6. Sepiolite and palygorskite

6.1. Introduction
Sepiolite and palygorskite are Mg-rich clay minerals, which form crystals with charac-
teristic fibrous or lath-like habits, attributed to their ribbon-like structure (Fig. 8). In the
USA the term fuller’s earth, which describes sorptive clays also includes palygorskite-
rich clays, although in the UK this term describes mainly Ca-rich bentonites. The term
‘attapulgite’ is often used as synonymous with palygorskite, though this is not rec-
ommended by the nomenclature committee of AIPEA. Other terms used in the past
to describe sepiolite and palygorskite are ‘meerschaum’ (meaning sea-froth in German)
and ‘hormites’. However, these terms are also not recommended by AIPEA.
Sepiolite and palygorskite have small degree of ionic substitutions in their structures
and therefore have limited CEC. They are considered to be sorptive clays because of
their large specific surface area, microporosity and sorptive capacity. These properties
are attributed to the presence of channels in their structure. They also have great mech-
anical strength and thermal stability. Finally, due to their fibrous shape they have less
tendency to flocculate and are used as suspending agents in suspensions with large elec-
trolyte concentrations.

6.2. Genesis of sepiolite-palygorskite deposits

Sepiolite and palygorskite form in a limited number of geological environments. Their
formation is controlled by Mg- and Si-activity, pH and the presence or lack of Al in the
pore waters (Meunier, 2005). Both minerals are favoured by high Mg-activity and a high
pH. Three main types of genetic envoirnoments can be distinguished, namely marine,
continental lacustrine and continental soils and calcretes (Velde & Meunier, 1987;
Jones & Galán, 1988). A small number of sepiolite and palygorskite occurrences
without economic significance are attributed to low-temperature hydrothermal altera-
tion of igneous rocks (Jones & Galán, 1988). It is generally accepted that both palygors-
kite and sepiolite in continental environments in general are authigenic phases. On the
other hand the presence of palygorskite in marine sediments may indicate the diagenetic
or sea weathering origin especially in abyssal sediments (Couture, 1977; Meunier,
2005), or may suggest a detrital origin in marine sediments adjacent to nearby land
masses (Jones & Galán, 1988). Authigenic palygorskite may also form in shallow,
lagoonal peri-marine environment.
The authigenic formation of palygorskite in marine environments necessitates a sub-
strate which is usually of smectite or of volcanic glass shards, and high Si-activity in the
392 G. E. Christidis

pore water (Couture, 1977; Weaver & Beck, 1977; Meunier, 2005; Christidis, 2006).
Replacement of smectite by palygorskite is an incongruent dissolution precipitation
process and excess silica is provided by dissolution of diatoms or radioalaria frustules
(Weaver & Beck, 1977; Christidis, 2006). Both smectite and volcanic glass also
provide the Al and Fe necessary for palygorskite. Sepiolite is not common in deep
marine sediments. A shallow lagoonal peri-marine environment may be important for
the formation of palygorskite deposits of economic significance. This is the case for
the Miocene deposits of Georgia and Florida, USA (Weaver & Beck, 1977).
Sepiolite and palygorskite often form in arid climates in inland saline lakes and
basins, in which detrital inputs consist of other clay minerals such as kaolinite, illite,
chlorite and Al-Fe-rich dioctahedral smectite. These environments display a minera-
logical zonation characterized by a decrease of the detritus and increase of the abun-
dance of neoformed Mg-silicates towards the centre of the basin (Velde & Meunier,
1987; Ece & Coban, 1994; Meunier, 2005). Sepiolite is the main phase formed by
direct precipitation in the centre of the basin in which detrital input and addition of
Al are minimized (Fig. 25). Palygorskite forms by dissolution of detrital minerals
which provide Al. Mg-rich smectites and trioctahedral smectites (stevensite and sapo-
nite) are also common in these environments. Their formation is favoured by the avail-
ability of Si in the lake water. The presence of amorhous Si which buffers mSiO2 retards
the formation of trioctahedral smectites, favouring sepiolite or/and palygorskite
(Meunier, 2005). Most sepiolite deposits of economic importance form in inland
saline lakes and basins. The Spanish sepiolite deposits and the large deposit of palygors-
kite, recently discovered in northern Greece belong to this category (Galán & Castillo,
1984; Kastritis et al., 2003).
Palygorskite and sepiolite commonly form in soils of semi arid and arid climates in
which the mean annual rainfall does not exceed 300 mm (Jones & Galán, 1988). The
presence of fibrous Mg-rich clays is closely related to the formation of calcite, which

