Bergaya F. Handbook of Clay Science

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showing a gradual increase of particle thickness, and a decrease in particle defect

density, lattice strain, and compositional variability (Peacor, 1992 ). The evolution of

dioctahedral clay minerals during diagenes is is smectite-random I/S-ordered

I/S-1M

illite-2M

muscovite.

The transformation of smectite to illite (‘illitization’) occurs when the sediment con-

taining I/S mixed-layer minerals is progressively buried in the basin, and the proportion

of illite layers increases. Potassium ions required for the illitization are supplied mainly

from the alteration of K

-feldspar within the sandstones, and from mica within the

shales (Furlan et al., 1996). In this sense the transformation of smectite to illite seems to

occur through a continuous series of interstratiﬁed I/S minerals. However, in the

Nankai Trough, Site 808 (an accretionary wedge where the Philippine plate subducts the

Eurasian plate) Masuda et al. (2001) observed that (i) authigenic K

-rich smectit e of

high Fe-content forms directly as an alteration product of volcanic glass at a depth of

E500 m below the seaﬂoor with no intermediate precursor; (ii) smectite is largely re-

placed by I/S minerals with Reichweite R ¼ 1 and minor illite and chlorite over a depth

of 500–700 m below the seaﬂoor. Most smectite and I/S minerals are derived from glass

alteration, rather than detrital as usually assumed; and (iii) I/S minerals with discrete

layer sequences (R ¼ 1) and illite co-exist, indicating discontinuities of the transforma-

tion from smectite to illite. This observation is similar to that of Dong et al. (1997),who

observed packets from Gulf Coast and other shales dominated by smectite, I/S minerals

with R ¼ 1(E50% illite), or illite. Thus, the smectite to illite transformation is not

simple and several mechanisms are possible.

On the other hand, S

rodon

(1999) used the concept of ‘fundamental particles’ to

interpret illitization. This concept was ﬁrst proposed by McHardy et al. (1982) and

further developed by Nadeau et al. (1984).S

rodon

suggested that the entities that

evolve during illitization are not mixed-layer crystals but fundamental particles.

The existence of interstratiﬁed I/S minerals in the diagenetic environment was

demonstrated by TEM, in apparent opposition to the ‘fundamental particle’ theory.

Thus, Nieto et al. (1996) were able to show by TEM the co-existence of chlorite, calcite,

and two I/S mixed-layer minerals (R ¼ 1 with 50% of illite layers and R>3 with more

than 90% illite), both thermodynamically incompatible, over distances of o100 nm

(Fig. 14.4). This assemblage suggests that the evolution of smectite to illite is an Os-

twald-type ripening process (Morse and Casey, 1988) in which phases reach equilib-

rium through progressive steps, with less heterogeneity and closer to the stable phases.

Therefore, two mechanisms of smectite-to-illite transformation are possible (i)

through dissolution- recrystallization without involving mixed-layer minerals (Inoue

and Kitagawa, 1994; Clauer et al., 1997; S

rodon

, 1999; S

rodon

et al., 2000) mostly

applicable to bentonite transformation and (ii) through I/S mixed-layer phases

(Eberl and S

rodon

, 1988; Inoue et al., 1988; Christidis, 1995; Nieto et al., 1996),

applicable to shales and some hydrothermally altered bentonites. Mechanism (ii)

involves solid-state layer-by-layer transformation, including random interstratiﬁc-

ation, transition to ordered mixed-layer minerals, and development of long-range

order (Bethke and Altaner, 1986).

