Clay and Clay Mineralogy

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GEOLOGIAN TUTKUSKESKUS M19/3232/2008/41 30.6.2008 Espoo

CLAY AND CLAY MINERALOGY

PHYSICAL – CHEMICAL PROPERTIES AND INDUSTRIAL USES Dr. Thair Al-Ani and Dr. Olli Sarapää

Kaolinite books from Litmanen

GEOLOGICAL SURVEY OF DOCUMENTATION PAGE

FINLAND Date / Rec. no.

Authors Type of report Thai Al-Ani and Olli Sarapää Commissioned by

Title of report Clay and clay mineralogy

Abstract

This book on Clays and Clay Minerals is related research reports conducted by the Geological Survey of Finland. One can look at the book as dealing with seven chapters: (l) Identification and occurrence of clays, (2) structure and composition of the clay minerals, (3) physical and geochemical properties of clay minerals, (4) methods of clay minerals investigation, (5) formation and alteration of clay minerals, (6) industrial uses of clays, (7) clays in Finland. Chapter 1 and 2 contain briefly summarizes the identity of clay minerals and their occurrence of clays and also describe a review of the structure and composition of the clay minerals. The subject of chapter 3 is on physical and geochemical properties of clay minerals and processing techniques are described. Also describe the equilibrium properties and ionic exchange of swelling clay minerals. Chapter 4 presents the details of the methods used for the identification and quantification of aluminosilicates include X-ray diffraction, electron microscopy, energy-dispersive X-ray analysis, infrared spectroscopy, differential thermal analysis, and scanning electron microscope. This chapter also including many practical examples for diffract grams of many clay samples from different deposits in Finland. Chapter 5 in a short chapter, generalizes about formation and alteration of clay minerals, outlines some generalizations regarding clay occurrences and environments of deposition.

Chapter 6 details the industrial uses of clay minerals especially for specific applications of kaolins, smectites, and palygorskite and sepiolite. This chapter also describes many clay deposits on a world scale; clay is of major economic significance, touching virtually every aspect of our everyday lives, from medicines to cosmetics and from paper to cups and saucers. It is very difficult to over-estimate its use and importance. Chapter 7 by Olli Sarapää presents the details examine of clays in Finland and summarize their origin and their mineralogical, chemical and industrial properties.

The value of this handbook to those outside the short course derives from its presentation of some interesting examples of diagenetic changes involving clay minerals.

The book contains much of the analytical data utilized analytical techniques including X-ray diffraction, infrared spectroscopy and electron microscopy. The figures and tables make this book very understandable and the references are adequate.

Keywords Clay, clay minerals, kaolinite, Virtasalmi, and Viittajänkä.

Geographical area GTK Espoo

Report serial Archive code M19/3232/2008/41

Total pages Language Price Confidentiality 94 english Unit and section Project code

Signature/name Signature/name Al-Ani Thair

Contents

Documentation page

CHAPTER ONE 1

1 INTRODUCTION TO CLAY MINERALOGY 1 1.1 Course Objectives 1 1.1.1 Preface 1 1.1.2 Clay 1 1.2 Identity of clay minerals 2 1.3 Occurrences of clay 4

CHAPTER TWO 5

2 STRUCTURE AND COMPOSITION OF THE CLAY MINERALS 5 2.1 Coordination Polyhedral 5 2.1.1 Tetrahedron 6 2.1.2 Octahedron 7 2.1.3 Layer types 8 2.2 Clay mineral Classification 10 2.2.1 Allophane 12 2.2.2 Kaolin Minerals 13 2.2.3 Smectite Minerals 16 2.2.4 Vermiculite 18 2.2.5 Illite minerals 21 2.2.6 Chlorite Minerals 23 2.2.7 Palygorskite and sepiolite 25 2.2.8 Mixed-layer clay minerals 28

CHAPTER THREE 29

3 PHYSICAL AND GEOCHEMICAL PROPERTIES OF CLAYS AND CLAY MINERALS 29 3.1 The main properties of particular clay minerals 29 3.1.1 Kaolin 29 3.1.2 Smectite (Bentonite) 30 3.1.3 Illite 31 3.1.4 Other clays 31 3.2 The Geochemistry of clay minerals 32 3.2.1 Equilibrium adsorption and Ion exchange 32 3.2.2 Surface charge properties 34 3.3 Measurement of Cation Exchange Capacity 37 3.4 Swelling Properties of smectite 39

CHAPTER FOUR 40

4 METHODS OF CLAY MINERALS INVESTIGATION 40 4.1 X-Ray Diffraction Experiment 40 4.2 Clay preparation for XRD 41 4.2.1 Procedure of Separation clay fraction from bulk sediments 41 4.2.2 Glycolation 43 4.2.3 Heating 43 4.2.4 Differentiation between Clays 43 4.3 Semi-quantitative analysis of clay minerals 45 4.4 Infrared spectroscopy 47 4.5 The Scanning Electron Microscope (SEM) 49 4.6 Transmission electron microscopy (TEM) 50

CHAPTER FIVE 51

5 FORMATION AND ALTERATION OF CLAY MATERIALS 51 5.1 Introduction 51 5.2 The clay cycle 52 5.3 Geological origin of clay minerals 53

CHAPTER SIX 57

6 INDUSTRIAL USES OF CLAYS 57 6.1 Kaolin 57 6.1.1 Paper industry 60 6.1.2 Other Applications 62 6.2 Smectite (Bentonite) 62 6.2.1 Bentonite barriers in sealing nuclear waste 64 6.2.2 Other uses of Smectite (bentonite) 64 6.3 Palygorskite – Sepiolite 65 6.4 Clay application for the future 67

CHAPTER SEVEN 68

7 CLAYS IN FINLAND 68 7.1 Clay minerals in fracture zones of the Finnish bedrock 68 7.2 Kaolin occurrences in Finland 70 7.2.1 Virtasalmi kaolins 70 7.2.2 Kainuu kaolins 76 7.2.3 Viittajänkä kaolin deposit 79 7.2.4 Other kaolins 811 7.2.5 Glacial and Postglacial clay deposits in Finland 81 7.2.6 Clay minerals in tills 82

8 REFERENCES 844

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CHAPTER ONE

1 INTRODUCTION TO CLAY MINERALOGY

1.1 Course Objectives

1.1.1 Preface This book contains material used to support a graduate-level course on clay minerals at Helsinki University and also support other researcher for clays. The content is presented in seven chapters that cover basic mineralogy and classification; physical and chemical properties (including aqueous solubility and ion exchange); X-ray powder diffraction methods for the identification and quantification of clay mineral assemblages; formation of clay minerals and geologic origin; industrial clays (including waste isolation uses) . The book covers varies aspects of clays such as clays in Finland, which is written by Dr. Olli Sarapää in chapter eight. This chapter examine the diversity of clays in fracture zones of Finnish bed rocks and summarize their origin and their mineralogical, chemical and industrial properties. The general objectives of this book are to give a greater understanding of clay mineral reactions in the environment and the processes controlling their geologic distribution. It produces an increased awareness of the relationship between structural/chemical characteristics of the diverse clay minerals present in rocks, soils and sediments and their physical and chemical properties. The contents of the book are useful in diverse fields of scientific and technical investigation.

1.1.2 Clay The meanings of the terms, “clays” and “clay minerals”, are important to be distinguishing before starting to read the book. A very brief note on general aspects of clay can be explained as follows:

The term "clay" refers to a naturally occurring material composed primarily of fine-grained minerals, which is generally plastic at appropriate water contents and will harden when dried or fired. Clay usually contains phyllosilicates, it may contain other materials that impart plasticity and harden when dried or fired. Associated phases in clay may include materials that not impart plasticity and organic matter.

Clay and sand both indicate a specific grain size; however, it is often used to refer to a specific mineralogical composition of sediments. Figure 1 shows the classification of siliciclastic sediments (unconsolidated, loose) that are based on average grain size. We advise to use it only for grain size. An important point in this figure is that the boundary between sand and silt is 0.06mm and smaller than 0.004mm is clay. The current ISO (International Organization for Standardiza- tion) Standard 14688:1996 placed the boundary at 0.06 mm between sand and silt (Geological Society, London, 2006).

The term “clay mineral” refers to phyllosilicate minerals and to minerals which impart plasticity to clay and which harden upon drying or firing. Clay minerals are layer silicates that are formed usually as products of chemical weathering of other silicate minerals at the earth's surface. They are found most often in shales, the most common type of sedimentary rock. In cool, dry, or tem-

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perate climates, clay minerals are fairly stable and are an important component of soil. Clay minerals act as "chemical sponges" which hold water and dissolved plant nutrients weathered from other minerals. This results from the presence of unbalanced electrical charges on the surface of clay grains, in which some surfaces are positively charged (and thus attract negatively charged ions), while other surfaces are negatively charged (attract positively charged ions). Clay minerals also have the ability to attract water molecules. Because this attraction is a surface phenomenon, it is called adsorption (which is different from absorption because the ions and water are not attracted deep inside the clay grains). Clay minerals resemble the micas in chemical composition, except they are very fine grained, usually under microscope. Like the micas, clay minerals are shaped like flakes with irregular edges and one smooth side. There are many types of known clay minerals. Some of the more common types and their economic uses are described here:

Figure 1. Grain size classification scheme

1.2 Identity of clay minerals

Kaolinite: This clay mineral is the weathering product of feldspars. It has a white, powdery appearance. Kaolinite is named after a locality in China called Kaolin, which invented porcelain (known as china) using the local clay mineral. The ceramics industry uses it extensively. Be- cause kaolinite is electrically balanced, its ability of adsorb ions is less than that of other clay minerals. Smectite: This clay mineral is the weathering product of mafic silicates, and is stable in arid, semi-arid, or temperate climates. It was formerly known as montmorillonite. Smectite has the ability to adsorb large amounts of water, forming a water-tight barrier. It is used extensively in

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the oil drilling industry, civil and environmental engineering (where it is known as bentonite), and the chemical industry. There are two main varieties of smectite, described in the following: Sodium smectite: This is the high-swelling form of smectite, which can adsorb up to 18 layers of water molecules between layers of clay. Sodium smectite is the preferred clay mineral for drilling muds, for creating a protective clay liner for hazardous waste landfills to guard against future groundwater contamination, and for preventing seepage of groundwater into residential base- ments. Sodium smectite will retain its water-tight properties so long as the slurry is protected from evaporation of water, which would cause extensive mud cracks. As a drilling mud, sodium smectite mixed with water to form a slurry which performs the following functions when drilling an oil or water well: 1) lubricates the drill bit to prevent premature wear, 2) prevents the walls of the drill hole from collapsing inwards, 3) suspends the rock cuttings inside the dense mud so that the mud may pumped out of the drill hole, and 4) when the dense mineral barite is added to drilling mud, it prevents blowouts caused by internal pressure encountered during deep drilling. So- dium smectite is also used as commercial clay absorbent to soak up spills of liquids. Calcium smectite: The low-swelling form of smectite adsorbs less water than does sodium smectite, and costs less. Calcium smectite is used locally for drilling muds. Illite: Resembles muscovite in mineral composition, only finer-grained. It is the weathering product of feldspars and felsic silicates. It is named after the state of Illinois, and is the dominant clay mineral in mid-western soils. Chlorite: This clay mineral is the weathering product of mafic silicates and is stable in cool, dry, or temperate climates. It occurs along with illite in mid-western soils. It is also found in some metamorphic rocks, such as chlorite schist. Vermiculite: This clay mineral has the ability to adsorb water, but not repeatedly. It is used as a soil additive for retaining moisture in potted plants, and as a protective material for shipping packages. Palygorskite (attapulgite): Palygorskite is synonymous terms for the same hydrated Mg-Al silicate material. The name specified by the International Nomenclature Committee is palygorskite. However, the name attapulgite is so well established in trade circles that it continues to be used by many producers and users. This mineral actually resembles the amphiboles more than it does clay minerals, but has a special property that smectite lacks - as a drilling fluid, it stable in salt water environments. When drilling for offshore oil, conventional drilling mud falls apart in the presence of salt water. Palygorskite is used as a drilling mud in these instances. Incidentally, palygorskite is the active ingredient in the current formula of Kaopectate.

Clay minerals form an important group of the phyllosilicates or sheet silicate family of minerals, which are distinguished by layered structures composed of polymeric sheets of SiO4 tetrahedral linked to sheets of (Al, Mg, Fe)(O,OH)6 octahedral. The geochemical importance of clay minerals stems from their ubiquity in soils and sediments, high specific surface area, and ion exchange capacities. Clay minerals tend to dominate the surface chemistry of soils and sediments. Fur- thermore, these properties give rise to a wide range of industrial applications throughout the history of mankind. The use of clay for mainly clay figures, pottery and ceramics was already known by primitive people about 25000 years ago (Shaikh and Wik, 1986). Today clay is an important material with a large variety of applications in ceramics, oil drilling, liners for waste disposal, and the metal and paper industry. Clay is furthermore used as adsorbent, decolouration agents, ion exchanger, and molecular sieve catalyst (Murray, 1991). Despite their importance, the clay minerals form a difficult group of minerals to study due to their small size, variable structural composition, and relative slow kinetics of formation and alteration.

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1.3 Occurrences of clay

Sedimentary rocks only make up 5% of the Earth's crust, but cover about 80% of the surface of the earth in which clays (including shales) form well over 40% of the sedimentary rocks. The raw material for sedimentary rocks comes from weathering. If we look at the volume of material at the earth's surface (Fig. 2), we see that clay minerals constitute about 16% of its total. 20 km is considered the surface of the earth because it is the region from which we extract natural resources (and dump our waste). Clay sediments are collected by the agencies of water (e.g. marine clays, alluvial clays, lacustrine clays), wind (Aeolian clays), or ice (e.g. glacial clay, till or boulder clay, as most clays in Finland). The majority of the common sedimentary clays, however, are the marine deposits typically comprising mixtures of coarser material with clay in which the clay mineral, illite, usually predominant (see chapter two for description of clay minerals). Clay mineral-rich deposits can be formed in two other principle ways:

• by weathering of parent minerals in situ to form a clay rich residual soil in which the clay mineral kaolinite frequently predominates, especially common in landscapes undergoing tropical weathering, and

• by ascending fluids, i.e. by hydrothermal alteration of the host rock. Cornish china clay is a good example, the feldspar of the local granite having been converted mainly into clay minerals of the kaolinite group. For full discussion see chapter five.

Figure 2. The volume of material at the earth's surface

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CHAPTER TWO

2 STRUCTURE AND COMPOSITION OF THE CLAY MINERALS

The structure and composition of the major clay minerals, i.e. kaolins, smectites, vermiculite, illite, chlorite and Palygorskite-sepiolite, are very different even though they are each comprised of octahedral and tetrahedral sheets as their basic building blocks. The arrangement and composition of the octahedral and tetrahedral sheets are account for most of the different in their physical and chemical properties.