Evaporation

Saline-alkaline lake

Legend
Basement rocks sm, pal

kaol, sm, ill,

sep, pal
chl (all clastic)

Fig. 25. Genetic model for the formation of sepiolite and palygorskite in inland saline-alkaline basins. Key:
kaol ¼ kaolinite, ill ¼ illite, chl ¼ chlorite, sm ¼ smectite, pal ¼ palygorskite and sep ¼ sepiolite.
 Industrial clays 393

forms through the evaporation of subsurface soil waters which migrate by capillary
action under highly evaporative conditions (Velde & Meunier, 1987). If soil waters
are rich in Si and Mg then Mg-silicates precipitate. Detrital soil minerals such as smec-
tite and illite are unstable in those soil horizons in which Si and Mg concentrate due
to evaporation and dissolve to form sepiolite and palygorskite along with other clay
minerals such as kaolinite (Meunier, 2005). An additional source of Mg can also be
high-Mg-calcite which recrystallizes to its low-Mg counterpart (Jones & Galán,
1988). In closed systems, soil sepiolite and palygorskite form in equilibrium with soil
waters and have certain Al2O3 contents. However, if the system is open then both min-
erals contain less Al2O3 (Meunier, 2005). Although genesis of palygorskite and sepiolite
in arid soils is a rather common process, it does not lead to the formation of economic
deposits because of the abundance of detrital impurities.
Minor occurrences of sepiolite and palygorskite have been reported in veins,
suggesting a hydrothermal origin (Jones & Galán, 1988). These occurrences have
been attributed in general to low-temperature hydrothermal activity and do not yield
deposits of economic importance.

6.3. Assessment of palygorskite-sepiolite deposits

Assessment begins with recognition in the field. Most palygorskite and sepiolite depos-
its contain smectites which are either detrital (dioctahedral) or authigenic (trioctahedral)
which create ‘popcorn’ textures similar to those observed for bentonites. The deposi-
tional environments of both minerals are indicative of sedimentation under arid climatic
conditions, usually terrestrial. Hence the deposits are characterized by the presence of
gypsum or calcretes (Velde & Meunier, 1987; Meunier, 2005). Finally, sepiolite and
palygorskite-rich clays have low specific gravity, due to the presence of channels in
their structures. These are diagnostic features, which allow us to distinguish them
from other industrial clays.
Confirmation of the identity is by mineralogical charactrization of the deposits
using XRD. Sepiolite and palygorskite are characterized by diffraction maxima at
12.4 Å and 10.6 Å, respectively, which do not migrate to lower angles after solvation
with ethylene-glycol vapours. In this manner, sepiolite is easily distinguished from
Na-smectite, which swells to 17 Å when subjected to ethylene glycol vapours.
Samples rich in palygorskite and sepiolite have smaller specific gravity and greater
BET specific surface areas, determined by adsorption of N2 gas, than other clay min-
erals. The BET specific surface areas of sepiolite and palygorskite are 300 m2/g
and 150 m2/g, respectively (Galán, 1996). Specific surface area is determined after
outgassing at 100– 1508C, when 10% of hygroscopic water is lost (Galán, 1996).
Heating at greater temperatures decreases the specific surface area due to micropore
destruction and structure folding (Serna et al., 1975; Post et al., 2007; Post &
Heaney, 2008).
Assessment for certain industrial applications includes the determination of certain
physical and chemical properties. Determination of rheological properties is the main
task for assesment of sepiolite and palygorskite in drilling fluids. Usually, suspensions
394 G. E. Christidis

of certain concentration are prepared and the apparent viscosity, plastic viscosity, and
filtrate loss are determined according to certain specifications, such as those of the
API (API, 1993), using Couette-type viscometers. Similar procedures are followed
during assessment of bentonites for drilling fluids, except for the fact that rheological
properties are determined in 40% (w/v) NaCl solutions (Murray, 2007). In the case
of sepiolite, viscosity measurements can also be carried out at higher temperatures.
The most important test for assessment as animal litters is water absorption using a
series of empirical absorption tests. The difficulty in this assessment is the lack of stan-
dardized water absorption tests, except for the Westinghouse test. Water absorption tests
are carried out both in raw and calcined samples. In general, palygorskite and sepiolite
have a greater water-absorption capacity per gram of clay compared with bentonites,
because of their low specific gravity. The ability for clumping (i.e. formation of wet
aggregates after addition of water, which can be separated from the dry clay) is an
important parameter, which is examined. Sepiolite and palygorskite have a lower
clumping ability than bentonites.