14.1. Geological Environments for Clay Formation 1143

Another interesting ﬁnding, illustrating the structural and chemical heterogeneity

of these I/S mixed-layer minerals, and their evolution in a basin, was described by

Drits et al. (2002). In oil-source shales from the North Sea (Upper Jurassic), oil

generation takes place simultaneously with diagenetic transformation of I/S. A link

between these two reactions can be the foll owing: NH

released during maximum oil

generation is ﬁxed as NH

cations in the NH

-bearing mica (tobelite) layers formed

from smectite or vermiculite layers of I/S minerals. This process leads to the for-

mation of four-component illite–tobelite–smectite–vermiculite (I-T-S-V) minerals in

a diagenetic interval, called ‘tobelitization window’. Drits et al. (2002) also found

that the amount of tobelite layers and ﬁxed NH

increased wi th diagenesis, while

the amount of K

ﬁxed in the I/S mixed-layer minerals as well as the proporti on of

-bearing illite layers remained constant, irrespective of sample location and depth

and degree of diagenetic transformation. They concluded that in these oil-source

shales, diagenesis is accompanied not by smectite illitization, as is generally the case,

but by smectite (or vermiculite) tobelitization. The constant proportion of illite

layers and the constant number of cations are considered as evidence of solid-phase

transformation of the I-T-S-V because any other mechanism would destroy this

Fig. 14.4. (A) TEM image of a sample UG-17 including chlorite, I/S mixed-layer minerals

with R ¼ 1 and R>3, and calcite. (B) lattice fringe image of the area marked in (A) showing

well-deﬁned individual packets of I/S minerals with R ¼ 1 and R>3. From Nieto et al. (1996).

Chapter 14: Genesis of Clay Minerals1144

constancy. Tobelitization of smectite in I/S minerals is probably typical of all oil-

source shales.

Contemporaneous with the smectite-to-illite transformation, signiﬁcant chlorite is

formed from smectite (through loss of Fe

and Mg

) or kaolinite (by diff usion of

and Mg

in the structure). The change of trioctahedral smectite to chlorite in

a Mg-rich environment can give a mixed-layer structure of the corrensite type.

During burial diagenetic I/S transformation, Mg

cations can also be ﬁxed in the

smectite layers. The formation of brucite-like sheets can give rise to an I/S-dioc-

tahedral chlorite structure (tosudite) (Lindgreen et al., 2002). Later diagenesis may

lead to neoformation of tosudite and trioctahedral chlorite-berthierine.

Diagenetic changes convert muds into lithiﬁed mudrocks, that is, mudstone-

shale-slate. The clay mineral reactions approach chemical equilibrium through

intermediate metastable products. Reactions rates in clays heated by igneous intru-

sion temperatures are many orders of magnitude faster and promote direct trans-

formation of smectite to muscovite without intermediate phases (Merriman, 2002).

Clay mineral reactions require relatively small activation energies because the new

minerals formed have similar compositions and structures to the reactants. As a

result, they can form sequences of reactants and products or reaction series

(Merriman and Peacor, 1999). Fig. 14.5 shows the reaction series recognized in

British Lower Paleozoic mudrocks (Merriman and Roberts, 1985, 2001; Fortey,

1989; Roberts et al., 1990, 1991), according to the basic phyllosilicate structure and

Fig. 14.5. Clay mineral reaction series observed in British Lower Paleozoic slate belts. Re-

action proceeds from top to bottom, indicated by heavy arrows. Diagonal arrows indicate

products contributed from one series to another. From Merriman (2002).

14.1. Geological Environments for Clay Formation 1145

increasing diagenetic-metamorphic conditions. For 2:1 dioctahedral clay minerals,

the series includes basically the smectite illitization, and considering the Na/K ratio,

this series evolves to muscovite (K

) or paragonite (Na

). For the 2:1 trioctahedral

clay minerals (Ca

,Mg

smectites), the smectite chloritization is the fundamental

transformation, usually through corrensite. Finally, for the most simple clay min-

erals, 1:1 dioctahedral ones (kaolinite), the basic transformation is the formation of

more ordered polytypes (mainly dickite), and the possibility of pyrophyllite forma-

tion in the anchizone–epizone conditions. These series described for the British

Lower Paleozo ic were also recognized in other belts, that is, in the Cantabrian

Cordillera (Spain) (Gala

n and Aparicio, 1980 ).