2.1 Coordination Polyhedral

Ionic bonds between oxygen and the cations, aluminum, silicon, magnesium, potassium and sodium, are the most important bonds in silicate minerals, including the clay minerals. The number of oxygen atoms surrounding the smaller cations in these mineral structures can be predicted from simple geometrical relationships. If the ions are considered as rigid spheres, then the ratio of the cation radius to the anion radius determines how many anions can be in contact with a cation, the coordination number. If the cation is very small, then only two oxygen anions may be in contact with it. As the relative size of the cation increases, more oxygen atoms may be coordinated. The limiting radius ratios, the coordination number, and the geometric configuration of the oxygen ions are showed on Table (1).

Table 1. The limiting radius ratios, the coordination number, and the geometric configuration of the oxygen ions R Cation / R Anion Coordination No. Arrangement Of Oxygen

< 0.16 2 corners of a triangle

0.16 – 0.23 3 opposite one another

0.23 – 0.41 4 corners of tetrahedron

0.41 – 0.73 6 corners of octahedron

0.73 – 1.00 8 corners of cube

> 1.00 12 close packed spheres

Clay minerals are hydrous aluminosilicates, thus the O-2(1.40Å), OH-1 (1.41Å), Al+3 (0.55Å), and Si+4 (0.41Å) ions are the most important ones to consider in developing an understanding of their crystal structure. From these approximate ionic radii, radius ratios can be used to predict coordination numbers. The RC: RA ratios for Si: O and Al: O are 0.29 and 0.39, respectively, which fall within the 0.23-0.41 limits predicted for tetrahedral coordination. Each silicon, or aluminium ion, should occupy the centre of a tetrahedron with an oxygen ion at each of the four corners. For Al, the radius ratio is very near to the lower limits of the range for predicted octahe-

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dral coordination (0.41-0.73) and it is frequently located at the centre of an octahedron with six oxygen ions at the corners. Aluminium may exhibit four-fold or six-fold coordination with oxygen or hydroxyl groups.

Magnesium and iron are two other common elements with radius ratios favouring octahedral (six-fold) coordination. Radius ratios are only approximate guides to structural arrangements as the actual coordination of the ions may influence their radii.

In addition to geometry, there are some other general considerations that can be applied to arrangements of atoms in the structures of minerals. Empirically, those relationships leading to the most favourable energy conditions in ionic bonded crystals can be represented by Pauling’s Rules (Pauling, L. 1960). You will observe how some of these rules are manifest when we consider structures in more detail. 1. Coordination polyhedral of anions form about a central cation. The coordination number and form are determined by the radius ratios (discussed above). 2. In stable structures, the total strength of the bonds from adjacent cations reaching the anion is equal to its valence. 3. Sharing of polyhedral edges and faces reduces the stability of ionic structures. 4. Cations with high charge and low coordination number do not usually share polyhedral elements. 5. The number of different kinds of constituents in a crystal tends to be small.

2.1.1 Tetrahedron The tetrahedron is one of the solid geometric forms used to represent the arrangement of atoms in clay mineral crystal structures. It is formed by connecting the centres of the four oxygen anions surrounding a central cation. In the clay minerals the predominant central cation of the tetrahedron is silicon. A limited number of tetrahedral are occupied by aluminium and occasionally ferric iron or other elements. A silicon, or aluminium, ion is surrounded by four oxygen ions to form a tetrahedron in the Figure (3a). The isolated tetrahedron has a net negative charge of -4 (Si with 4+ charges and four O with 2- charges). The tetrahedral rest on triangular face and the four triangular faces of the tetrahedron are formed by joining the centres of the anions. Only two of the faces are visible in the polyhedral illustration on the Figure (3a). In clay minerals, the three oxygens at the base of the tetrahedron are shared with adjacent tetrahedral and only the apical oxygen retains a charge of -1. Al may freely substitute for the silicon ions.

A supplemental view on the Figure 4 emphasizes that the cation (blue) occupies the centre of the tetrahedron (blue lines), and that it is bonded (brown rods) to four anions which would be located at the corners of the tetrahedron.

The tetrahedral sheet is formed by sharing each of the three oxygen atoms at the base of a tetrahedron with neighbouring tetrahedral. When you view the tetrahedral sheet from the side, the planar distribution of atoms and the sum of charges in each plane are readily apparent (Figure 5a). Each atomic plane in the sheet has a unique composition and charge. The composition of the sheet is: Si4O10 and it has a net charge of -4.

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2.1.2 Octahedron

The second structural unite is the octahedral sheet, in which the hydroxyl atoms (OH) in the corners and cations in the centre. The cations are usually aluminium (Al), iron (Fe), and magnesium (Mg) atoms. The octahedral sheet is comprised of closely packed oxygens and hydroxyls in which Al, Fe, and Mg atoms are arranged in octahedral coordination. The net charge on an isolated Al-OH octahedron is -3 (Al 3+ and six OH with 1- charge). In the octahedral sheet the charge is reduced through the sharing of anions by adjacent octahedral. A single octahedron can be recognized by following the bonds from the small blue balls (Al atoms). Three of them are directed upwards and are each connected a hydroxyl group as indicated by the arrows (Fig. 3b). The remaining three bonds are directed downwards to other hydroxyl groups. When aluminium with positive valence of three (Al +3) is present in the octahedral sheet, only two-thirds of the possible positions are filled in order to balance the charges. When only two octahedral sites filled with trivalent cations is a dioctahedral sheet. When magnesium with a positive charge of two (Mg+2) is present, all three positions are filled by divalent cations is a trioctahedral sheet.

(a) Tetrahedron (T)

(b) Octahedron (O)

Figure 3. Tetrahedron and octahedron geometric forms.

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Figure 4. View emphasizes that the cation (blue) occupies the centre of the tetrahedron

The octahedral sheet is formed by sharing all hydroxyl groups at the corners of an octahedron with neighbouring octahedral. When you view the octahedral sheet from the side (Figure 5b), they contain four aluminium atoms, six lower plane hydroxyls, and six upper plane hydroxyls. These rectangles are the same size as the planar motif outlined for the tetrahedral sheet. The formula for this unit is: Al 4 (OH) 12 and the net charge is ZERO. a

4 apical O -8 4 Si +16 6 basal O -12

6 OH -6 4 Al +12 6 OH -6

Figure 5. Side view showing the tetrahedral and octahedral sheets.

2.1.3 Layer types

These basic building blocks are linked in clay minerals to form sheets of silica tetrahedral and aluminium or magnesium octahedral. The silica tetrahedral sheet (T) and the octahedral sheet (O) are joined in two possible way: (1) The 1:1 layer silicate structure, 1(T) +1(O) sheet so that the apical oxygen of the tetrahedral sheet replaces one hydroxyl of the octahedral sheet to form what termed the 1:1 clay mineral layer as kaolinite (Fig. 6). The second way the 2:1 layer silicate

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structure, 1(O)+2(T) sheets so that 2/3 of hydroxyls in the octahedral sheet between 2 tetrahedral sheets are replaced by apical oxygens of the tetrahedral sheet to form 2:1 clay mineral layer (e.g. illite) as seen in (Fig. 7). Typical results for 1:1 and 2:1 dioctahedral minerals with no substitutions are, respectively: (none)(Al4)(Si4)O10(OH)8 and (none)(Al4)(Si8)O20(OH)4.

Clay minerals are not simply pure hydrous aluminium silicates. Mg may substitute for Al in the octahedral sheet and other ionic substitutions are possible. The major constraint is that the replacing ion must have a radius that will not require a different coordination, nor seriously disrupt the structure of the sheet. The replacing element may have a different ionic charge. Composi- tional variations may alter the physicochemical properties of the clay.

The most common isostructural substitutions are: • Al and ferric Fe for Si in the tetrahedral sheet, and • Interchange of Mg, Al, ferrous Fe, ferric Fe, and Mn in the octahedral sheet.

In the tetrahedral sheet, the common substitutions replace Si with an ion of lower valence. This is a point that you must remember when calculating the charge on the cation plane of the tetrahedral sheet. A tetrahedral sheet with Si3Al or Si3Fe (One of the 4 Si4+ atoms replaced by either Al3+ or Fe3+.) will have 15+ charges associated with the cation plane rather than 16+. In the octahedral sheet, similar reductions in charge occur when Al3+ is replaced by Mg2+ or Fe2+. However if Mg2+ is replaced by Al3+ or Fe3+, the result is an extra + charge. An octahedral sheet with (Al3Fe3+ 0.5Mg 0.5) would have a total charge associated with the cation plane of +11.5.

Figure 6. Structure of 1:1 layer silicate (kaolinite) illustrating the connection between tetrahedral and octahedral sheets.

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Figure 7. Structure of 1:2 layer silicate (illite) illustrating the connection between sheets.

2.2 Clay mineral Classification The important structural and chemical differences among the clay minerals are the basis for the individual mineral species names and the arrangement of the species in groups. The basis for classification of clay minerals is established in a series of articles by CMS and AIPEA Nomen- clature Committees. Don't read about nomenclature until you have completely digested the general article by Brindley and G. Brown (1984).

The important structural and chemical differences among the clay minerals are the basis for the individual mineral species names and the arrangement of the species in groups. Planar hydrous phyllosilicates, the common clay minerals, are classified according to the layer type, the magni- tude of the net layer charge, the type of interlayer material, the character of the octahedral sheet, and the composition or structure of individual species.

These groups are kaolinite, smectite, illite, chlorite and palygorskite.

A useful classification of the clay minerals (Table 2) was proposed and used by Grim in his book "Clay Mineralogy" (1968), which is a basis for outlining the nomenclature and differences between the various clay minerals. The phyllosilicates are divided into groups, each containing dioctahedral and trioctahedral subgroups. These subgroups are composed principally of either two or three sheets of atoms of two main kinds; one of silicon and oxygen atoms (silica layer) and the other a combination of aluminium with oxygen or hydroxyl atoms (the alumina or aluminium hydroxide layer). These layers are united chemically in either alumina-silica pairs (the kaolinite group), or in silica-alumina-silica trios (the montmorillonite and illite groups).

Clay minerals are part of the larger class of silicate minerals: the phyllosilicates. Included in the phyllosilicate family are the larger true micas, which include the familiar minerals muscovite and

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biotite and the brittle micas, which includes the less-familiar mineral margarite (a calcium rich member of the mica group of the phyllosilicates with formula: CaAl2(Al2Si2)O10(OH)2). We have learned much of what we know about clay minerals from the macroscopic (i.e., single crystal) study of the true micas. The true micas will be included in our discussion because they are well characterized and serve as a good model by which to understand clay structures (Fig 8).

Table 2. Classification of the clay minerals

Amorphous -Allophane group Crystalline A. Two layer type (one silica-tetrahedrons and one alumina-octahedrons) 1. Eqidimentional Kaolinite group Kaolinite, nacrite, dickite 2. Elongate Halloysite B. Three layer type (two layers of silica and one central layer of alumina) 1. Expanding lattice a. Eqidimentional Smectite group Na-montmorillonite, Ca-montmorillonite and beidellite Vermiculite b. Elongated Smectite Nontronite, saponite, hectorite 2. Non-expanding lattice Illite group Chlorite group Dioctahedral chlorite (donbassite) Di, trioctahedral chlorite (Cookeite, sudoite) Trioctahedral chlorite (Clinochlore, chamosite, nimite)

C. Mixed-layer types (ordered stacking of alternate layers of different types)

D. Chain structure type (hornblende-like chains of silica tetrahedral linked together by octahedral groups of oxygens and hydroxyls containing Al and Mg atoms) Palygorskite (attapulgite), sepiolite,

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Figure 8. Classification of silicates and clay minerals (Bailey, 1980b; Rieder et al., 1998).

In this chapter, a general review of the structure and composition of the various clay minerals are given. Those who are interested in more detailed discussions of the structures as Bailey (1980, 1988), Moore and Reynolds (1997) and Murray (2002, 2007). The physical and chemical properties of a particular clay mineral are dependent on its structure and composition.

2.2.1 Allophane Allophane is a series name used to describe clay-sized, short-range ordered aluminosilicates associated with the weathering of volcanic ashes and glasses. Allophane commonly occurs as very small rings or spheres having diameters of approximately 35 - 50 Å. This morphology is characteristic of allophane, and can be used in its identification.

Allophanes have a composition of approximately Al2Si2O5·nH2O. Some degree of variability in the Si: Al ratios is present: Wada K., (1989) reports Si: Al ratios varying from about 1:1 to 2:1. Because of the exceedingly small particle size of allophane and the intimate contact between allophane and other clays (such as smectites, imogolite, or non-crystalline Fe and Al hydroxides and silica) in the soil, it has proven very difficult to accurately determine its composition. Con- sequently, there is always some potential error associated with the compositional ratios has reported.

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The term allophane has been applied to numerous other materials in the past, including imogolite, any non-crystalline aluminosilicate, or any clay-sized material exhibiting structural randomness. Current usage is limited to short-range ordered aluminosilicates having Si: Al ratios between 1:2 to 1:1 possessing a spherical or ring-shaped morphology. Allophane usually gives weak XRD peaks at 2.25 and 3.3 Å. Identification is commonly made by infrared analyses or based on transmission electron morphology.

A limited amount of isomorphism substitution occurs in allophane. The most common type is the substitution of Fe for Al. Little permanent charge is assumed to be present. The majority of the charge is variable charge, and both cation and anion exchange capacities exist, with the relative amounts depending on the pH and ionic strength of the soil chemical environment. Wada reports CEC values of 10 - 40 cmol kg-1 at pH 7.0 and AEC values of 5 - 30 cmol kg-1 at pH 4.0. Other studies have measured CEC values as low as 10 and as high as 135 cmol kg-1 at pH 7.0. The surface area of allophane has been calculated to be about 1000 m2 g-1, while values measured with ethylene glycol monoethyl ether are in the range of 700 - 900 m2 g-1.

2.2.2 Kaolin Minerals

The kaolin group minerals comprise kaolinite, nacrite, dickite and halloysite, and are among the most common clay minerals in nature. They have a 1:1 layered structure, that is, each layer con- sists of one tetrahedral silicate sheet and one octahedral sheet, with two-thirds of the octahedral sites occupied by aluminium. Kaolinite, nacrite and dickite all have the ideal chemical composition: Al2Si205(OH)4, they differ from one another only in the manner in which the 1:1 layers are stacked. Halloysite, in its fully hydrated form, has the ideal chemical formula Al2Si205(OH)4.2H20 and the theoretical chemical composition is SiO2, 46.54%; Al2O3, 39.50%; and H2O, 13.96%. Kaolinite differs from the other three members of the group by including molecular water in the interlayer.

Within the kaolin group minerals, kaolinite is the most abundant and has received most attention in terms of its structure, properties and industrial applications. However, because of its close similarity with the aforementioned polytypic, many of the properties and uses described for kaolinite apply equally to the other polytypic. Consequently, for the purposes of expediency, the following disclosure will be restricted primarily to kaolinite and halloysite but it should be borne in mind, as it will be readily appreciated by those skilled in the art, that the invention applies equally to nacrite and dickite. Kaolinite is the most important member of the group. It occurs in residually weathered material and is a common constituent of soil .Naturally occurring kaolins typically have a wide range of particle sizes, particle crystallinity, minor element compositions and chemical reactivity for intercalation reactions. Kaolins sorted into a size range of 0.5 - 2.0 mm typically have a specific surface of about 5 m2 g-1 and a cation exchange capacity of 10 meq./100 gm or less. These and other properties, such as opacity and viscosity, make kaolins suitable for a wide range of uses including paper coatings and fillers, pottery, porcelain and sanitary ware production and fillers in paints and rubbers. These properties however do not allow kaolins to be readily utilised in other uses as described hereinafter. However, if their specific surface and/or cation exchange capacities could be increased, their usefulness would be increased and thus they could then be used in many other applications including use as catalysts, metal scavengers, carriers and absorbents. For more explanation see chapters three and six.