6.4. Mining and processing of sepiolite and palygorskite

Sepiolite and palygorskite are mined using the typical opencast methods used for ben-
tonites (Murray, 2007). Processing does not involve expensive beneficiation techniques
but includes mainly crushing, drying and grinding. The crushed material goes either
directly for drying or for extrusion. Sometimes MgO is added during extrusion to
improve viscosity. Drying may take place at low and high temperature (2008C and
6008C, respectively) according to the final application of the raw materials. Products
designed for the drilling industry are dried at low temperatures, whereas those designed
for adsorbents are dried at high temperatures. Drying is followed by grinding and sizing.
Certain varieties of sepiolite and palygorskite are used as bleaching earths. However,
contrary to bentonites, they do not undergo acid activation prior to use. This is because
both sepiolite and palygorskite dissolve readily even after mild acid treatment, and loose
their decolorization properties similar to Mg-rich smectites (Suarez-Barios et al., 1995;
Myriam et al., 1998; Balcı, 1999).

6.5. Distribution of sepiolite and palygorskite deposits

Although sepiolite and palygorskite are rather common clay minerals, there is a limited
number of palygorskite and sepiolite deposits of economic significance throughout
the world. Important palygorskite deposits occur in Florida and Georgia, USA, China,
Senegal and Greece. Sepiolite deposits of economic significance occur in Spain
and Turkey.
The palygorskite deposits in Florida and Georgia are of Early to Middle Miocene age
and dominate the world’s production. Palygorskite formed at the expense of Wyoming
montmorillonite and stevensite formed locally as a by-product (Weaver & Beck, 1977).
Montmorillonite provided Al and to a lesser degree Fe and additional Mg and Si were
supplied by solution. In this sense, dissolution of smectite may be considered an incon-
gruent dissolution-precipitation process (Jones & Galán, 1988). The excess Si was
 Industrial clays 395

mainly biogenic and was supplied by dissolution of sponges and diatoms. The deposi-
tional environment was shallow ‘peri-marine’ with estuaries and marine lagoons which
varied from saline to nearly fresh-water in a shallow low-energy environment (Krekeler
et al., 2004).
An important palygorskite deposit of Middle Miocene age occurs in Guanshan in
Anhui province, China (Zhou & Murray, 2003). The palygoskite bed is 3– 6 m thick
and formed at the expense of basaltic ash deposited in a lacustrine environment. In
Senegal, 3 – 6 m thick palygorskite-rich beds of Early Eocene Age overlie an Al phos-
phate deposit. Finally, in Greece, a palygorskite deposit was discovered recently in
Ventzia Basin, western Macedonia (Kastritis et al., 2003). The deposit formed in a
lacustrine environment and displays zonation, with palygorskite occurring at the
centre of the basin (Fig. 26). Palygorskite is Fe-rich and formed via diagenetic alteration
of detrital smectite which originated from the nearby ophiolite complex of Vourinos and
the smectite-bearing sands of the Mesohellenic trench (Kastritis et al., 2003). Alteration
of smectite was facilitated by Si-rich pore fluids.
Spain is the main producer of sepiolite but palygorskite outcrops are common as
well. Important lacustrine deposits of sepiolite of Miocene age occur at Vicalvaro
near Spain. According to Galán & Castillo (1984) four different types of sepiolite and
palygorskite can be distinguished: (1) sepiolite deposits of Miocene age in distal alluvial
fans and in perennial lacustrine sediments associated with trioctahedral smectite; (2)
palygorskite occurrences formed by alteration of chlorite by dissolution-precipitation;
(3) palygorskite diagenetic cement in sandstones; and (4) sepiolite and palygorskite
Pliocene deposits formed in a brackish lacustrine environment, in which sepiolite
occurs at the lower part of the deposit with carbonates, and palygorskite is present in
the higher sectors with more detrital input. It is interesting that the Spanish lacustrine
sepiolite deposits do not display the typical concentric zoning with sepiolite occupying
the centre of the basin (e.g. Meunier, 2005) but is distributed in marginal areas of the
basin (Jones & Galán, 1988).