Clay mineral assemblages found in slates can also be related to the geotectonic

setting and thermal history of the basin (Merriman, 2002). Extensional basins are

characterized by high-heat ﬂow (>35 1C/km) and hydrothermal activity, an d tend to

have a greater diversity of transformed clay minerals containing both K

- and Na

rich products of the 2:1 dioctahedral reaction series. Pyrophyllite, rectorite, corren-

site, and paragonite can be recognized but kaolinite is rare. In contrast, clay

assemblages that evolve in convergent low-heat ﬂow basins generally contain fewer

mineral species, simply K

-white mica and chlorite, and Na

-mica and pyrophyllite

are rare or absent. In extensional basins Na

ions may come from low-temperature

mixing of hydrothermal ﬂuids and sea water. Such ﬂuids are unavailable in con-

vergent basins because of a lack of volcanic activity.

Over the last few decades, clay mineral assemblages and some parameters, such as

the Ku

bler index for illite, smectite–il lite reactions, chlorite composition, and the

Arkay index for chlorite, were used as geothermometers and geobarometers for

assessing the conditions in basins that were submitted to burial diagenesis and very

low-grade metamorphism. However, as Essene and Peacor (1995) pointed out, the

use of such systems and parameters does not provide accurate geother mometers

because most of them are not based on equilibrium reactions.

Recent research demonstrated that most clay minerals are out of equilibrium with

their environment. Clay mineral transformations are equivalent to Ostwald-ripening

steps, driven by the potential for minimum free energy, and clay minerals may reach

equilibrium only in very low or low-grade metamorphism. The application of clay

mineral assemblages is based on physical calibration as observed for natural clay-bearing

systems. The repeatability of clay mineral assemblages in space and time, as a function of

increasing P–T conditions, is often assumed to imply equilibrium. However, this ob-

servation is not evidence for equilibrium although it is a necessary condition (Essene and

Peacor, 1995). Most clay mineral associations are metastable, and the chemical reaction

kinetics depend on a great variety of parameters and circumstances, such as time, ﬂuid/

rock ratio, tectonic history (deformation), starting material, pressure, and temperature.

On the other hand, ‘retrograde’ reactions can occur where I/S mixed-layer min-

erals form by replacement of metamorphic illite (Jiang et al., 1990), smectite derives

from chlorite (Nieto et al., 1994), or smectite and highly expandable I/S minerals

from illite (Zhao et al., 1999).

Chapter 14: Genesis of Clay Minerals1146

Another interesting formation of clay minerals in this environment is due to the

devitriﬁcation of volcanic ash, leading to the formation of massive smectite (bento-

nite). The process seems to consist of slow ionic diffusion caused by a slow burial

rate, and hence a low compaction rate in an open chemical system. In many bento-

nite beds I/S mixed-layer minerals with high K

content are found in contact with

the enclosing sed imentary rock, indicating that the diffusion process gradually

transforms smectite into I/S mixed-layer minerals over a long period of time.

A special case of great practical interest is the formation by diagenesis of clay

minerals in sandstone pores. Connected pores in sandstone create permeability.

Since sandstones can serve as possible hydrocarbon reservoirs, porosity and perme-

ability are two important factors. Both parameters can decreas e due to the diagenetic

formation and growth of clays in pores and in the passages connecting them, hin-

dering the passage of ﬂuids and reducing the economic value of the reservoir. Ka-

olinite is the most frequent clay mineral formed, often as vermicular ‘books’ of

stacked layers but I/S interstratiﬁed clay minerals, berthierine, and lath-shaped illite

are also common.

Diagenetic kaolinite can form by ﬂushing sandstones with meteoric water. Ka-

olinite-feldspar assemblages are stable until 120–140 1C and react to form illite.

Huang et al. (1986) conﬁrmed the role of ﬂuid/rock ratio in altering feldspars into

kaolinite or illite. Ehrenberg et al. (1993) and Ruiz Cruz and Andreo (1996) doc-

umented the kaolinite-dickite transition at E120 1C.