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Dickite and nacrite result from hydrothermal or pneumatolytic alteration (pneumatolytic is the process by which ores and minerals are formed from the action of vapours produced by igneous magmas). They are rare in clay materials and virtually of no importance commercially. They differ from kaolinite in the way the silicate units are stacked.

Halloysite is similar to kaolinite but the microscopic crystalline particles are elongate rather than equidimensional in shape. Halloysite occurs in two forms: one hydrated, in which there is a layer of water molecules between the layer, and one dehydrated. The hydrated form has a basal spacing of 10 Å and dehydrated form, 7.2 Å (Fig. 9). The shape of halloysite is elongated tubes, whereas the shape of kaolinite is pseudo-hexagonal plates and stacks (Fig. 10). The International Nomenclature Committee has recommended the terms 7 Å halloysite and 10 Å halloysite to des- ignate the two forms. The elongated tubular form according to Papoulis et al. (2004) is made up of overlapping curved sheets of kaolinite type. Halloysite has a similar structure to kaolinite but contains a single layer of water (2.9Å) in the interlayer space. The layer thickness is therefore, 10Å. There is also lots of disorder between layers. The sheets of water are trapped during crystallization. The presence of water between the layers alters the distribution of stresses within the mineral lattice such that the layers curve to form a tubular structure (Joussein, et al. 2005). That is mean the halloysite occurs as cylinders or spherical shapes (due to hydrogen bonding with water molecules, see Fig.11). So10Å halloysite is stable under the ground water level conditions that include water saturation and when the Al: Si ratio is approximately 1:1 (Churchman and Carr, 1975).

Figure 9. XRD patterns of oriented powder mount of the <2μm fraction two samples from Kahdeksai- siensuodeposit, showing the appearance of hkl peaks kaolinite and halloysite without mica/illite (after Thair Al-Ani, et al. 2004).

The halloysite content in some parts of Virtasalami deposits may indicate both hydrothermal alteration and in situ weathering of the plagioclase feldspars in quartz-feldspar gneiss parent rocks probably formed the Virtasalmi kaolin. Such processes have been described by Schwaighofer

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and Muller (1987). Halloysite develops in a changing geochemical environment caused by a constant alteration of rainy and dry seasons. Based on the occurrence of the elongate or halloysite particles as observed by SEM, it is apparent that at least most the elongated halloysite were formed later than kaolinite and then it dehydrated into metahalloysite. Some studies indicate that the 10-Å form is unstable, always transforms into 7-Å halloysite with the time.

Figure 10. SEM from platelets of Virtasalmi (Litmanen) kaolinite and halloysite consisting dominantly of long tubes(after Thair Al-Ani, et al. 2004).

Figure 11. Development of halloysite from kaolinite by weathering kaolinite sheets (Robertson and Eggleton, 1991).

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2.2.3 Smectite Minerals

The major smectite minerals are Na-montmorillonite, Ca-montmorillonite, saponite (Mg- montmorillonite), nontronite (Fe-montmorillonite), hectorite (Li-montmorillonite), and beidellite (Al-montmorillonite). Smectite minerals are composed of two silica tetrahedral sheet with central octahedral sheet and are designated as a 2:1 layer mineral (Fig. 12). Water molecules and cations occupy the space between the 2:1 layers. The theoretical charge distribution in the Smec- tite layer without considering substitutions in the structure is as shown in Table (3).

Table 3. Charge distribution of the Smectite layer.

6O2- 12- 4Si4+ 16+ 4O2- + 2(OH) - 10-(Plane common to tetrahedral and octahedral sheets) 4Al3+ 12+ 4O2- +2(OH) - 10-(Plane common to tetrahedral and octahedral sheets) 4Si4+ 16+ 6O2- 12-

The theoretical formula is (OH)4Si8Al4O20.NH2O (interlayer) and the theoretical composition without the interlayer material is SiO2 , 66.7%; Al2O3, 28.3%; and H2O, 5%. However, in smectite, there is considerable substitution in the octahedral sheet and some in the tetrahedral sheet. In the tetrahedral sheet, there is substitution of aluminium for silicon up to 15% and in the octahedral sheet, magnesium and iron for aluminium (Grim, 1968). If the octahedral positions are mainly filled by Al, the smectite mineral is beidellite; if filled by Mg, the mineral is saponite; and if by Fe, the mineral is Nontronite.

Figure 12. Diagrammatic sketch of the structure of smectites (USGS).

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The most common smectite mineral is Ca-montmorillonite, which means that the layer charge deficiency is balanced by the interlayer cation calcium and water. The basal spacing of the Ca- montmorillonite is 14:2Å. Na-montmorillonite occurs when the charge deficiency is balanced by sodium ions and water and basal spacing is 12.2Å. Ca-montmorillonite has two water layers in the interlayer position and Na-montmorillonite has one water layer.

It is well established that beidellites are more common in soil environments, whereas montmorillonite are more typical of geologic materials (Wilson 1987). Moreover, high-charge beidellites were reported as the dominant clay mineral in the highly smectitic soils known as Vertisols (Badraoui et al.1987; Badraoui and Bloom 1990). Soil beidellites are thought to exist as the weathering product of micas and chlorites, because these already have the tetrahedral substitution required for the beidellite structure. The evidence for the formation of smectites from soil solution is difficult to establish. According to Borchardt (1989), beidellites would be crystallized from soil solutions if the pH is less than 6.7, in which case exchangeable A1 is present, whereas montmorillonite would be formed in base-saturated soils with low organic matter content and high pH (>6.7), which is the case in Vertisols. However, Kounetsron et al. (1977) showed that beidellite was formed in a Vertisols from the weathering of mica through dissolution and recrys- tallization processes rather than by simple transformation of the mica structure.

The smectite mineral particles are very small and because of this, the X-ray diffraction data are sometimes difficult to analyze. A typical electron micrograph of sodium montmorillonite is shown in Fig. 13. Smectites, and particularly Na-montmorillonite, have a high Base Exchange capacity as is described later in this book.

Figure 13. Authigenic smectite (montmorillonite) overgrow on pore spaces and authigenicly-overgrown quartz grains in sandstone.

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2.2.4 Vermiculite

Vermiculite is a member of clay minerals, produced by the decompositions of micas and occurs as quite large crystals of mica-like appearance. It has a layer structure, and the interlayer contains water molecules and exchangeable cations, mainly Mg2+ ions.

The structure of Mg-saturated vermiculite resembles talc in that it contains a central octahe- drally-coordinated layer of Mg ions that lie between two inwardly pointing sheets of linked tetrahedral. These silicate layers are normally separated by two sheets of interlayer water molecules arranged in a distorted hexagonal fashion (Fig. 14). The chemical composition of vermiculite in weight percentage was: SiO2 (44.62); Al2O3 (9.18) Fe2O3 (5.46); CaO (0.78); MgO (20.44); Na2O (0.11); K2O (0.48) with loss of weight after heating at 1273 K or at 1000°C (18.93). Based on the data, the typical structural formulae of the vermiculite is (MgFe,Al)3(Al,Si)4O10(OH)2·4H2O .

Because vermiculite and smectite have smaller layer charge than micas the attractive forces between the 2:1 layers and the interlayer cation is less. The hydration energy of the interlayer cation may then be sufficient to overcome the attractive forces of the layer to the cations and allow water to hydrate the interlayer cation which causes swelling normal to the plane of the layers (Brindley & Brown, 1984). The ability of vermiculite and smectite to swell in water allows cation exchange between the interlayer cation and cations in an external solution. Both groups of minerals can also sorbs organic cations by cation exchange and other organic molecules by sal- vation of the interlayer cations. In fact a widely used diagnostic test for identifying vermiculite and smectite is based on the amount of swelling when ethylene glycol or glycerol is sorbed between the 2:1 layers. This statement was confirmed by XRD patterns (Fig. 15) that shows swelling-lattice vermiculite is very common and in X-ray analyses it has typically a strong 16 Å peak after ethylene glycol treatment. Vermiculite is termed as swelling when the basal spacing at 14 Å shifts to below 16.70 Å after ethylene glycol treatment while smectites shifts to 16.70-17.1 Å (Grim 1968, Thorez 1975, Brindley & Brown 1984). KCl-treatment shifts vermiculite between 12.40-12.80 Å to 10 Å (Thorez,1975). After KCl- and heat treatment 14 Å peak shifts to 10 Å, this is a typical reaction for vermiculite. Generally, vermiculite swells less than smectite because the interlayer cation to 2:1 layer attractive forces is greater. The CEC of vermiculite was 135 meq/100 g and BET surface area was 16 m2 g− 1.

Vermiculite is identified on the basis of its strong 14.60-14.00 Å peak, which shifts to 10 Å after heating to 450 - 550 ºC. Progressive removal of this interlayer water results in a series of less hydrated phases that include the 14.36 Å lattice with two sheets of water molecules, a 11.59 Å lattice with a single sheet of water molecules, and a 9.02 angstrom lattice from which all water has been removed . This statement was confirmed by XRD patterns (Fig. 16) that made possible to determine changes in interlayer spacing that served as the radon diffusion channels: for the ungrounded sample before the heat treatment d002 was 14.4Å, it decreased to 11.6Å after sample heating at 100°C, and to about 10Å for the sample heated at 300°C. In the temperatures above 800°C caused a decrease in the interlayer spacing due to the collapse of layers, yielding a talc-like structure, which is characterized by the stack layers without water or cations in the interlayer space Inasmuch as the layers are electrically neutral and interlayer cations occupy only about one-third of the available sites, cohesion between the layers is typically weak (V. Balek. et al., 2007). It was detected by high temperature XRD that on further heating to 900°C and above this temperature new crystalline phases, i.e. enstatite and spinel, were formed.

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Figure 14. Diagrammatic sketch of the structure of vermiculite (USGS).

Vermiculites are usually formed in sediments by the alteration of micaceous minerals (biotite and chlorite to trioctahedral vermiculite; muscovite to dioctahedral vermiculite (Moore and Rey- nolds, 1997). However, vermiculites formed through the alteration mica are comparatively rare in marine sediments because the K of sea water readily contracts them (Deer et al., 1975). While present, marine vermiculites are probably derived from volcanic material, chlorite, and hornblende.

In the studied vermiculite in Finnish tills, Pulkkinen (2004) point out that swelling-lattice vermiculite is a common clay mineral in the clay fraction of Finnish tills. The studying by transmission electron micrographs point out that vermiculite is fine grained about ~1 μm (Fig.17).

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Figure 15. XRD patterns of the K-saturated material noted ‘A’: (a) heating overnight at 110°C and solvated with ethylene glycol vapour at room temperature (K-110°C-EG vapor-rt); (b) solvated with ethylene glycol vapour at room temperature (K-EG vapor-rt); (c) solvated with ethylene glycol vapour at 65°C (K-EG vapor-65°C); (d) solvated with liquid ethylene glycol; (e) air-dried (K-sat-AD). (After Régine Mosser et al., 2005)

Figure 16. XRD patterns of vermiculite samples heated at different temperatures a – un-ground sample, b –ground sample for 2 min. E – enstatite, S – spinel (after V. Balek. et al., 2007)

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Figure 17. TEM-photograph of Chlorite/vermiculite in clay fraction of Suomussalmi, Isopalo and Kianta moraine. Identification is based on chemical elements (EDS) (after Pekka Pulkkinen, 2004).

2.2.5 Illite minerals

Illite is a clay mineral mica, which was named by Grim et al. (1937). The structure is a 2:1 layer in which the interlayer cation is potassium (Fig.18). The size, charge, and coordination number of K is such that it fits snugly in hexagonal ring of oxygens of the adjacent silica tetrahedral sheets. This gives the structure a strong interlocking ionic bond which holds the individual layers together and prevents water molecules from occupying the interlayer position as it does in the smectite. Simply it might say that illite is a potassium smectite.

Illite differs from well-crystallized muscovite in that there is less substitution of Al3+ for Si4+ in the tetrahedral sheet. In muscovite, one-fourth of Si4+ is replaced by Al3+ whereas in illite only one-sixth is replaced. Also, in the octahedral sheet, there may be some replacements of Al3+ by Mg2+ and Fe2+. The basal spacing d(001) of illite is 10Å. A more detailed discussion of the structure of illite and its variable composition can be found in Moore and Reynold (1997). The charge deficiency, because of substitution per unite cell layer, is about 1.3-1.5 for illite contrasted to 0.65 for smectite. The largest charge deficiency is in the tetrahedral sheet rather than in the octahedral sheet, which is opposite from smectite. For this reason and because of the fit, potassium bonds the layer in a fixed position so that water and other polar compound cannot readily enter the interlayer position and also the K ion is not readily exchangeable. Illites, which are the dominant clay minerals in argillaceous rocks, form by the weathering of silicates (primarily feldspar), through the alteration of other clay minerals, and during the degradation of muscovite (Deer et al., 1975). Formation of illite is generally favoured by alkaline conditions and by high concentrations of Al and K.

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Figure 18. Diagrammatic sketch of the structure of illite (USGS).

Figure 19. SEM of illite from Lappajärvi clays(after Al-Ani & Vaarma, 2005).

Most studies of fine clay fraction of till in Finland indicated that the most abundant clay minerals are illite, vermiculite, kaolinite and swelling-lattice vermiculite. The most important factors controlling the mineralogical and geochemical composition of the fine and clay fractions of the tills in Finland are the composition of the bedrock and the possible occurrence of an old weathering crust (Pulkkinen, 2004). Al-ani and Vaarma (2005) were studied Lappajärvi Impact Crater in

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western Finland and found that illite is the most common clay mineral associated with kaolins and smectite as shown in the SEM and X-ray diffraction of fines in Figure(19 and 20).

Figure 20. XRD pattern of illite and other clay minerals in Lappajärvi quaternary deposits (after Al-Ani & Vaarma, 2005).

2.2.6 Chlorite Minerals

The chlorite group members contain a 2:1 layer with variable x and an interlayer hydroxide sheet. In some references, they are referred to as 2:1:1 mineral. The octahedral sheets may both be dioctahedral (di/di) or trioctahedral (tri/tri), or mixed (di/tri, or tri/di). The interlayer hydroxide sheet may have a positive charge.

This projection of a chlorite structure reveals the basic sequence of 2:1 layers separated by a complete interlayer hydroxide sheet (Fig 21). All of the octahedral in the hydroxide sheets are filled, making both sheets trioctahedral, but they have different compositions. All of the 2:1 octahedral positions are occupied by Mg ( dark yellow octahedral). In the interlayer, Mg and Al ( Yellow octahedral) are both present.

This chlorite also has one more Al in the tetrahedral sheet than the common micas (Fig 22). Each tetehedral sheet cation plane contains Al2Si2. The octahedral sheet in the 2:1 layer is Mg- filled, but the interlayer hydroxide sheet has Al substituting for Mg. The interlayer sheet thus has a net positive charge.