Fig. 26. Schematic cross section of the palygorskite-smectite deposit in the Ventzia basin, Greece. The
deposit is characterized by zonal distribution of the various clay minerals (after Kastritis et al., 2003).
396 G. E. Christidis

The Turkish sepiolite deposits occur in western Anatolia, near Eskisehir. The deposits
are of Middle– Upper Miocene age and were formed in a saline to alkaline lacustrine
environment (Ece & Coban, 1994; Ece, 1998). Sepiolite occurs in the form of beds
with variable thickness and nodules. Sepiolite beds have formed via precipitation in
the centre of the basin. Three types of sepiolite beds are distinguished, black sepiolite
rich in organic matter, brown sepiolite poor in organic matter and with 5% dolomite,
and white-pale yellow sepiolite with 20– 40% dolomite (Ece & Coban, 1994).
Nodular sepiolite occurs in beds at the margins of the lake. Sepiolite formed via diage-
netic alteration of magnesite (Ece, 1998). Sepiolitization of magnesite has not been
observed in the centre of the basin.

6.6. Applications of sepiolite and palygorskite

Due to their structural features (inversion of tetrahedra and formation channels) and aci-
cular particle morphology palygorskite and sepiolite have numerous industrial appli-
cations. Both sepiolite and palygorskite develop important thixotropic and rheological
properties and are utilized in drilling fluids as competitors of bentonites. Their suspen-
sions are stable in the presence of brines and electrolytes. Rheological properties can be
improved by addition of 1 – 2% MgO. The viscosity of sepiolite is unaffected by pH for
pH , 8 (Alvarez, 1984). At pH , 4, both minerals tend to dissolve and the stability of
their suspensions deteriorates gradually. At pH .9, the sepiolite suspensions have
Newtonian characteristics. The optimum pH for rheological properties is 8.5, at which
sepiolite buffers the aqueous medium.
Sepiolite and palygorskite-rich clays are utilized as adsorbents, because they display
greater sorptive capacity for water and oil than the other clays. Due to their low specific
gravity, they can adsorb up to 100% of their dry weight in water and up to 80% of their
dry weight in oil. Important applications include animal litters and industrial floor adsor-
bents. Although they do not display clumping properties, their performance can be
improved by adding small amounts Na-bentonite (Murray, 2007). In order to increase
their sorptive properties, they are heated to 200– 3008C to remove zeolitic water from
their channels, although heating at such a high temperature may cause micropore
destruction and structure folding (Serna et al., 1975).
Sepiolite and palygoskite can replace expensive organic thickeners in emulsion paints
(Murray, 2007). The acicular particles prevent settling of chemicals and pigments. The
thixotropic properties of palygorskite and sepiolite also reduce sagging of paints. Those
properties also do not allow the growth of fungi, and the viscosity of the paints is unmo-
dified by hard water or temperature (Alvarez, 1984). An additional application of sepio-
lite and palygorskite includes pharmaceuticals, both as an excipient, on which the active
molecules are retained, and as an adsorbent of toxins, bacteria and liquid, in the treat-
ment of diarrhoeal processes. These applications are attributed to the large specific
surface area of both minerals (Alvarez, 1984). Similarly, both minerals are used as
binders and additives for animal nutrition, whereby they act as growth promoters, as
components of Amide concentrate supplements, as carriers for supplements, as
binders of feed and for stimulation of production. Moreover, they function in a
similar way to pharmaceuticals, i.e. they adsorb toxins and bacteria.
 Industrial clays 397

Additional uses of sepiolite and palygorskite include among others the fabrication of
carbonless paper, adhesives and caulks, asphalt, ceramics, cosmetics, foundry sand
binders (competitors of bentonites), laundry washing powders (competitors for synthetic
zeolites) etc. (Murray, 2007). Finally, although sepiolite and palygorskite have fibrous
morphology similar to asbestos, epidemiological studies on the workforce exposed to
the fibres on a daily basis, have shown their detrimental effect to human health is
minimal, contrary to asbestos (Santaren & Alvarez, 1994).