According to Beaufort et al. (1998) the kaolinite-to-dickite reaction proceeds by

gradual structural changes concomitant with crystal coarsening and a ch ange from

booklet to blocky morphology (Fig. 14.6). The crystallization of dickite follows two

different pathways: (i) accretion of new material from the dissolution of unstable

kaolinite particles and/or detrital minerals, or coarser metastable kaolinite and (ii)

neoformation of ordered dickite by dissolution–crystallization process. They suggest

that the kaolinite-to-dickite reaction has a petrogenetic signiﬁcance, useful for re-

constructing the burial diagenesis of sedimentary basins, and for petroleum and gas

exploration.

Lanson et al. (2002) reviewed the authigenic minerals formed in sandstones during

burial diagenesis. While early precipitation of kaolinite is generally related to ﬂush-

ing by meteoric waters and as a consequence of feldspar dissolution in an open

system, subsequent diagenetic kaolinite-to-dickite transformation probably results

from invasion by acidic ﬂuids of organic origin. This conversion is kinetically con-

trolled and dickite is the most stable polytype in sandstone. However, dickite is not

formed exclusively by diagenetic transformation of kaolinite. It can also form from

dissolution of K

-feldspars and other Al-rich sil icates with increasing temperature,

probably in the presence of organic acids. On the other hand, illite crystallization at

the expense of kaolinite implies that an energy barrier is overcome either by an

increase in K

activity ratio in solution or by a temperature increase. Illitization

of kaolinite is not necessa rily coupled to dissolution of K-feldspars as usually de-

scribed, but a potassium source is required.

14.1. Geological Environments for Clay Formation 1147

Fig. 14.6. Schematic model of the kaolinite-to-dickite reaction involving both morphological

and structural changes as a result of water–rock interaction for increasing burial depths in

sandstone reservoirs. Arrows indicate the transfer of matter due to material redistribution

involved in the dissolution–crystallization processes. After Beaufort et al. (1998).

Chapter 14: Genesis of Clay Minerals1148

In summary, we can say that early burial diagenesis and the passage to low-grade

metamorphism (anchimetamorphism) transform clay minerals into other minerals

until environmental equilibrium is reached. Typical clay minerals formed during

diagenesis are the I/S interstratiﬁed minerals, which are intermediate phases in

smectite illitization. This transformation was studied in much detail, because the

reaction can indu ce petroleum migration and build up geopressures by liberating

water from smectite interlayers. Soluble SiO

,Fe

, and Mg

liberated during

illitization may cause cementation in the pores of sedim entary rocks (see Altaner and

Bethke (1988) and references therein). With increasing burial other phyllosilicates

can be formed, such as pyrophyllite and paragonite. These minerals are in metastable

equilibrium with I/S mixed-layer minerals in the anchimetamorphism zone. Chlorite

can also be formed from smectite or kaolinite. The ﬁrst stage of transformation can

give rise to corrensite. On the other hand, massive smectite can be formed from the

devitriﬁcation of volcanic ash.

Kaolinite is a typical clay mineral formed in sandstone pores by direct preci-

pitation. At 120–140 1C it can transform into illite (by reaction with feldspars) or

dickite. Dickite can also form directly from dissolution of feldspars (Lanson et al.,

2002).

XRD was fundamental to following the evolution of clay minerals in a basin as a

function of depth, time, and tectonics. The application of other instrumental tech-

niques, notably HRTE M, provided much valuable information, and novel insights

into clay mineral formation and transformation. More studies are still necessary, in

particular, on the relationship between these process es and their geotectonic setting

(Merriman, 2002).

14.1.4. Hydrothermal Alteration

As mentioned before, intrusion of plutonic rocks and the deposition of volcanic

events lead to wall-rock alteration and transformation of some rock-forming min-

erals into clayey materials. Clays can be genetically associated with late-stage de-

posits (chloritization), pegmatites, albitite-greissen deposits (musco vitization,

chloritization), and hydrothermal veins. The veins are most important because crys-

tal-chemical variations of clay minerals are particularly related to the ﬂuid compo-

sition and controlled by the type of altered rock.