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Figure 21. Diagrammatic sketch of the structure of chlorite (USGS)...

Figure 22. Chlorite structure (USGS).

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Dioctahedral chlorite is dioctahedral in both the 2: 1 layer and the interlayer hydroxide sheet. An example is donbassite (Brindley and Brown, 1984). Trioctahedral chlorites should be named according to the dominant divalent octahedral cation present. Recommended species names are 2+ clinochlore for Mg-dominant [end member = (Mg5Al)(Si3Al)O10(OH)8], chamosite for Fe - 2+ dominant [end member = (Fe5 Al)(Si3Al)O10(OH)8], nimite for Ni-dominant [end member = 2+ (Ni5Al)(Si3Al)O10(OH)8], and pennantite for Mn -dominant [end member = 2+ (Mn5 Al)(Si3Al)O10(OH)8]. All other species and varietal names should be discarded because arbitrary subdivisions according to octahedral and tetrahedral compositions have been shown to have little or no structural significance. Tetrahedral compositions and trivalent octahedral cations are not considered in the recommended species names, nor is the distribution of octahedral cations between the 2: 1 layer and the interlayer. Adjectival modifiers, such as those of Schaller (1930), may be used to indicate either important octahedral cations other than the dominant cation or unusual tetrahedral compositions. Bayliss (1975) gives modifiers appropriate for many of the chlorite species listed in other nomenclature systems. Figure (23) shows the SEM image of chlorite with quartz grain.

Figure 23. SEM image of chlorite (after Bayliss, 1975)

2.2.7 Palygorskite and sepiolite Palygorskite and sepiolite have similar fibrous or lath-like morphologies, but palygorskite exhibits more structural diversity and, although both minerals are Mg silicates, palygorskite has less Mg and more Al than sepiolite (Moore and Reynolds, 1997). Structurally, (palygorskite and sepiolite) consist of blocks and channels "ribbon-like" sheets extending in the c-axis direction. Each structural unit is built up of two tetrahedral silicate layers and a central trioctahedral layer. In the octahedral layer Mg+2 ions occupy two different structural positions: (1) on the borders of the structural blocks, coordinated to water molecules; and (2) in the interior of the blocks, linked to hydroxyl groups.

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The structures are modulated from the ideal by the periodic reversal of tetrahedral apices along Y. Channels between ribbons may contain exchangeable cations and water. Differences in the structures are due to the width of the ribbons.

In a (100) projection of palygorskite (Fig 24a), ribbon-like strips of the 2:1 layer are apparent due to inversion of tetrahedral (arrows). The ribbons or strips are five octahedral cations wide. A similar projection of the sepiolite structure (Fig 24b) reveals ribbons that are eight octahedral cations wide (Bailey, 1980).

(a)

(b)

Figure 24. Diagrammatic sketch of the structure of palygorskite (a) and sepiolite (b) (USGS).

These minerals have structural attributes that are similar to those exhibited by pyriboles and micas. Modulation produces a structure with ribbons that are wider than in the pyriboles but are not continuous enough to be micas. Both sepiolite and palygorskite clay minerals are Mg-silicates, but palygorskite has higher alumina content. A general formula for palygorskite is (OH2)4(OH2)Mg5Si8O20.4H2O. A general formula for sepiolite is (OH2)4(OH)4Mg8Si12O30.8H2O.

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These two clay minerals contain two kinds of water, one coordinated to the octahedral cations and other loosely bonded in the channels, which is termed zeolitic water. These channels may also contain exchangeable cations. Unique chemical properties are a result of the "zeolite-like" channels associated with these modulated materials. (a) (b)

Figure 25. Transmission electron micrograph of (a) sepiolite and (b) palygorskite

Figure (25) shows an electron micrograph of palygorskite and sepiolite. Both palygorskite and sepiolite are elongated in shape and often occur as bundles of elongated and lath-like particles. Usually, the sepiolite elongated are longer than Palygorskite elongates (10-15Å for sepiolite and >5Å for Palygorskite). The morphology of these two clay minerals is a most important physical attribute crystal (Galan, E. and Ferrero, A., 1982)

The fibrous nature of sepiolite and palygorskite precludes the production of oriented aggregate mounts to enhance the 001 reflection for X-ray powder diffraction (Fig. 26). However, strong reflections from the 011 planes yield intense peaks at 12.2 angstroms in sepiolite and at 10.5 angstroms in palygorskite. These peaks are unaffected by solvation with ethylene glycol, but change during heat treatments. After heating to 400 C, the 001 peak of both minerals is reduced in intensity and new palygorskite peaks occur at 9.2 and 4.7 angstroms (Singer, 1989). After heating to 550 C, the original 011 palygorskite and sepiolite peaks are completely destroyed, but now new peaks for sepiolite occur at 10.4 Å.

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Figure 26. XRD difractograms of (a) powder XRD pattern of the bulk sample; (b) air-dry, EG-solvent, and heated 550 ºC, for palygorskite and sepiolite from Nile Valley in Egypt (Carlos, et al.,1998).

2.2.8 Mixed-layer clay minerals

Mixed-layered or interstratified clay minerals refer to remarkable phyllosilicates structures, characterized by a vertical stacking sequence of two or more types of single layers. The layers involved can be of 2:1, 2:1:1 and even 1:1 types. Mixed layer clays are common, and consist of clays that change from one type to another through a stacking sequence. They mainly form through weathering or middle-late diagenesis, but also characterize some hydrothermal and sedimentary environments. The sequences can be ordered and regular or high unordered and irregular. The mixed-layer minerals show in X-ray analyses characteristics that are intermediate between those of the individual minerals involved. Typical peaks for mixed-layer clay minerals between 11-14 Å (Chamley 1989). The most common mixed-layer clay minerals are mixtures of vermiculite, chlorite, illite and swelling-lattice vermiculite. Mixed-layered clay minerals were identified mostly by X-ray diffraction analysis of glycolated and orientated sample fractions at less than 2 um and less than 0.2 um. Numerous researches obtained the infrared absorption spectra and differential thermal analysis to study mixed-layered clay minerals such as Keonu and Avasur (1965).

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CHAPTER THREE

3 PHYSICAL AND GEOCHEMICAL PROPERTIES OF CLAYS AND CLAY MINERALS

3.1 The main properties of particular clay minerals

The physical and chemical properties of particular clay and clay minerals are dependent on the structure and composition. The structure and composition of the major industrial clays, i.e. kaolins, smectites, illite and palygorskite-sepiolite, are very different even though each is comprised of octahedral and tetrahedral sheets as their basic building lock. However, the arrangement and composition of the octahedral tetrahedral sheets account for most differences in their physical and chemical properties.

The important physical and chemical characteristics that relate to the applications of the clay material are shown in Table 4. In most applications, the clays are used because of the particular physical properties that contributed to the end product, i.e. kaolin for paper coating or bentonite in drilling muds. In some cases, the clay is used for its chemical composition, i.e. kaolin for use as a raw material to make fibreglass or clays and shales in the mix to make cement.

Table 4. Important physical and chemical characteristics of clay materials

Particle size, shape and distribution Mineralogy Surface area, charge, and chemistry PH Ion exchange capacity Brightness and colour Sorption capacity Reheology Ceramic properties Dispensability

The physical and chemical characteristics of kaolinite, smectite (bentonite), illite, palygorskite and other clay minerals are discussed in the following sections.

3.1.1 Kaolin

Kaolinite, the main constituent of kaolin, is formed by rock weathering. It is white, greyish- white, or slightly coloured. It is made up of tiny, thin, pseudohexagonal, flexible sheets of tri- clinic crystal with a diameter of 0.2–12 µm. It has a density of 2.1–2.6 g/cm3. The cation exchange capacity of kaolinite is considerably less than that of montmorillonite, in the order of 2– 10 meq/100 g, depending on the particle size, but the rate of the exchange reaction is rapid, almost instantaneous (Grim, 1968). Kaolinite adsorbs small molecular substances such as lecithin, quinoline, paraquat, and diquat, but also proteins, polyacrylonitrile, bacteria, and viruses

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(McLaren et al., 1958; Mortensen, 1961; Weber et al., 1965; Lipson & Stotzky, 1983). The adsorbed material can be easily removed from the particles because adsorption is limited to the surface of the particles (planes, edges), unlike the case with montmorillonite, where the adsorbed molecules are also bound between the layers (Weber et al., 1965). Upon heating, kaolinite starts to lose water at approximately 400 °C, and the dehydration approaches completeness at approximately 525 °C (Grim, 1968). The dehydration depends on the particle size and crystallinity.

3.1.2 Smectite (Bentonite)

Smectite feels greasy and soap-like to the touch. Freshly exposed bentonite is white to pale green or blue and, with exposure, darkens in time to yellow, red, or brown. The special properties of smectite are an ability to form thixotrophic gels with water, an ability to absorb large quantities of water with an accompanying increase in volume of as much as 12–15 times its dry bulk, and a high cation exchange capacity. Substitutions of silicon by cations produce an excess of negative charges in the lattice, which is balanced by cations (Na+, K+, Mg2+, Ca2+) in the interlayer space. These cations are exchangeable due to their loose binding and, together with broken bonds (approximately 20% of exchange capacity); give montmorillonite a rather high (about 100 meq/100 g) cation exchange capacity, which is little affected by particle size. This cation exchange capacity allows the mineral to bind not only inorganic cations such as caesium but also organic cations such as the herbicides diquat, paraquat and s-triazines (Weber, 1970), and even bio-organic particles such as rheoviruses and proteins (Lipson & Stotzky, 1983), which appear to act as cations. Variation in exchangeable cations affects the maximum amount of water uptake and swelling. These are greatest with sodium and least with potassium and magnesium. Interstitial water held in the clay mineral lattice is an additional major factor controlling the plastic, bonding, compaction, suspension, and other properties of montmorillonite-group clay minerals. Within each crystal, the water layer appears to be an integral number of molecules in thickness. Physical characteristics of bentonite are affected by whether the montmorillonite composing it has water layers of uniform thickness or whether it is a mixture of hydrates with water layers of more than one thickness. Loss of absorbed water from between the silicate sheets takes place at relatively low temperatures (100–200 °C). Loss of structural water (i.e., the hydroxyls) begins at 450–500 °C and is complete at 600–750 °C. Further heating to 800–900 °C disinte- grates the crystal lattice and produces a variety of phases, such as mullite, cristobalite, and cordi- erite, depending on initial composition and structure. The ability of montmorillonite to rapidly take up water and expand is lost after heating to a critical temperature, which ranges from 105 to 390 °C, depending on the composition of the exchangeable cations. The ability to take up water affects the utilization and commercial value of bentonite (Grim, 1968). Montmorillonite clay minerals occur as minute particles, which, under electron microscopy, appear as aggregates of irregular or hexagonal flakes or, less commonly, of thin laths (Grim, 1968). Differences in substitution affect and in some cases control morphology.

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3.1.3 Illite

Illite, together with chlorite, is the main component of common clay and shale. It is also an important impurity in limestone, which can affect the properties and thus the value of the stone for construction and other purposes (Carr et al., 1994). Despite the widespread occurrence of illite in nature, large deposits of high purity are quite rare. Illite usually occurs as very small (0.1–2 µm), poorly defined flakes commonly grouped into irregular aggregates. Lath-shaped and ribbon-shaped illite particles up to 30 µm in length and 0.1– 0.3 µm in width have also been described (Srodon & Eberl, 1984), but their existence is contro- versial. Velde (1985) stated unqualifiedly that these so-called filamentous illites are mixed-layer structures. Srodon & Eberl (1984), however, drawing on the same references plus their own data, concluded that these filaments in some cases are mixed-layer structures but in other cases are composed only of illite, and they further supported their view with scanning electron microscopic photographs of lath-shaped crystals of what they identified as illite. The special properties of illite are derived from its molecular structure. The balancing cation is mainly or entirely potassium, and charge deficiency from substitutions is at least twice that of smectites (i.e., 1.3–1.5 per unit cell layer) and is mainly in the silica sheet and close to the surface of the unit layer rather than in the octahedral layer as in smectites (Grim, 1968). These differences from smectites produce a structure in which interlayer balancing cations are not easily exchanged and the unit layers are relatively fixed in position and do not permit polar ions such as water to readily enter between them and produce expansion. Illite reacts with both inorganic and organic ions and has a cation exchange capacity of 10–40 meq/100 g, a value intermediate between those of montmorillonite and kaolinite (Grim, 1968). Ion exchange capacity is reduced by heating. The potassium in the interlayer space is "fixed" to a considerable degree, making it not readily available to plants, a matter of importance in soil science and agriculture. A portion of the interlayer potassium, however, can be slowly leached, leading to the degradation of the illite. Such degradation, however, can be reversed by the addition of potassium. Wilken & Wirth (1986) stated that Fithian illite from Illinois, USA, and adsorbed hexachlorobenzene suspended in distilled water with a sorption partition coefficient of 2200–2600 and that more than half of this adsorbed hexachlorobenzene could be desorbed by further contact with distilled water. However, the fithian illite used in the experiment had a composition of 30% quartz, 19% feldspar, 11% kaolinite, 1% organic carbon, and 40% illite, making it impossible to know how much of the measured adsorption could be ascribed to illite. The dehydration and other changes in illite with heating have been studied by several investiga- tors, with inconsistent results (Grim, 1968). Some of the inconsistency in findings may result from differences in the period at which samples were held at a given temperature, since dehydration is a function of both time and temperature (Roy, 1949). It is also probable that small differences in particle size, crystal structure, and molecular composition among samples of what were ostensibly the same mineral contributed to the inconsistencies. Dehydration takes place either smoothly or in steps between about 100 and 800 or 850 °C for both biotite and muscovite illites. Loss of structure by the various illite minerals occurs between about 850 and 1000 °C.

3.1.4 Other clays Another important property of clay minerals, the ability to exchange ions, related to the charged surface of clay minerals. Ions can be attracted to the surface of a clay particle or take up within

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the structure of these minerals. The property of clay minerals that causes ions in solution to be fixed on clay surfaces or within internal sites applies to all types of ions, including organic molecules like pesticides. Clays can be an important vehicle for transporting and widely dispersing contaminants from one area to another.

3.2 The Geochemistry of clay minerals

3.2.1 Equilibrium adsorption and Ion exchange

The typically small grain size (<2 µm) of clay minerals results in the presence of large surface areas. These surface areas are available for exchange of ions and molecules between the solids and surrounding solutions. Exchange of ions involves adsorption and desorption which are commonly fast (on geological time scales). Adsorption takes place because of the attraction of ions to a surface. The strength of the bonding varies from weak van der Waals (physical adsorption) to moderate absorption (electrostatic adsorption) to strong chemical bonds (chemisorp- tions), henceforth simply referred to as adsorption. This process involves neutral species (H2O, H4SiO4, organic molecules) and ions.

Example of kaolinite: Notice in the schematic diagram below that for 1:1 structures, positive ions are attracted to the light-blue tetrahedral basal oxygen surface. At the same time, negative ions are attracted to the dark-blue octahedral hydroxyl surface.