7. Common clays and shales

7.1. Introduction – terminology
Common clays and shales are used for the manufacture of building and engineering
bricks, tiles and agricultural land drains (Ridgeway, 1982). The end clay products are
fired, unglazed, structural clay ware. These materials are used for construction purposes;
therefore they are often called ‘structural clay products’ and the raw materials are called
‘structural clays’, ‘heavy clays’ or ‘brick clays’ although the manufactured products do
not include only bricks. Common clays and shales are sediments consisting of clay min-
erals and other silicates, carbonates, sulphides, sulphates and oxides-oxyhydroxides.
The type and abundance of the various minerals occurring in common clays and
shales, depend on the nature of the parent rocks, the climate, the depositional environ-
ment and diagenesis.
There are three main types of bricks depending to their properties and end uses (Rid-
geway, 1982). Common bricks (known as ‘commons’ in the UK), which are suitable for
general building purposes, ‘facing bricks’ (or facings) specially made for their attractive
appearance, which is achieved by variations in colour and surface texture, and finally
‘engineering bricks’, which display high strength and low water absorption. Similarly,
there are three types of tiles, namely ‘roofing tiles’, used for construction of the roofs of
buildings, ‘floor tiles’ used for covering floors and ‘quarry tiles’, which are resistant to
acids, oils and frost, i.e. they resemble engineering bricks. Finally, there are two main
types of clay pipes, the ‘porous pipes’ used mainly for agricultural field drains and ‘vitri-
fied pipes’.
Increasing temperature during firing causes progressive partial melting of the clay
bodies known as vitrification. With increasing vitrification the melt increases and the
apparent porosity of the fired product decreases progressively. The temperature interval
between the beginning of vitrification of a ceramic body and the temperature at which
the body begins to deform is known as vitrification range (Ridgeway, 1982). Vitrifica-
tion range depends on the chemical composition of the clay body. The presence of Mg-
rich minerals, such as chlorite, increases the vitrification range of the ceramic bodies
(Grimshaw, 1971).

7.2. Formation of deposits of common clays

The weathering products of igneous and metamorphic rocks in the Earth’s crust may
remain in situ as residual clays or more commonly they may be transported and
398 G. E. Christidis

deposited in sedimentary basins. Additional modifications to the sediment may be

caused during diagenesis. The type of clay mineral assemblage which may form
depends on the climate and the lithology of the source area. Acid tropical-subtropical
soils and soils of warm temperate climates which receive .635 mm of rain are domi-
nated by kaolinite (Dixon, 1989; Wilson, 1999). Smectite forms under neutral or alka-
line conditions. Illite forms in many temperate soils including podzols, whereas chlorite
occurs in soils of cold and arid regions. Finally, sepiolite forms in soils of semi-arid and
arid climates in which the mean annual rainfall does not exceed 300 mm and in inland
saline lakes and basins (Jones & Galán, 1988). In arid and semi-arid environments,
evaporation of groundwater leads to precipitation of carbonates (mainly calcite) in
the soil profile, i.e. the formation of calcretes (Velde & Meunier, 1987). The influence
of rainfall and lithology on clay mineral assemblages is shown in Figure 27.
Thick beds of mudstones form in aqueous environments in which current velocity is
low and decreasing (Ridgeway, 1982). Deposition of clay minerals is favoured by floc-
culation, when rivers discharge their sedimentary load to the marine environment.
Characteristic depositional environments include river floodbasins, delta tops, estuaries,
lagoons, continental shelves and ocean basins. In environments with bottom stagnant, or
hypersaline waters such as those found in lakes, floodbasins and tidal esuarines, lami-
nation of the clay sediments may be well preserved. In contrast, in well oxygenated
marine environments, benthic organisms produce mottled or homogeneous texture.
Optimum conditions for thick accumulations of argillaceous sediments occur where
basins bordered by slowly rising shield areas and drained by rivers at the maturity
stage, are characterized by warm-humid climates. The clay minerals develop by weath-
ering of silicates forming thick soil horizons in these shield areas. Erosion of these soils

<2 µm fraction of soils on <2 µm fraction of soils on

siliceous igneous rocks basic igneous rocks
Relative clay mineral content (%)