Hydrothermal alteration involves water–rock interaction at temperature above

E50 1C. In general, the altered material forms at higher temperatures than those of

the hydrothermal ﬂuids, hence the minerals are unstable in the presence of water at

the temperatures considered. The ﬂuids are aqueous and contain gaseous compo-

nents and dissolved materials.

Hydrothermal systems are open to components where convecting ﬂuids produce

maximum water/rock ratios, and depths may be shallow (Henley, 1985). Very

different clay minerals may form at the same time, but at different temperatures. At

300–400 1C the clay alteration facies include sericite or mica, potassium feldspars and

14.1. Geological Environments for Clay Formation 1149

chlorite, very simila r to that of the greissen deposits. As temperature decreases illite

and kaolinite predominate with some I/S mixe d-layer minerals. At still lower tem-

peratures smectite or kaolinite can be the most important clay minerals. The most

striking examples of hydrothermal alteration are the production of pure kaolin and

silica. These minerals can occur on the scale of kilometers, and are of economical

interest.

In epithermal ore deposits, found in the upper (o1–2 km) geothermal systems

where predominantly meteoric ﬂuids convect, interaction with the host rock takes

place at temperatures of 200–300 1C. In general, these systems are relatively short-

lived, remaining active over periods of hundreds of thousands of years or less (Henley

and Ellis, 1983). Nonetheless, interesting clay (e.g., kaolin) deposits can be produced.

The following three cases illustrate the diversity of clay minerals formed in this

environment under the inﬂuence of altered rock type, P and T, and ﬂuid chemistry.

In the Golden Cross epithermal Au–Ag deposit at Waihi, New Zealand hosted by

andesitic and dacitic lavas, hydrothermal alteration of the rocks gives rise to quartz,

adularia, calcite, clay minerals (chlorite, illite, kaolinite, I/S mixed-layer minerals,

and smectite), and others. Accor ding to Tillick et al. (2001) all phyllosilicates are

neoformed directly from the ﬂuid s. I/S mixed-layer minerals (R ¼ 1) and micas form

simultaneously without intermediate series, as Masuda et al. (2001) observed for the

burial diagenes is of mudsto nes. The illite content in the I/S minerals increases with

depth and proximity to the central vein system. Smectites co-exist with (R ¼ 1),

rather than (R ¼ 0), I/S minerals.

The second case is the hydrothermal alteration of ben tonites of the Tsantili de-

posit in Milos Island, Greece. The original Wyoming-type montmorillonite was il-

litized along a 40 m vertical proﬁle, forming I/S mixed-layer minerals with different

proportions of expandable layers and different ordering (Fig. 14.7). The alteration is

characterized by a massive incorporation of K

and removal of Si

,Na

,Ca

,andFe

from the rock. The hydrothermal alteration is probably associated

with the emplacement close to the parent bentonite of barite veins at o200 1C.

Illitization is controlled by temperature, K

availability, and ﬂuid/rock ratio. It can

be considered to be the result of K

-metasomatism. As the potassium is derive d

from the dissolution of authigenic K

-feldspar, the I/S mineral particles are appar-

ently formed by solid-state transformation. According to Christidis (1995), the ill-

itization and spatial distribution of expandable miner als could be used as a guide for

the exploration of mineral deposits.

The third case concerns the intrusion of a magmatic rock in a sedimentary series,

leading to contact metamorphism in the host rock. Subsequent hydrothermal alter-

ation usually yields clayey materials. For instance, in the Priego de Co

rdoba area

(Subbetic zone of the Betic Cordilleras, Spain), a lacolith of stratiform dolerite

intruded marly sediments at quite shallow depths below the ocean ﬂoor during the

intracontinental rifting phase. In the ﬁrst stage, contact metamorphism caused crys-

tallization of Ca

-silicates at about 500 1C. During a process of hydrother mal

alteration superimposed on the contact aureole, chlorite was formed in the area

Chapter 14: Genesis of Clay Minerals1150

closest to the volcanic rocks in a ﬁrst step. In the zone farthest from the contact area

corrensite was formed, but a later cooling phase led to crystallization of saponite in

the host rock, and di- and trioctahedral smectites inside the volcanic rock (Abad

et al., 2003 ). Temperature can decrease from E300 1C (chlorite) to 270–200 1C

(chlorite–smectite interstratiﬁcations) in the prograde phase, and to 180–130 1C

(saponite) in the retrograde phase (Stakes and O’Neil, 1982; Dudoignon et al., 1997).