Example of vermiculite or smectite. The case for low-charge 2:1 structures is notably different from 1:1 structures. The schematic diagram below shows that 2:1 structures have mostly positive ions are attracted to the light-blue tetrahedral basal oxygen surfaces.

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Clay minerals have high sorption capabilities. They may absorb large quantities of compounds in the interstices between the particles much like a paper towel absorbs water. Clays also adsorb ions due to the electrostatic charges generated by atomic substitutions in their crystal structure and other processes. Adsorbed ions may be firmly attached to clay surfaces or readily hydrated and exchanged. In some instances, ions may be fixed in structural positions within the silicate layers. Organic matter and oxides/hydroxides also contribute significantly to the sorption capacity of clay-rich materials.

Cation exchange reactions often dominate the adsorption reactions. They are mostly dependent on the permanent negative charge of the 2:1 type layers. Johnston (2002) has listed six of the active sites influencing the sorption of organic molecules on clays. 1) The neutral siloxane surfaces of the 2:1 clays where no isomorphism substitution has occurred function as a Lewis base site. 2) Permanently charged sites (usually negative) are localized where isomorphism substitutions occur in the tetrahedral or octahedral sheets. 3) Exchangeable metal cations or metals in isomorphism substitution sites may interact directly with adsorbed molecules. 4) Polarized water molecules coordinated to exchangeable cations or under coordinated metal ions at flake edges serve as sites of surface acidity. 5) Organic molecules adsorbed on clay surfaces may create hydrophobic surfaces or serve as molecular pillars for exchange molecules. 6) Broken edge sites associated with under coordinated metal ions form surface hydroxyl groups which are among the most active of sites.

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3.2.2 Surface charge properties

They are responsible for the pH dependent charge in soils and sediments. At low pH they develop a positive charge due to the adsorption of protons. At higher pH they may be neutral sites and eventually develop a negative charge. This pH-dependent charge (illustrated below) may be a significant factor in clays with small particle sizes and limited isomorphism substitution. One way in which surface charge can develop is by adsorption of an ion where the solid acts as an electrode. (e.g., H+ and OH- on the surfaces of clays). In clay-aqueous systems the potential of the surface is determined by the activity of ions (e.g., H+ or pH) which react with the mineral surface. The simultaneous adsorption of protons and hydroxyls as well as other potential determining cations and anions, leads to the concept of zero point of charge or ZPC, where the total charge from the cations and anions at the surface is equal to zero. The charge must be zero and this does not necessarily mean the number of cations versus anions in the solution is equal. For clay minerals the potential determining ions are H+ and OH- and complex ions formed by bonding with H+ and OH-.

For example on the basal oxygen surface of kaolinite or illite, the O2- ions bond with H+ to form hydroxyls. These surfaces further react as either acids or bases with other protons or hydroxyls: The surface charge is therefore, pH dependent. The broken Si-O bonds and Al-OH bonds along the surfaces of the clay crystals result in hydrolysis.

Low pH

+ + MOH + H --> MOH2

High pH

- - MOH + OH --> MO + H2O

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The pH that corresponds to the ZPC is referred to as the pHZPC or the isoelectric point. 1. With pH's below the pHZPC the solid has would have anion exchange capacity?

At low pH the solution in contact with the basal oxygen surface of the tetrahedral sheet will contain excess protons.

The surface will then exhibit an anion exchange capacity

2. pH's at the pHZPC , the solid would have no exchange capacity.

At a pH the corresponds to the ZPC (isoelectric point) the solution in contact with the basal oxygen surface of the tetrahedral sheet will contain a balance of protons and hydroxyls.

The surface will then exhibit no exchange capacity

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3. pH's above the pHZPC, the solid would have cation exchange capacity.

At high pH the solution in contact with the basal oxygen surface of the tetrahedral sheet will contain excess hydroxyls.

The surface will then exhibit a cation exchange capacity

Table of pH for zero point of charge for clay minerals. Note that in Georgia Piedmont soils, the typical pH is below 4, in Georgia estuarine systems typical pH is 7.5.

Mineral pHZPC Gibbsite 10 Hematite 4.2 - 6.9 Goethite 5.9 - 6.7 Na-feldspar 6.8 Kaolinite 2 - 4.6 Montmorillonite <2 - 3 Quartz 1 - 3

Note that Al and Fe hydroxides have a high pHZPC. Kaolinite and montmorillonite have low PHZPC.The ZPC is determined from a titration curve where pH is varied. There are numerous models devised to predict the distribution of surface-species. One of the simplest models that give a reasonable representation of the ion distributions is the eeffect of pH on the ion exchange properties of allophane and smectite (Warren and Rudolph, 1997) are compared in Figure (26). (Allophane is a noncrystalline hydrous aluminosilicate.) Allophane has a well defined PZC (point of zero charge) near pH 7 where the adsorption of Cl- and Na+ are equal. At lower pH, the allophane has a net positive charge and Cl- is adsorbed. At high pH, the surface

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(broken bonds) is negative and Na+ is adsorbed. Allophane may have considerable anion exchange capacity at low pH and a large cation exchange capacity at high pH. The large negative charge due to isomorphous substitutions dominates the ion exchange characteristics of the smectite at all pHs. Adsorbed Cs+ always exceeds 700 mmol/kg. Changes in the character of the bonds at the edges of the flakes produce a small anion exchange capacity at low pH (Cl- adsorption) and create a small increase in adsorbed Cs+ at high pH values.

Figure 27. The effects of pH on the ion exchange properties of allophane and smectite (after Warren and Rudolph, 1997).

3.3 Measurement of Cation Exchange Capacity To measure exchangeable bases of the clay minerals, samples were saturated with ammonium chloride so that the exchangeable bases (calcium, magnesium, potassium and sodium) were replaced by ammonium ions. The exchanged bases in the solution were then measured using the many technical procedures.The measurement of cation exchange characteristics is not a simple task due in large measure to the changes occurring in some systems as a function of pH.

The main problems to be aware of in the measurement of CEC include. 1. The presence of soluble salts and carbonates. 2. The effect of cation and anion type. 3. The pH effect and the use of buffered solutions. 4. Ionic strength effects. 5. Methods used to account for, or to remove entrained electrolyte.

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Three general approaches have arisen as detailed in the chapter by Sumner and Miller (1996). 1. Cation Exchange Capacity, CEC, by Direct Method (Cation Exchange Capacity, cmoles or mmoles/kg). CEC is determined when 5 g of air dried clay (1 gm peat) is leached with 60 mL 1 M NH4OAc, pH 7, to saturate exchange sites with ammonium ions. Excess free ammonium ions are rinsed from the clay with isopropyl alcohol. The remaining ammonium ions held on cation exchange sites are replaced by leaching the soil with successive aliquots of a solution of 10% KCl acidified to 0.005 N HCl. Ammonium is determined on the KCl leached by distillation and titration (Tecator-Kjeltec Auto 1030 Analyzer), or by colorimetric on a Technician Auto Analyzer using the salicylate /nitroprusside method. 2. CEC is determined by the summation of exchangeable Ca, Mg, Na, K, and H+. Base cations are extracted by leaching 3 g air dried clay (1 g for peat) with successive aliquots of 1 M NH4OAc, pH 7, to total 60 mL. The concentrations of the base cations in the leached are determined by ICP-AES. The CEC is calculated from the sum of the base cations and exchangeable H+. 3. The third procedure is Buchner Funnel Procedure (Page eds., 1982). In this procedure solutions to be analyzed for calcium and magnesium were diluted by 20 percent with 5 percent lanthanum chloride to reduce interferences with dissolved silica oxide. Base cation concentrations were analyzed with a Perkin Elmer Model 3030 Atomic Adsorption spectrophotometer. Absorp- tion mode was used for measuring calcium and magnesium whereas sodium and potassium were measured by emission. Calibration standards for each of the cations were made in 1 mg/L, 2 mg/L, 5 mg/L and 10 mg/L concentrations and the values (in mg/L) were calculated in relation to the standards (Equation 1). The samples had to be diluted with deionized water by a factor of 10 to find the exchangeable potassium, by a factor of 20 to find the exchangeable magnesium and a factor of 50 to find the exchangeable calcium.

Equations 2 and 3 show how the dilutions and the lanthanum chloride dilutions were taken into account in calculating the exchangeable bases. No dilutions were necessary to determine the exchangeable sodium.

Equation 1: Calculating exchangeable base concentration (meq/100g) meq/100g = [(Extract volume (mL)) / (Soil weight (g)] * 0.1 * mg/L

Equation 2: Calculating concentration with Lanthanum dilution

Ca and Mg meq/100g = 0.8 * (meq/100g)

Equation 3: Calculating concentration based on dilutions

X: 1 dilution = X * (meq/100g)

See in below website the interactive animation of cation exchange capacity measurement. http://hintze-online.com/sos/1997/Articles/Art5/animat2.dcr

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Typical CEC values in cmoles/kg for common clay minerals reflect their basic layer structure and layer charge differences are shown in Table (5). When applied to industrial minerals, the CEC is therefore determined mainly by the amount and type of clay minerals present in the deposit (Van Olphen, 1963).

Table 5. CEC values for clay minerals and organic material

Clay minerals CEC (meq/100 g)

Vermiculite 120-150

Montmorillonite 80-120

Illite 20-40

Chlorite 20-40

Kaolinite 1-10

Organic matter 100-300

3.4 Swelling Properties of smectite

The interlayer in smectites is not only hydrated, but it is also expansible; that is, the separation between individual smectite sheets varies depending on: (1) the type of interlayer cations present (monovalent cations like Na+ cause more expansion than do divalent cations like Ca2+), (2) the concentration of ions in the surrounding solution, and (3) the amount of water present in the soil. Because the interlayer is expansible, smectites are often referred to as "swelling clays". Soils having high concentrations of smectites can undergo as much as a 30% volume change due to wetting and drying or these soils have a high shrink/swell potential. These dramatic changes in soil volume are responsible for the properties of soils in the Vertisols order, which form deep cracks upon drying.

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CHAPTER FOUR

4 METHODS OF CLAY MINERALS INVESTIGATION The methods used for the identification and quantification of aluminosilicates include X-ray diffraction, electron microscopy, energy-dispersive X-ray analysis, infrared spectroscopy , differential thermal analysis, and scanning electron microscope. X-ray diffraction leads to an understanding of the structural characteristics of the clay mineral. Infrared absorption has limited usefulness but as more knowledge is gained in this field is becoming a more important adjunct to other tests. The electron microscope is a powerful tool in clay mineralogy but involves a very special technique and requires expensive equipment. Differential thermal analysis reveals information on the behaviour of clay minerals and many times, but not always, permits identification. For additional information on these identification techniques X-Ray Diffraction and the Identifica- tion and Analysis of Clay Minerals see Moore and Reynolds (1997).

4.1 X-Ray Diffraction Experiment

X-rays are electromagnetic radiation of wavelength about 1 Å (10-10 m), which is about the same size as an atom. They occur in that portion of the electromagnetic spectrum between gamma-rays and the ultraviolet. The discovery of X-rays in 1895 enabled scientists to probe crystalline structure at the atomic level. X-ray diffraction has been in use in two main areas, for the fingerprint characterization of crystalline materials and the determination of their structure. Each crystalline solid has its unique characteristic X-ray powder pattern which may be used as a "fingerprint" for its identification. Once the material has been identified, X-ray crystallography may be used to determine its structure, i.e. how the atoms pack together in the crystalline state and what the interatomic distance and angle are etc. X-ray diffraction is one of the most important characterization tools used in solid state chemistry and materials science. We can determine the size and the shape of the unit cell for any compound most easily using the diffraction of x-rays (Fig. 27)

Figure 28. Reflection of x-rays from two planes of atoms in a solid.

The path difference between two waves:

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2 x wavelength= 2dsin(theta)

For constructive interference between these waves, the path difference must be an integral number of wavelengths: n x wavelength= 2x

This leads to the Bragg equation: n x wavelength = 2dsin (theta) EXAMPLE 1 Unit Cell Size from Diffraction Data The diffraction pattern of copper metal was measured with x-ray radiation of wavelength of 1.315Å. The first order Bragg diffraction peak was found at an angle 2theta of 50.5 degrees. Cal- culate the spacing between the diffracting planes in the copper metal. The Bragg equation is n x wavelength = 2dsin (theta) We can rearrange this equation for the unknown spacing d: d = n x wavelength/2sin (theta). Theta is 25.25 degrees, n =1, and wavelength = 1.315Å, and therefore d= 1 x 1.315/ (2 x 0.4266) = 1.541 Å In this lab you will measure the x-ray powder diffraction pattern from a single crystal. Your TA will give you the sample to be measured and show you how to set up the Miniflex x-ray diffract meter. You should measure all the values of 2theta from the chart, and after converting those into d values calculate the repeat distance in your unit cell. In your lab note book list all the 2theta values with their corresponding values of n and d. Then calculate the mean and median values of the unit cell. Normally in powder X-ray diffraction studies we would want the mineral grains to be oriented randomly on the glass slide. But for clay minerals, the most diagnostic "d" spacing is between the {001} planes. So, when the grains are placed on the glass slide they are usually placed in a few drops of water so that they will settle onto the slide with their {001} planes parallel to the slide. Thus, when we X-ray them, we get diffraction predominantly off of the {001} planes and can measure the "d" spacing between these planes.

4.2 Clay preparation for XRD

4.2.1 Procedure of Separation clay fraction from bulk sediments

After preliminary removing of sand fraction (>63µm) with wet sieve, clay is separated from silt by centrifugation or sedimentation from suspensions. To do this, we first disaggregate the sample and place it in a settling tube filled with water. Particles will settle in the water according to Stokes Law:

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2 V = 2/9(ρg - ρw) g r /η Where: V = the settling velocity 3 ρg = density of the mineral grain (2.6 - 2.8 g/cm for clay minerals) 3 ρw = density of water (1g/cm ) g = acceleration due to gravity (980 cm/sec2) r = radius of the mineral particle (10-4 cm for clays) η = viscosity of water (10-2 gcm/sec2).

For deflocculating of particles, usually a disaggregating agent is added to the water to keep the individual particles from adhering to one another. The particles are placed in a large glass cylinder filled with water and the disaggregating agent, and the mixture is stirred. Three dispersing agents, 1% sodium polyphosphate (Petschick et al., 1996) and 0.0005 mol/l sodium hexametaphosphate (Shirozu, 1988) and 0.0005 mol/l sodium pyrophosphate (Moore and Reynolds, 1989), were attempted to compare effectiveness for particle separation and influence for clay minerals. Although Na-polyphosphate shows better ability for dispersion than that of Na- hexametaphosphate, it obviously attacked carbonate minerals (calcite peak was disappeared). Silt fraction (2-63µm) and clay fraction (<2µm) were separated by centrifuging at 700 rpm three time using modified Stock’s law (Shirozu, 1988). All extracted clay fraction were gathered in one tube and then washed by pure water for three times to remove dispersing agent and got clay suspension. There is several method of preparation oriented mount of the clay from the suspensions. (a) The easiest, quickest and probably most commonly used method of preparing oriented clay was the Stokes law to figure out how far particles of clay size will settle in a given time. That distance is measured on the cylinder, and that amount of water is then poured off and collected. It is then poured through a filter to separate the clay minerals from the water. The filter is then dried and the clay minerals are placed on a glass slide ready for X-ray diffraction analysis. The preparing oriented clay mounts for XRD is to drop clay suspension on a glass slide and let it dry at room temperature. Having arrived at the point of having a clay suspension, the next step is the production of an orientated clay film. The ideal is to have the (00l) planes exclusively orientated parallel to the sample surface. This is rarely if ever achieved. A high degree of preferred orientation allows the basal spacing’s to be used for the diagnostic tests that are required for clay analysis. Non-clay fines tend to disrupt the degree of preferred orientation. (b) Another method but not so easy is the filtration technique (Drever, 1973). The clay suspension is vacuum filtered on (0.45 µm pore diameter) filter and laid face down on a glass slide. First filtration to drain all fluid was carried out within 5 minutes to avoid grain sedimentation on the filter. After filtration was completed entirely, clay film on the filter was carefully transferred onto a glass slide. At this time, one drop of pure water was previously put on the slide to remove air bubble between the clay film and the slide. Fi- nally, the clay film was dried for 15 minutes at 50 ºC with the filter. The filter was slowly removed after it become white from semitransparent.