Relative clay mineral content (%)

80 80 Annual temperature: 10–15.5°C

Annual temperature: 10–15.5°C
K S
60 60
K
S

40 I V 40
V

20 20
G
G
0 0
0 500 1000 1500 2000 0 500 1000 1500 2000
Mean annual precipitation (winter), mm Mean annual precipitation (winter), mm

Fig. 27. Variation in the clay mineralogy of Californian soils developed on acidic and mafic igneous rocks,
with mean annual precipitation (modified after Barshad, 1966). Key to the symbols: S ¼ smectite,
K ¼ kaolinite, I ¼ illite, V ¼ vermiculite, G ¼ gibbsite.
 Industrial clays 399

yields a sedimentary load, which also contains a considerable fraction of coarse-grained

sediments such as sand and gravel. The coarse fraction is trapped closer to the source
area, whereas the fine-grained sediments move downstream forming thick deposits of
silt and clay in the more distal parts of the basin.
When the suspended matter is transported to the sea, the greater fraction is deposited
in the continental shelf and a smaller fraction is transported by to the deeper parts of the
ocean basins by turbidity currents. In general, thick clay deposits in the continental shelf
are deposited during marine transgression periods. In areas with low rates of clastic sedi-
mentation and warm climates, deposition of carbonates takes place. In fact most clays
contain variable amounts of carbonates, mainly calcite. In lagoonal environments, the
Mg/Ca ratio of the water may be large enough to allow precipitation of dolomite.
The composition of soils at the source areas affects the composition of marine sedi-
ments in the various oceans. For example, kaolinite is the main clay mineral in sedi-
ments around Australia and those areas of the central Atlantic, fed by the main rivers
of Africa and South America (Chamley, 1989). On the other hand, chlorite is a main
clay mineral in marine sediments at high latitudes adjacent to glaciated continental
areas. Illite is associated with weathering at temperate climates and is abundant in the
north Atlantic and north Pacific supplied by rivers from Eurasia and North America.
Smectites are abundant in areas remote from the main influxes of continental sediments,
in which volcanic input may be significant, and are characteristic of the south and central
Pacific. Older argillaceous sediments, which constitute common clay deposits, are
expected to display a similar distribution of clay minerals, which reflect palaeoclimatic
and palaeogeogrphic influence. However, the original clay mineral distribution may be
modifed by diagenetic reactions, the most important of which are illitization or chlori-
tization of smectite and kaolinite with burial depth. Hence the abundance of illite and
chlorite increase with time and that of kaolinite and expandable clays decreases accord-
ingly. These diagenetic changes affect the physical properties of the clays.

7.3. Assessment of common clays and shales

The assessment of common clays and shales differs from the other clays because effort is
focused on the fired end products rather than the raw materials. Although the mineralo-
gical composition of the raw material controls the physical and chemical properties of
the bricks and tiles, the physical properties of the fired products cannot be predicted
from the mineralogical and chemical composition of the raw materials. Assessment
begins with determination of the mineralogical and chemical composition of the raw
materials, which will assist to estimate the mineralogical composition after firing at
various temperatures (Dunham, 1992). The most important minerals present, other
than clay minerals, are quartz, feldspars, carbonates, Fe-bearing minerals (oxides, oxy-
hydroxides, carbonates, and sulphides) and organic matter. Minor phases include
anatase, gypsum and apatite. Gypsum may cause problems – undesirable colourations
in the form of stains known as scumming. Organic matter may provide an important
fuel element during firing as in the Oxford Clay, UK, or may remain unburned
and form undesirable black spots known as ‘black hearts’ which cause bloating,
400 G. E. Christidis

40 i.e. swelling of the ceramic. Carbonates

Plastic Limit (PL)