Deep-sea hydrothermal alteration is rather different from the alteration described

above. Hydrothermal alteration in basic and acidic rocks tends to produce Mg-and

Fe-free clay mineral assemblages, whereas alteration of deep-sea basalts at very low

temperatures yields Fe-rich clays. With decreasing temperatures from E300 1C the

Fig. 14.7. (A) Schematic cross-section of the south sector of the Tsantili bentonite deposit

(Greece). Ii (i ¼ 0  85) is the illite content of the I/S mixed-layer minerals. Ca in ¼ presence

of calcite incompatible with jarosite; Ja in ¼ presence of jarosite incompatible with calcite; (B)

Simpliﬁed model proposed for the illitization of Wyoming-type montmorillonite in the Tsantili

deposit in relation to barite vein (vertical line represents the cross-section of (A)). K,

Heat ¼ K

availability and temperature are decreasing from the deposit. After Christidis

(1995).

14.1. Geological Environments for Clay Formation 1151

sequence of clay minerals in the altered basalts is serpentine-saponit e -celado-

nite-nontronite. The last mineral is found in the ‘weathering’ zone on basalts, that

is, at the sea water–rock interface.

In summary, clay minerals formed by hydrothermal activity are not related to

sedimentary layers or weathering crusts. In contrast with the other environm ent

described, zonation from the heat source is a characteristic of this complex envi-

ronment, which can be very variable over short distances. Clay minerals formed can

have an economic interest, and a detailed study can provide information about their

use as indicators for exploring ore deposits.

14.2. ORIGIN OF CLAY DEPOSITS OF ECONOMIC INTEREST

Commercial clays, that is, clays used as raw material in industry rank as the

leading industrial rocks in both tonnages and total value. Clay is a rock historically

linked to humans in their agricultural, industrial, and cultural development. It is a

relatively abundant raw material with many different applications, encompass ing a

wide spectrum of products. These range from low-value materials, produced with a

minimal degree of processing (e.g., certain plastic clays for bricks) to high added-

value materials (e.g., reﬁned kaolins for paper coating) (see Chapter 10.1). About

90% of production comprises building clays, or miscellaneous clays for structural

products. Only 10% of deposits are largely composed of a clay mineral (kaolin,

bentonite, fuller’s earths, etc.); these are generally referred to as ‘special clays’.

The origin of commercial clays is of great basic interest because the uncommonly

massive concentrations of very pure clay material are produced under special con-

ditions. Geological criteria shown by genetic studies can be used for the exploration

and mining of new deposits.

Here we give a brief overview of the genesis of the principal clay deposits, other

than structural clays. The latter are sedimentary formations (marine and continental)

abundantly distributed throughout the world from the Carboniferous to the Cenozoic.

14.2.1. Kaolins

Kaolinite can be formed by weathering (residual kaolins) and hydrothermal activity

(hydrothermal kaolin), or occur as an authigenic sedimentary mineral. Residual and

hydrothermal kaolins, called primary kaolins, are formed in situ by surface and

underground waters or hydrothermal ﬂuids. Sedimentary kaolins, called secondary

kaolins, are composed of kaolinized material from a source area that was eroded,

transported, and deposited in a continental or coastal environment. Most kaolinites

in sedimentary kaolins are formed before deposition, in a primary environment, but

some can be authigenic after the deposition of the arkosic or sandy material, usually

by the action of underground water.

Chapter 14: Genesis of Clay Minerals1152