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The whole x-ray peaks of clay appear between 2-30 2θ X-ray diffractometry. Figure (28) shows the outline of the sample processing to separated clays from bulk sample and to analysis the clay minerals in both procedures.

4.2.2 Glycolation Organic liquids, primarily ethylene glycol and glycerol, are extensively used as an auxiliary treatment to expand swelling clays. Whether or not a mineral expands and the amount of expansion can provide essential supplementary information aiding clay-mineral identification. Swell- ing clays include smectites (e.g. montmorillonite, nontronite, and beidellite), some mixed-layer clays, and vermiculite. Two methods are presented below: a vapor treatment and a rapid method. The advantage of the vapor treatment is fewer disturbances of the sample and less amorphous scattering of X-rays by excess liquid than in the case of the rapid method.

The vapor treatment method is to keep the clay aggregate for long enough time in ethylene glycol (EG) vapor at elevated temperature (24 h at 80 ºC). Longer times will not hurt samples. Do not remove mounts until they are to be run on the X-ray diffractometer. The rapid method has to be done by adding of ethylene glycol directly to the surface of wet clay aggregate and letting it almost dry. This can disturb the orientation and the quality of pattern is often poor, because the aggregate has to be run wet. Glycolation shifts the low charge vermiculite and Smectite 001-peak position from 12-15Å to 16-17Å.

4.2.3 Heating Heat treatments at various temperatures are commonly used to help identify clay minerals by re- vealing changes in crystal structure spacing’s or loss of the structure. Depending on the temperature and the mineral species, these treatments can collapse the structure by dehydration, or in the case of other minerals destroy the crystal structures. However, it is important for the analyst to remember that some of the changes caused by the heat treatments may be temporary, and that partial or complete rehydration may occur during cooling. The heated treatment apply at 350 and 550 °C and the processing of heating as follows: Place the oriented aggregate mount in the furnace using the tongs. Leave sample in the furnace not less than one half hour at 350 °C; Remove mount by pulling it forward with the wire hook until the edge of the mount can be grasped with the tongs. Do not remove mounts until they are ready to be run on the diffractometer; X-ray the sample and repeat the above procedure at 550 C.

4.2.4 Differentiation between Clays The approximate basal spacing (in Angstroms) for the major clay mineral groups are shown in Table (6). The basal spacing of peaks on your scan can be read from the bottom of the screen. The distinction between kaolinite and chlorite may be difficult, particularly if the (001) and (003) chlorite reflections are weak. Usually, however, if both minerals are present, the chlorite (004) at about 3.5 Å is slightly offset from the kaolinite (200) reflection.

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Table 6. Values of basal reflections of clay minerals in Å, N (untreated samples), EG (ethylene glycolated) and 500(heated sample to 550ºC (after, Thorez, 1975).

Clay mineral symbol 7 8 9 10 11 12 13 14 15 16 17 Å Kaolinite K N EG Illite I N EG 500 Smectite M 500 N EG Chlorite C N EG 500 Palygorskite Pal N EG 500

Clay Group (001) (002) (003) (004) Kaolinite 7.16 to 7.17 3.57 to 3.59 2.38 to 2.39 ~1.8 Mica(illite) 9.98 to 10.10 4.94 to 4.98 3.31 to 3.36 ~2.5 Smectite 14.00 to 14.40 7.18 to 7.20 4.79 ~3.5 Chlorite 14.00 to 14.30 7.05 to 7.18 4.68 to 4.76 ~3.5

o Heat treatment

With heating the following changes take place:

1. The kaolinite structure is destroyed.

2. The micas are unaffected.

3. The basal spacing of the smectites collapses to 9.5-10.4 Å.

4. in chlorites (001) will commonly increase in intensity and may shift to about 13.8 Å. The (002), (003), and the (004) reflections decrease in intensity.

o Glycolation

With glycolation only the smectite group and the smectite layers in mixed-layer illite-smectite are affected. Their basal spacing increases to approximately 17 Å.

Notes:- a. Iron-rich chlorites (where Fe fills more than ~30% of the octahedral sites) have relatively weak (001) and (003) reflections and strong (002) and (004) reflections and are thus easily con-

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fused with kaolinite. The 4.7 Å peak of chlorite, even if quite weak, is usually diagnostic, unless the chlorite is quite high in Fe, in which case it may be missing altogether. b. If the kaolinite structure does not break down at 550 °C, do not heat to a higher temperature than 600 °C or you will melt the glass slide. Try increasing the heating time by one hour; however, do not exceed approximately six hours as water will eventually leave the other clay structures also. c. The 14 Å peak of chlorite tends to be sharper than that of a smectite. d. Examples of powder X-ray patterns for the common clay minerals are displayed in the X-ray lab along with brief explanations. Use these to help you with your interpretations.

4.3 Semi-quantitative analysis of clay minerals

Semi-quantitative assessments make the identification of individual components in polymineralic samples much more valuable. Unfortunately, the intensity of a mineral's diffraction peaks can not be directly used as an accurate measure of abundance because sample mounts and X-ray ma- chine conditions vary, and because different minerals, different atomic planes within a mineral, and different samples of the same mineral do not have the same ability to diffract X-rays (Bis- caye, 1965). However, Biscaye (1965) also found that useful semi-quantitative comparisons can be made between samples by means of various ratios of peak areas. These ratios vary in part due to mineralogy and in part due to scattering factors inherent to X-ray diffraction. For example, a 17-angstrom peak will have four times the intensity of a ten-angstrom peak if a two theta com- pensating device is not used (Borchart, 1989).

Semi-quantitative mineral analyses used XRD data obtained from the random mounts and the ‘external standard’ method of Chung (1974). This procedure involves preparing oriented clay mounts for XRD. XRD analysis of oriented clay samples were made after air drying at room temperature and ethylene-glycol solvated conditions. The intensities of selected XRD peaks (through 2-40º 2θ) characterizing each clay mineral present in the size fraction (e.g. kaolinite 7.1Å, smectite 14Å, illite 10Å,chlorite 14.3Å, Palygorskite 10.5Å and sepiolite 12.2Å) are measured for a semi-quantitative estimate of the proportion of clay minerals present in the size- fractions <2µm and 2-16µm. According to Moore and Reynolds (1989), a good way to approach the interpretation of XRD patterns is to work one peak at a time. Note its position on the air- dried sample diffract gram and look for changes with EG saturation and heating.

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Sample

Normal sedimentation By Centrifugation

<2 µm fraction <0, 2 µm fraction

Air-dried clay films

Heat-treatment EG - Solvation At 300 C and 500 C

Clay minerals identification

Semi quantitative Qualitative result

Figure 29. Sample processing flow chart to prepare clay films and analysis clay minerals.

Remember that diffraction peaks should appear for all the other d (00l) reflections within the two theta interval scanned if there is no mixed layer. If you observe a 10Å peak, look for 5.0 and 3.5Å peaks, or even a 20Å peak. Comparison of peak positions on the two diffractograms (Fig. 29) for sample 353 from the Lanson and Bouchet (1995) article produces a preliminary determination of the minerals present in a sample from the subsurface. The 7.18Å peak on the air-dried sample (AD) is unchanged (7.17A) on the EG saturated sample (EG) suggesting kaolinite and/or chlorite. The 10.03 and 9.94Å peaks indicate an illite that is more than 90% illite layers The 14.44 and 14.39Å peaks can be assigned to chlorite and/or vermiculite. The new 16.74Å peak on the EG pattern and reduced diffracted intensity between the 10 and 14Å peaks are indicative of smectite-rich materials. Departure from ideal peak positions for the pure layer types may be due to compositional differences or randomly interstratified mixed layers. Resolution of the “and/or” assignments above requires further testing. For example, kaolinite and chlorite may be distinguished by scanning the 3.5A region where the kaolinite (002) and chlorite (004) peaks are re- solved.

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Figure 30. Decomposition with 3 elementary peaks of the XRD pattern obtained from sample (AD) and with 4 elementary peaks obtained from sample (EG). These peaks are associated with minerals I/S (16.74Å), chlorite (14.38Å), illite (9.9Å) and kaolinite (7.18Å).

4.4 Infrared spectroscopy Infrared spectroscopy (IR) is widely used to determine and investigate the structure of the various mineral phases. It provides information ranging from the detection and identification of specific or minor mineral constituents, hardly accessible from X-ray diffraction techniques, to the determination of the stacking order and ordering pattern of substituting cations in clay minerals (Farmer, 1974). In the field of clay mineralogy, study of infrared absorption spectra has lately made a remarkable progress, as exemplified by numerous researches on the absorptions in the OH region, on the relationships between the chemical composition and the variation of position of absorption bands, and on the after-heating variation of absorption bands. Concerning the OH region a large number of researches have been made, particularly on absorption of kaolin minerals (Yoon, et al. 1992).

The clay minerals powder (0.5gm) were mixed with 300 mg of KBr disk, previously dried at 110 ºC and then pressed in vacuum by the oil press. The pressure exerted on the specimen was 10 tons per 1 cm2. Thus, a disk of 10-13 mm in diameter and 1-2 mm in thickness was made. Russell (1987) showed that better resolution can be obtain by using KBr instead KI as dispersant for the preparation of the disks.

The advantages of IR analysis compared to X-ray powder diffraction in principle, are that with IR technique both the crystalline minerals and amorphous materials can be investigated, especially for minerals that could be found in trace percentage and out of XRD sensitivity. For more careful work, and particularly for the quantification of clay minerals such as kaolinite and chlorite that have overlapping reflections (particularly if the (001) and (003) chlorite reflections are weak), the IR technique is best method used to distinction between kaolinite and chlorite. In

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spite of these advantages, the IR methods are still less widespread for quantitative determination of clay and other minerals. The greatest problem is that the material of equivalent chemical composition but of different particle size and crystalline state has considerably different spectra (Gadsden, 1975). It is remaining best to use the two methods to complement rather than substitute one another. The absorbing bands for most clay minerals and associated minerals as given by Farmer (1974) shown in Table (7).

Table 7. Absorption bands of some clay minerals, data from Farmer (1974) and Gadsden (1975).

Mineral Absorption bands (cm-1)

Kaolinite 3695, 3660, 3625, 1035, 1020, 915 Halloysite 3696, 3624, 3414, 1035, 1005, 910 Montmorillonite-Illite 3635, 3400, 1640, 1130, 1020, 920 Chlorite 3586, 3560 3436, 1004, 980

The infrared technique has used in the studying of kaolinite from Virtasalmi to identify the crystallinity and purity of kaolin. Band assignment is listed in Table (8) were determine by the base line method according to Hlavay et al. (1977) and Characteristic absorption value was calculated.

IR spectra clearly showed predominance of kaolinite in all studied samples, and does not exhibit any peak for impurities minerals such as Smectite (Fig. 30). The spectral region between 800 and 750 cm –1 is very sensitive against the crystallinity and purity of the kaolinite mineral. Pure kaolinite exhibits two peaks (790 and 755 cm –1) in this region and a well-crystallized kaolinite exhibits two sharp peaks at 3690 and 3620 cm –1 (Giese, 1988). In well crystalline kaolinite the 3670 and 3650 cm –1 band being much weaker than those at 3690 and 3620 cm –1 or no existing, but the broading of the 3670 and 3650 cm –1 band are typical of disordered kaolinite. The IR spectrum was identified that the most of Virtasalami kaolin was pure and well crystalline kaolinite.

Table 8. Band assignment of pure kaolinite (Hlavay et al. 1977).

Wave number (cm –1) Assignments

3690,3660,3620 O-H Stretching vibrations 1120, 1040, 1020 Si-O Asymmetrical stretching vibrations

700, 420, 432 Si-O Bending vibrations 940, 929 Al-OH Bending vibrations

795, 760 Si-O-Al Compounded 540 Si-O-Al Compounded

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Figure 31. IR spectra of untreated kaolinite fraction <2μm from Virtasalmi kaolin deposits (after Thair Al-Ani et al., 2004).

4.5 The Scanning Electron Microscope (SEM)

The Scanning Electron Microscope (SEM) is a microscope that uses electrons rather than light to form an image. There are many advantages for using the SEM instead of a light microscope. The SEM method provides detailed images of individual grains of clay minerals and EDS provides detection of major and minor elements at points on grain surfaces, allowing highly reliable identification from crystal form and composition, as well as direct observation of particle packing and size. The SEM also produces images of high resolution, which means that closely spaced features can be examined at a high magnification >3000 times or range. Preparation of the clay samples is relatively easy since most SEM require the sample to be conductive. A few thin sections of clay-size fractions coated with an Au-Pd conductor were examined morphologically under Scanning Electron Microscopy (SEM) with Link (EDS). The thin sections of representative samples were analyzed for semiquantitative determination of primary and secondary minerals.

Morphologies, as observed by scanning electron microscopy (SEM), can also be useful in identification of clay minerals. Kaolinite shows a variety of morphologies, including platy, pseudohexagonal particles, booklets and vermicular stacks. Halloysite also shows generally tubular morphologies although spherical particles are also common. Illite appears as sheets or large flack crystals or as fibres and chlorite clusters of bladed or platy crystals.

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4.6 Transmission electron microscopy (TEM)

TEM can be used for observation of single clay mineral grains. Stacking of layers, thickness of domain and nature of interlayering can be observed, provided that resolution is good enough (couple of Å).

Clay-size fractions were prepared for a TEM study by dispersing the material in alcohol. The samples were placed on a carbon - coated metal grid gives preferred orientation and allows observation of clay flaks and examined with TEM equipped with the same microanalysis electron microscope. Various magnifications were used to obtain suitable micrographs of clay minerals. Electron microscopy is the only method capable of measuring the size of individual single particles. The method allows for direct measurement of the several dimensions of the particles and thus also the shape of the particles is to be taken into account (Bates, 1971).

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CHAPTER FIVE

5 FORMATION AND ALTERATION OF CLAY MATERIALS

5.1 Introduction The formation and alteration of clay minerals and their accumulation as clay materials can occur by a very wide range of processes. In one way or another, however, most of these processes and the environments in which they operate involve the chemical actions and physical movement of water. As such, clay minerals can be considered the characteristic minerals of the Earths near surface hydrous environments, including that of weathering, sedimentation, diagensis/low-grade metamorphism and hydrothermal alteration (Fig. 31).