35 may also act as colour modifiers creat-

30 Suitable zone ing bricks with lighter colours after
25 firing or may form CaO grains, which
Optimum
20 zone may also cause bloating (Grimshaw,
15 1971).
The most important physical property
0 5 10 15 20 25 30 35 40 45 determined for raw clay materials is
Plasticity Index (PI) plasticity, which is assessed by determi-
Fig. 28. Prediction of the extrusion behaviour of nation of the Atterberg limits (plastic
clays according to the Atteberg Limits (modified limit, liquid limit, and plasticity
from Bain & Highley, 1978). index). In fact the plasticity index is
used as a guide to the workability of a
clay (Fig. 28). The plasticity of clays is usually assessed either by the Casagrande
method or by the cone penetrometer method (BSI 1377, 1990). Particle-size distribution
is also determined, and the raw material is classified according to its sand, silt and clay
content in suitable empirical triangular diagrams (Fig. 29). Such diagrams may be used
for a rough estimation of the possible final fired products (Winkler, 1954). Often it is
useful to examine the various size fractions with a binocular microsope.
The assessment of structural ceramic products is carried out on test specimens usually
produced with vacuum extruders. The specimens are fired at various temperatures
usually in the temperature range 700 – 11008C. For each fired specimen the mineralogi-
cal composition and certain physical properties are determined. The mineralogical com-
position of the fired products depends on temperature, the firing cycle and the
composition of the raw materials. Typical high-temperature minerals include melilite
(usually gehlenite), pyroxene, wollastonite and anorthite (Dunham, 1992). The pyrox-
ene has fassaitic composition and is known as “ceramic pyroxene” (Dondi et al.,

100%
<2 µm
1 = Solid bricks
2 = Vertically perforated bricks
3 = Roofing tiles, lightweight blocks
4 = Thin-walled hollow bricks

3 4
2
1

100% 100%
<20 µm <2–20 µm

Fig. 29. Technological classification of bodies of structural clay products according to Winkler (1954).
 Industrial clays 401

1998; Bauluz et al., 2004). Hematite forms from dehydration of goethite and/or oxi-
dation of pyrite or marcasite present in the orignal clay. Mullite may form at high temp-
erature in ceramic bodies which come from Al-rich clays, usually rich in kaolinite
(Onike et al., 1986). The mineralogical changes occurring in the most common miner-
alogical constituents of brick clays are summarized in Figure 30.
The physical properties usually determined in test specimens are drying and firing
linear and volume shrinkage, porosity, water absorption and mechanical properties
such as modulus of rupture and compressive strength. The colour of the fired bodies
is also recorded, either according to the Munsel colour chart or to the CIELab
system, using a colourimeter. Colour depends on the presence of certain minerals.
Hematite provides red colour to the fired bricks. The presence of free, fine-grained
CaO from decomposition of calcite yields yellowish colours. The colour may be modi-
fied by the dominance of reducing atmosphere in the kiln or the presence of certain addi-
tives, such as Mn. In general Mn yields dark colours to the fied ceramic products
(Grimshaw, 1971). Compressive strength and water absorption are the main properties
used to classify the bricks as engineering and common (Table 9). Water absorption and
porosity decrease with firing temperature because of the vitrification of the clay. Usually

Kaolinite

Illite
Key to the symbols
trans-vacant cis-vacant
Smectite Evolution of volatiles
Interlayer or adsorbed water
Sepiolite Structural water
Spinel formation
Chlorite
Mullite formation
SO2 Corundum formation
Enstatite formation
Pyrite
Olivine formation
CO2 Cristobalite formation
Oxidation
Calcite

Organic
Matter
a ß
Quartz

0 200 400 600 800 1000 1200

°C
Fig. 30. Schematic representation of the mineralogical transformations taking place in the brick clays during
firing. Transformations were considered up to 12008C.
402 G. E. Christidis

Table 9. Classification of bricks by compressive strength and water absorption (BSI 3921: 1985).
Class Compressive strength Water absorption
(N/mm2) (% by mass)
Engineering A 70 4.5
Engineering B 50 7.0
Damp-proof course 1 5 4.5
Damp-proof course 2 5 7.0
All other types 5 No limits
There is no direct relationship between compressive strength and water absorption and durability.
Damp-proof course 1 bricks are recommended for use in buildings whilst damp-proof course 2 bricks are recommended for
use in external works.

the physical properties of the fired products are presented vs. temperature in diagrams
(Bain & Highley, 1978; Artigas et al., 2005).