Figure 32. The geological time-temperature in formation and alteration processes of clay minerals (after Reeves et al., 2006).

Simply defined, the weathering environment is that rocks and the minerals they contain are altered by processes determined by the atmosphere, hydrosphere and the biosphere. Soil formation, also known as pedogenesis, occurs in the weathering environment. The sedimentary environment is the zone in which, soil, weathered rock and mineral (and biogenic) materials are eroded, mixed and deposited as sediments by water, wind and ice. Diagenesis involves all those physical and chemical processes that occur between sedimentation and metamorphism, whilst hydrothermal alteration encompasses the interactions between heated water and rock. Some of the important environmental controls, processes, and sediment components involved in early diagenetic reactions, along with some of the more distinctive clay material products that may result are summa- rized in Figure (32).

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Environmental observation of clay mineral transformations and neoformations in shear zones, it is concluded: Three successive alteration stages can roughly be distinguished in terms of kaolinization, illitization and montmorillonitization. The driving force of alteration is the chemical energy of low-temperature solution transfers. The main controlling factor of clay mineral transformation sequences is the decrease in permeability as a consequence of argillation (Riedmiiller, 1978).

EARLY DIAGENESIS

ENVIRONMENTAL CONTROLS PROCESSES COMPONENTS SOME CLAY PRODUCTS Oxidation Fe/Mn Oxides Brown/red clay Sedimentation/burial rate Reduction Black pyritic shales Marine or Non Marine Carbonates Fermentation Sideritic mudstones Sulphides Bioturbation Tonstein (kaolin) Stagnant or ventilated Authigenic clays Mechanical Bentonite (smectite) Compaction Organic matter

Figure 33. Summary of early diagenesis in terms of the main environmental controls, important processes, solid components affected, and examples of particular products that may result (after Reeves et al., 2006)

In this chapter, the origins of the various clay minerals that may occur in each of these environments are reviewed along with the processes that may lead to their accumulation and alteration, usually together with other components, to form clay materials. In many instances, clay materials are formed in one environment by the accumulation or alteration of clay minerals formed in others. Thus the geological history of a clay material, and consequently its properties and behaviour, may depend on many environments.

5.2 The clay cycle

Ultimately, the clay cycle (Fig. 33) begins with the formation of clay minerals and their accumulation in soils by the weathering of primary rock forming minerals and glassy volcanic ash. Ero- sion and transport of clay minerals from soils and weathering profiles followed by their selective sorting, segregation and deposition by physical processes of sedimentation lead to their further accumulation as the main components of muds in sedimentary basins. Burial of mud by further mud, or other sediments, may follow and in response to the physical and chemical changes that accompany burial (known as diagenesis) a progressive transformation occurs to change muds into mudstones and shales. Eventually, if tectonic forces are involved, mudstones and shales are transformed to slates. Inevitably, over geological time, tectonic uplift brings buried mudrocks (mudstone, shale and slate) back to the Earth's surface where the agents of weathering begin the cycle once again.

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Figure 34. Schematic clay cycle, whereby clay minerals formed in one environment are frequently recy- cled into others (after Reeves et al., 2006).

5.3 Geological origin of clay minerals

The interpretation of the origin of clay minerals is one of the most interesting aspects of clay mineralogy. Clays and clay minerals occur under a fairly limited range of geologic conditions. The environments of formation include soil horizons, continental and marine sediments, geo- thermal fields, volcanic deposits, and weathering rock formations. Most clay minerals form where rocks are in contact with water, air, or steam.

Recall that the nature of clay formed during the weathering process depends upon three factors: 1. The mineralogical and textural composition of the parent rock. 2. The composition of the aqueous solution. 3. The nature of the fluid flow (i.e., rate of water flow and pore network) The contact of rocks and water produces clays, either at or near the surface of the earth” (from Velde, 1985). Rock +Water → Clay For example, the CO2 gas can dissolve in water and form carbonic acid, which will become hydrogen ions H+ and bicarbonate ions, and make water slightly acidic. CO2+H2O → H2CO3 →H+ +HCO3- The acidic water will react with the rock surfaces and tend to dissolve the K ion and silica from the feldspar. Finally, the feldspar is transformed into kaolinite. The rock mineral weathering is one of the main natural sources of clay minerals and metal concentrations in the soil. The soils are open system. Accordingly, the faster the flow rate, the shorter the contact time of solution with the primary minerals. The stability diagram (Fig. 34) shows that clay minerals are stable under conditions near the surface. As shown in soils developed in Hawaii display the effect of different weathering products from a parent (basalt), smectites are formed at low rainfall, while kaolinite formed at moderate – high rainfall.

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Figure 35. Diagram shows the effect of rain fall versus the percentage of clays formed in soils (after Sherman, 1952).

The physical and chemical conditions under which the kaolin minerals form are relatively low pressures and temperatures. These minerals are typical of three main environments: 1) weathering profiles; 2) hydrothermal alterations; and 3) sedimentary rocks. The most common parent minerals from which kaolin minerals develop are feldspars and muscovite. The transformation of potassium feldspar into kaolin minerals occurs according to the equation:

2 KAlSi3O8 + 3 H2O Al2Si2O5(OH)4 + 4 SiO2 + 2 K (OH)

Solubility of the several chemical species is pH dependent (Mason, 1952). The pH values of the natural waters normally lie between 4 and 9; alumina is not soluble in this range; silica solubility increases parallel to the pH and the alkalis and alkaline earth elements are soluble and mobile (Figure 35). Thus, kaolinite is easily formed and is widespread in soils developed under hot-wet, intertropical climates (Chamley, 1989). As a consequence, detrital kaolin minerals are important components of sedimentary rocks deposited near these areas. In addition, kaolin minerals frequently grow, from the same phases (feldspars and white mica), during the early diagenesis. The major tasks in understanding the geologic significance of clay minerals are to isolate and to identify the signatures of the various factors and the affects of geologic processes responsible for the present-day properties of the clay-rich materials. The composition may represent a single stage of clay formation or be the result of multiple generations. The accurate study of clay mineral resources requires, however, the knowledge of the main methods of identification and differentiation among the several clay minerals as we explain in this book. These methods include: X-ray diffraction and infrared spectroscopy, scanning electron microscopy (SEM) study and transmission electron microscopy (TEM),

The reading material in this chapter is only intended to provide a few examples of specific studies concerning the geologic origin of clay minerals. Each should be considered with respect to the general themes outlined in the Introductory Section.

Virtasalmi Kaolins in south eastern Finland at the most basic level of discrimination, the residual materials of the weathering. According to the Sarapää (1996) Virtasalmi kaolin was formed in situ by chemical weathering processes of parent rocks. Mineralogical, chemical and micro tex-

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tural analysis of the studied kaolin deposits showed that the progressive weathering of plagioclase and biotite of the parent rocks are the main source of halloysite and kaolinite in the studied area. The colour and quality of kaolin are controlled by the type of the weathered bedrock. White kaolin mainly derives from quartz-feldspar gneiss and mica gneiss, whereas coloured kaolin from amphibolites and diopside amphibolites.

Other clues to their origin may be revealed by unique textural, mineralogical or chemical characteristics. Clay minerals have diverse origins that make them very useful for paleoclimatic recon- structions, interpretation of the tectonic history of a region, determination of provenance, definition of the timing and nature of digenetic reactions, and many others.

Figure 36. Superposed curves of the solubility of silica and of alumina as functions of pH (after Kraus Kopf, 1956). Maria Dolores Ruiz Cruz (2006) in the studying of the genesis and evolution of the kaolin-group minerals during the digenesis and the beginning of metamorphism concluded four processes that are responsible for the evolution of kaolinite minerals during the burial- or tectonic digenesis.

1. The kaolinite → dickite → nacrite transition is well documented in natural environments and occurs at increasing temperatures. 2. Thermodynamic and experimental data indicate, however, that the stable phase, at the common P-T conditions, is kaolinite. Thus, the described transformations are not only controlled by temperature. 3. Strain and solutions composition appear to be important factors controlling the dickite and nacrite formation.

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4. Futures researches must be focused in natural fluid characterization by means of stable iso- topic and fluid inclusions studies. Parallel synthesis in well-controlled chemical systems must supply additional information.

Walter Keller in several review articles identified six processes that are responsible for the formation of clay minerals (i.e., Keller, W. D., 1964). 1) “Crystallization from solution.” Essentially equivalent to clays considered authigenic or a product of neoformation. 2) “Replacement by clay minerals” (no specific reference to the replacement process). 3) “Weathering of silicate minerals and rocks (not clay minerals) dependent largely on relative activity of H. 4) Weathering of other clay minerals. 5) Diagenesis, reconstitution, and ion exchange. 6) Hydrothermal alteration of minerals and rocks. The geochemical and mineralogical relationships associated with degradation, rejuvenation, ag- gradations and neoformation can be considered as subtractive or additive as indicated in a similar presentation from Millot's (1979), (Fig. 36). As long as the temperature, pressure, and solution composition are constant a clay assemblage may not change. If conditions of the environment are altered, particularly solution composition, the minerals will evolve by adding or subtracting elements from exchangeable and structural positions. The environments associated with evolution by subtraction include weathering and soils, subaqueous alteration (halmyrolysis) and degrading diagenesis. The model (Figure 36) enclosing weathering and soils on the lower left is there to emphasis that this environment is the most important of those associated with subtractive processes. Processes resulting in addition of chemical constituents occur in soils, during sedimentation, and during diagenesis. Additive changes are most important in sedimentation and diagenesis. Mineral changes inferred in these diagrams often involve mixed-layer clays and the nature of the mixed-layers formed by the various processes may be different. These diagrams should help to envision the diverse origins and potential complexity of clay mineral assemblages. They may give some good ideas about how to approach the geologic interpretation of minerals in fine- grained materials, a significantly greater challenge than those facing students working on sand- stones or carbonates.

Figure 37. Origin of clay minerals by Millot

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CHAPTER SIX

6 INDUSTRIAL USES OF CLAYS Clay minerals are some of the most important, if not the most important, of our industrial minerals. Millions of tons are utilized annually in a large variety of applications. These applications include uses in geology, the process industries, agriculture, environmental remediation and construction. This book focuses on the major applications of the clay minerals today and looks into the future growth and applications of certain specific clay minerals. Why are certain clay minerals used in specific applications? The reason is that the physical and chemical properties of a particular clay mineral are dependent on its structure and composition. As we explained in chapter two, the structure and composition of kaolins, smectites, and palygorskite and sepiolite are very different even though they each have octahedral and tetrahedral sheets as their basic building blocks. However, the arrangement and composition of these octahedral and tetrahedral sheets account for major and minor differences in the physical and chemical properties of kaolin, smectites and palygorskite.

The important characteristics relating to the applications of clay minerals are particle size and shape, surface chemistry, surface area, surface charge, and other properties specific to particular applications, including viscosity, colour, plasticity, green, dry and fired strength, absorption and adsorption, abrasion and pH. In all applications, the clay minerals perform a function and are not just inert components of the system. Several authors those discuss clay mineral applications and, although some are historical, they are essential to our present understanding of how and why the clay minerals have such an extensive industrial utilization. Some of these are Murray & Lyons (1956), Grim (1962), Clem & Doehler (1963), Haden (1963), Jordan (1963), Konta (1995), Murray (1984, 1991, 1986), Grimshaw (1971), Grim & Guven (1978), Robertson (1986), GalaÂn (1996), Elzea & Murray (1994), Heivilin & Murray (1994), Pickering & Murray (1994), Keith & Murray (1994) and Murray (2007).

Table (9) shows some of the properties of kaolin, smectite and palygorskite that account for many of their applications. Specific applications are discussed under the headings kaolin, smectite, and palygorskite and sepiolite.

6.1 Kaolin

Kaolin is soft, white plastic clay consisting mainly of the mineral kaolinite which is a hydrated aluminium silicate Al2 Si2 O5 (OH)4. It is formed by the alteration of feldspar and muscovite. Kaolin deposits are classified as either primary or secondary. Primary kaolins result from residual weathering or hydrothermal alteration and secondary kaolins are sedimentary in origin. Kao- lin is an important industrial mineral, which is used in many industrial applications.

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Table 9. Some properties of clay minerals that can be related to their applications. Kaolin Smectite Palygorskite

1:1 layer 2:1 layer 2:1 layer inverted White or near white Tan, olive green, white Light tan Little substitution Octahedral and tetrahedral Octahedral substitution Substitution Minimal layer charge High layer charge Moderate layer charge Low base exchange High BEC Moderate BEC capacity Pseudo-hexagonal flakes Thin flakes and laths Elongate Low surface area Very high surface area High surface area Very low absorption High absorption capacity High absorption capacity capacity Low viscosity Very high viscosity High viscosity

Three kaolin production areas dominate the world markets. These are the sedimentary Kaolins in Georgia and South Carolina in the United States; the primary kaolins in the Cornwall area of south western England; and the sedimentary kaolins in the lower Amazon basin in Brazil. These kaolins are of high quality in that they have high brightness and relatively low viscosity at high solids concentration (70%). This means that they can be used for paper coating which is the largest application. Other kaolin deposits, which are regionally important, are located in Australia, Argentina, Czech Republic, China, France, Germany, Indonesia, Iran, Mex- ico, South Korea, Spain, Turkey, and Ukraine. The physical and chemical properties of kaolin has led to its extensive use as filler, extender, paper coater, ceramic raw material, pigment, and also it is an important raw material for the refractory, catalyst, cement, and fibber glass industries. The total world production is currently estimated to be 39 million tons per year as distributed in Table (10)

Table 10. The total world production of kaolin per year . Paper Filling and Coating 45% Refractories 16% Ceramics 15% Fiberglass 6% Cement 6% Rubber and Plastics 5% Paint 3% Catalyst 2% Others 2% *Roskill Information Services, Ltd. A The Economics of Kaolin@ 10th Edition

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The major world production of kaolin was estimated by the USGS to be 37,800,000 tons, in 2007. The major production by country is shown as in Table (11)

Table 11. The major world production of kaolin USA 7,330,000 China 600,000 Brazil(beneficiated) 2,500,000 MEXICO 490,000 UK 2,100,000 Spain 460,000 Czech republic 1,049,225(beneficiated) TURKEY 450,000 Czech republic 5,183,000 (raw) Argentina 300,000 Germany 3,800,000 France 300,000 South Korea 2,400,000 Ukraine 300,000 Commonwealth of Inde- 6,000,000 Indonesia 250,000 pendent States(CIS) Iran 900,000 Australia 250,000 *USGS

All kaolin is mined using open pit methods utilizing shovels, draglines, and backhoes. The normal economic overburden to kaolin ratio is 6.5 to 1 or less. Two processing methods are used in the production of kaolin - dry and wet. The dry process is rather simple and a typical flow sheet is as follows:

MINING > CRUSHING > DRYING > PULVERIZATION > CLASSIFYING > BAGGING & LOADING.