7.4. Application of TTT diagrams in the brick industry

The mineralogical composition of bricks is usually examined with the use of phase dia-
grams. This approach considers that the system under study is at equilibrium. In general,
liquid-to-solid transformations follow changes which are described by fast diffusion and
thus they can be displayed in equilibrium phase diagrams. However liquid-to-solid
transformations are not the main type of mineralogical transformations observed in
brick clays. Instead, solid-state transformations are the main mechanisms driving
high-temperature reactions in the common industrial practice. For solid-state transform-
ations, time is an additional important parameter due to the very slow solid-state diffu-
sion. Equilibrium may require exceptionally long periods of heating which are not
economically viable for the industry. Thus, metastable equilibrium phases may form
in preference to the stable state predicted by equilibrium phase diagrams. Moreover
the structural ceramic products may have optimum physical and mechanical properties
to meet certain specifications, although their mineralogical constituents need not have
reached equilibrium. An alternative approach to studying such systems is the kinetic
study of the mineralogical transformations in the clay matrix during firing, using
‘Time Temperature Transformation’ or TTT diagrams (Dunham, 1992; Dunham
et al., 2001).
The TTT diagrams describe the disequilibrium process of a transformation on a graph
of temperature, T, vs. time, t. A logarithmic time scale is generally used (Putnis &
McConnell, 1980). They are non-equilibrium phase diagrams, because they describe
processes as a function of thermal history. They can be used to describe spinoidal,
homogeneous and heterogeneous nucleation processes, as well as transformation
during increasing or decreasing T (Putnis & McConnell, 1980). Different products in
any given starting composition may be observed, resulting from different transformation
mechanisms depending on the actual temperature. Under certain circumstances, the
stability fields of mineral assemblages in various temperatures can be determined.
The TTT diagrams should be used with equilibrium phase diagrams.
A typical TTT diagram used to evaluate brick clays is shown in Figure 31. The stab-
ility fields of the various high-temperature phases under the specific experimental setup
 Industrial clays 403

depend both on temperature and time. Water absorption

An important application of TTT dia- Firing temperature
and time for
grams is that they may also contain associated commercial
information about important properties F product
of the fired clays (Fig. 31). In this E
manner the optimum set of firing temp-
erature and time can be estimated. These
composite diagrams are of industrial

Temperature (°C)
significance because often the end-
products may have physical properties
which are acceptable to the industry, D
although the firing cycle may be con-
siderably shorter than that dictated by C
the common industrial practice (Energy Target
mineralogy zone Fired shrinkage
Efficiency Practise, 1993). Hence the
TTT diagrams may suggest an energy-
efficient way of producing structural
clay products.
B

8. Concluding remarks A

Industrial clays have been utilized as

Time (h) log-scale
raw materials since the Palaeolithic era
and have been associated with important Fig. 31. Schematic TTT diagram of a brick clay.
advances in civilization of the human Phases A, B, C, D, E and F are hypothetical. The
species. Although initially they were diagram contains also information about some
important physical properties of bricks such as fired
used mainly as ceramic products, with
shrinkage and water absorption. The shadowed area
time, the range of applications has indicates firing conditions which yield end products
broadened and today they play more with optimum physical properties, whereas the solid
functional roles in various industrial circle indicates typical industrial firing temperature-
applications. Their importance is due time conditions. Using TTT diagrams may assist in
reducing production costs.
to the presence of clay minerals, which
have remarkable physical and chemical
properties. These properties can be modified with various inorganic and organic
reagents. New industrial products with specific properties and large added value
based on natural raw materials are currently produced by the chemical industry and
find application in various industries. A typical example is the formulation of nanocom-
posites. Also, synthetic clay minerals and clay-like materials such as LDH are prepared
and much effort is focused on the production of pure materials without admixtures.
From an industrial point of view, the assessment of industrial clay deposits is the
vital step in the characterization of the raw materials. Assessement involves three
steps: (1) identification of the clay minerals and associated minerals present; (2) deter-
mination of important physical and chemical properties of the clays; and (3) comparison
404 G. E. Christidis

with similar commercial products available for the same application. It is evident that
industrial clay research has a significant fundamental research component and an
applied research component, which are closely interrelated. The need to produce new
large-added-value products and/or synthetic products involves the coming together of
several scientific disciplines rendering the study of clays a clearly multidisciplinary
task. This has been evident from the presentation of the different types of industrial
clays in this chapter. The role of the clay scientist who deals with industrial clay deposits
is to combine the fundamental and the applied components and to collaborate with
scientists with different scientific backgrounds such as chemistry, soil science,
physics, materials science, etc. This is a challenging and fascinating task.

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