The wet process is more complex with a typical flow sheet shown as follows:

MINING

BLUNGING

DEGRITTING

Fractionation and particle size separation. Selective Flocculation Magnetic Separation Delamination Flotation

LEACHING Surface Treatment

DEWATERING

Apron Drying DEFLOCCUATION High solids slurry

Calcination SPRAY DRYING

BAGGING & LOADING

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6.1.1 Paper industry

The largest single user of kaolin is the paper industry, which used approximately 20 million tons per year as shown in Table (10). Kaolin is uniquely suited for this use because of its fine particle size, platy particles, good viscosity, low abrasion, good opacity, white colour, high brightness, and good print quality. Several grades which vary in brightness and particle size are available to the paper coater. Prices (FOB) range from 4 to 8 cents per pound. An analysis of a sheet of paper in the National Geographic magazine would show that approximately 35% of the weight would be kaolin.

A major value added product produced by the US kaolin industry is calcined kaolin. Calcined kaolins have a high brightness and opacity and are used to extend titanium dioxide which is expensive ($1.00 per lb.). Titanium dioxide is a prime pigment with exceptionally high brightness and opacity. Calcined kaolin can replace up to 60% T O in paper coating and filling formulations and also in paint formulations without any significant loss in brightness and opacity. Calcined kaolins are produced by heating spray dried fine particle kaolins to temperatures in the range of 1,000 C. The kaolinite becomes anhydrous and transformed to mullite (Al2 Si2 O5) and SiO2. The price of calcined kaolin products range from 15 cents to 25 cents per pound.

The major competitive mineral is calcium carbonate, which has made severe inroads in paper filling. However kaolin still dominates the paper coating market. Kaolin is an excellent ceramic raw material and is a necessary ingredient in most white ware and sanitary ware applications because it fires white, is plastic, and has good shrinkage and strength properties.

Over 400,000,000 tons of kaolin has been produced from the deposits in Georgia in the USA. Within the next few years the kaolin production in Georgia will be downsized because of the depletion of reserves. The same is true in England but fortunately there are several hundred million tons of high quality kaolin reserves in Brazil, which will become the world production leader in this century.

Critical issues involved in the mining of kaolin vary with the geographic location of the deposits. Major issues in mining the kaolin in Southwestern England are the disposal of the waste material produced in the processing and the disposal of the 80 to 85% of the host rock from which the kaolin is extracted by hydraulic methods. Another problem is the depth of the ore body as mining progresses because these deposits are funnel shaped and become smaller in circumference with depth. Another issue with the English kaolins is their relatively high viscosity in comparison with the Georgia and Brazil sedimentary kaolins. As the paper coating machines run faster in order to produce more coated paper to improve productivity, the viscosity becomes more critical. Diagrammatic representations of particle packing in kaolin-water systems (Fig. 38) are useful to summarize the geometrical parameters that influence viscosity and emphasize the effect packing in papermaking. A well-sorted size distribution of kaolinite crystals is usually an indication of high viscosity. In schematic representation the lacks of bimodal distribution can be improved by blending small percentage of the fine kaolin during processing to get lower viscosity.

In Georgia, most of the low overburden kaolins have been mined. This means that as the overburden ratio increases, the cost of mining increases. Another critical issue is that the availability

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of high quality, low viscosity kaolins are rapidly disappearing so the industry must devise methods to improve the viscosity in order to use marginal reserves of high viscosity kaolins.

In Brazil, a severe problem is the heavy rains, which flood the mines during the rainy season. Also the remoteness of the mines from population centres creates difficulties in labour and management turnover.

Kaolin is an environmentally safe material with no adverse health problems as long as the fine particle dust is controlled. The open pit mines in Georgia and Brazil are reclaimed, so that the land can be used for agriculture, forestry, or recreational projects.

Kaolin particularly that which is used for coating paper is traded in a global market. Kaolin which is used for ceramics and fillers is usually restricted to regional or local markets. The paper coating kaolins are produced by large global companies but the ceramic and filler grade kaolins are produced by smaller companies who employ local marginalized people who are dependent on them for their livelihood. Kaolins are currently and in the future a large contributor to the economy of developing countries such as Brazil, Argentina, Tanzania, Indonesia, Suriname, and

Figure 38. Schematic representation of the packing of kaolin particles (Bundy et al., 1965). Dashed lines represent sphere of adsorbed water. (A) Monodisperse. (B) Polydisperse system.

India. Paper is a very necessary commodity and as a country develops, more kaolin will be needed to improve the printability of the paper particularly for colour printing. China is a good example where 15 years ago there was no paper produced in China that could be used for colour printing. Today China is the 3rd largest producer of paper and large quantities of kaolin are imported to produce high quality coated paper for internal consumption.

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6.1.2 Other Applications

Kaolin has many other industrial applications; some of the many uses of kaolin are shown in Ta- ble (12). Table 12. Industrial uses of kaolin.

Paper coating Cement Food additives Paper filling Pencil leads Bleaching Extender in paint Adhesives Fertilizers Ceramic raw material Tanning leather Plaster Filler in rubber Pharmaceuticals Filter aids Filler in plastics Enamels Cosmetics Extender in ink Pastes and glues Crayons Cracking catalysts Insecticide carriers Detergents Fibreglass Medicines Roofing granules Foundries Sizing Linoleum Desiccants Textiles Polishing composing

Another very large user of kaolins is the ceramics industry, particularly in white ware, sanitary ware, insulators, pottery and refractories. Both primary and secondary kaolins can have excellent ceramic properties (Murray, 1986). Kaolin which does not have the physical and chemical properties for use in a paper-coating application can have excellent ceramic properties. Halloysite, one of the kaolin minerals, is used as an additive in high quality dinnerware to provide translu- cency and strength. The major source of halloysite is on the North Island of New Zealand (Murray et al., 1977).

From the above, it can be seen that kaolins are indeed a valuable and versatile industrial mineral. Only a few kaolins in the world can be used for paper coating because of the stringent requirements for low viscosity and good colour. Many more deposits, however, can be utilized for ceramics and filler applications, so a careful evaluation of a kaolin deposit must be made in order to determine whether or not the material can be processed for some industrial use or use

6.2 Smectite (Bentonite)

Smectite is the mineral name given to a group of Na, Ca, Mg, Fe, and Li-Al silicates. The mineral names in the smectite group which are most commonly used are Na-montmorillonite, Ca- montmorillonite, saponite (Mg), nontronite (Fe), and hectorite (Li). The rock in which these smectite minerals are dominant is bentonite. Bentonite is smectite clay formed from the alteration of siliceous, glass-rich volcanic rocks such as tuffs and ash deposits.

The sodium, calcium, and magnesium cations are interchangeable giving the montmorillonite a high ion exchange capacity. The industrial bentonites are generally either the sodium or calcium variety.

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Bentonites are important and essential in a wide range of markets including drilling mud, foundry sand binding, iron ore pelletizing, pet waste absorbents, and civil engineering uses such as waterproofing and sealing. Bentonites have excellent rheological and absorbent properties. So- dium bentonite has a high swelling capacity, and forms gel-like masses when added to water. Calcium bentonite has a much lower swelling capacity than sodium bentonite but this can be improved by treatment with soda ash to produce sodium exchanged bentonite. Normally these sodium exchanged bentonites do not have as high swelling capacity as the natural sodium bentonites.

The largest sodium bentonite deposits are located in the Western United States in Wyoming, Montana, and South Dakota. These sodium bentonites are also called Western or Wyoming bentonite which means a high swelling sodium bentonite. Other smaller sodium bentonite deposits occur in Argentina, Canada, China, Greece, Georgia Republic, India, Morocco, South Africa, and Spain. Calcium bentonite deposits are much more common than sodium bentonite ones. In the United States calcium bentonites occur in Georgia, Alabama, Mississippi, Texas, Illinois, and Missouri. Elsewhere calcium bentonites occur in England, Germany, Spain, Italy, Greece, Tur- key, Georgia Republic, Czech Republic, Ukraine, Japan, Algeria, Morocco, South Africa, China, India, Japan, Argentina, and Brazil. The world production of all types of bentonite was estimated by the USGS to be 11,800,000 tons, in 2007. The major production by country is shown in Table (13). Table 13. The major world production of bentonite. USA 5,070,000 JAPAN 415,115 GREECE 1,100,000 GERMANY 360,000 Turkey 1,000,000 UKRAINE 300,000 CIS 750,000 BRAZIL 250,000 ITALY 470,000 Czech republic 220,000 MEXICO 450,000 OTHERS 1,300,000 *USGS

Bentonite has several important physical and chemical properties which make it important in a wide range of markets. In addition to its rheological and absorbent properties, bentonite has excellent plasticity and lubricity, high dry bonding strength, high shear and compressive strength, good impermeability and low compressibility. The many uses of smectites are shown in Table (14). Table 14. Industrial uses of smectites.

Drilling mud Medical formulations Crayons Foundry bond clay Polishing & cleaning agents Cement Pelletizing iron ores Detergents Desiccants Sealants Aerosols Cosmetics Animal feed bonds Adhesives Paint Bleaching clay Pharmaceuticals Paper Industrial oil absorbents Food additives Fillers Agricultural carriers De-inking of paper Ceramics Cat box absorbents Tape-joining compounds Catalysts Beer and wine clarification Emulsion stab Pencil leads

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The most important uses of bentonite will explain as follows:

6.2.1 Bentonite barriers in sealing nuclear waste

Over twenty-five years ago, bentonite was identified as a suitable material for use in sealing nuclear waste repositories in crystalline rock. The consideration of bentonite for this application was a natural consequence of its widespread use as a sealant in petroleum and civil engineering. The important properties of swelling, plasticity and low hydraulic conductivity were identified as key requirements for in filling materials that could be used around waste containers, as tunnel backfills and as seals for boreholes, drifts and shafts. Since that time, a greatly improved understanding has been developed of the basis for design and long-term performance assessment of bentonite barriers in geologic repositories.

Bentonite barriers are essential components of many repository designs for the geological disposal of highly radioactive waste. Considerable resources and time have been devoted to their study and testing, focusing on diverse aspects at different levels of detail, and different scale and time horizons. This effort is being pursued in a continuous pursuit of confidence building and optimisation, justified by the importance of bentonite barriers for the safety, and for the design, operation and closure of the repository systems.

Several European countries have integrated their skills, knowledge and experience in the BENIPA (Bentonite Barriers in Integrated Performance Assessment) project, within the Fifth Framework Programme of the European Union. The project was carried out from September 2000 to August 2003. Participants in the project are two National Agencies, responsible for national HLW management (ENRESA as Project Co-ordinator, from Spain and NAGRA, from Switzerland) and six Research Centres (GRS, from Germany; IRSN, from France; NRG, from The Netherlands; SCK-CEN, from Belgium; VTT, from Finland and ZAG, from Slovenia). For more information of uses the bentonite in technology of barrier system in backfilling and sealing of disposal, you can visit the website of Posiva (www.posiva.fi) to learn more.

6.2.2 Other uses of Smectite (bentonite)

Drilling mud, or drilling gel, has bentonite as a major component. Drilling mud is crucial in the extraction of drill cuttings during the drilling process. Bentonite, when mixed with water, forms a fluid (or slurry) that is pumped through the drill stem, and out through the drill bit. The bentonite extracts the drill cuttings from around the bit, which are then floated to the surface. Taconite, a low grade iron ore, has been developed as an economic source for iron. During processing, the taconite is ground into a very fine powder. The ground taconite is then mixed with small amounts of bentonite which serves as a binder to the taconite. This mixture is processed into balls or pellets. The process is finished when these pellets are sintered in rotary kilns that give the pellets a hard surface.

The metal casting industry needs bentonite as an economical bonding material in the molding processes associated with the metal casting industry. Bentonite, when mixed with foundry molding sands, forms a pliable bond with the sand granules. Impressions are formed into the face of the bentonite/sand mixtures. Molten metal is pored into the impressions at temperatures exceed-

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ing 2,800 F. The unique bonding characteristics of bentonite insure the durability of the mold during these high temperatures. Once the process is complete, the bentonite/sand mold can then be broken away from the casting face and reused. In recent years, bentonite has become a major component in the manufacturing of cat litter. Be- cause of the unique water absorption, swelling, and odor (smell) controlling characteristics of bentonite, it is ideal for use in "clumping" types of cat litters, 95% of all cat litter is made from clay (Fig. 39). Clumping cat litter has become widely accepted as an economical alternative to conventional non-clumping type cat litters. Because bentonite forms clumps when wet, the clumps can easily be removed and disposed of. The remainder of the unused material stays intact and can continue to be used. Clumping cat box litters will last longer with less frequency of changing. For many years bentonite has been used as a binder in the feed pelletizing industry. Small amounts of bentonite can be added to feed products to insure tougher, more durable pellets. Ben- tonite has also proved helpful in sealing freshwater ponds, irrigation ditches, reservoirs, sewage and industrial water lagoons, and in grouting permeable ground. In addition, it has been used in detergents, fungicides, sprays, cleansers, polishes, ceramic, paper, used as a base for cosmetics and medicines, and applications where its unique bonding, suspending or gallant properties are required.

Figure 39. Side-by-Side cat litter Comparison.

6.3 Palygorskite – Sepiolite

Palygorskite is a term that is synonymous with attapulgite. The term attapulgite is largely used industrially even though the international mineral nomenclature committee ruled that palygorskite was first used and therefore is the preferred term. Both palygorskite and sepiolite are hydrated magnesium aluminium silicates. Sepiolite has higher magnesium content than palygorskite and has a slightly larger unit cell size. Both of these minerals are thin elongate chain type structures. When dispersed in water these elongate crystals are inert and non-swelling and form a random lattice capable of trapping liquid and providing excellent thickening, suspending, and gelling properties. These clays do not flocculate with electrolytes and are stable at high tempera-

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tures, which make them uniquely applicable for many uses. The term fuller=s earth is a term used for highly absorbent and natural bleaching clays. Thus the term includes both attapulgite and calcium montmorillonite, so there is a definite overlap in the use of the term fuller=s earth, with both attapulgite and calcium bentonite.

Palygorskite (attapulgite) and sepiolite deposits are relatively rare in comparison with the other industrial clays. A major occurrence of palygorskite is in the south eastern USA in southern Georgia and northern Florida, where major deposits of Palygorskite (attapulgite) occur. The major occurrence of sepiolite is near Madrid in Spain. Two other large occurrences of palygorskite are in Senegal near Theis, ~100 km east of Dakar and in China in Anhui Province, ~120 km north- west of Nanjing. There are also occurrences of sepiolite in Turkey.

The applications of palygorskite and sepiolite are many (Table 15). The major uses of palygorskite and sepiolite, are in drilling muds, paints, liquid detergents, adhesives, car polish, flex- ographic inks, cosmetics, floor absorbents, potting mixes, oil-spill clean up material, carriers for fertilizers, pesticides, or hazardous chemicals, decolorize various mineral, vegetable and animal oils, as a receptor coating on carbonless copy paper, and for pet litter. Because these clays are relatively unaffected by electrolytes their viscosity is retained whereas bentonites flocculate and lose their high viscosity. Both palygorskite and sepiolite are used as a binder for pelletized animal feed. Certain studies have reported increased feed efficiency and improved digestive hy- giene. Another use is as an additive in cement where because of its elongate shape and absor- bency it strengthens the resulting concrete. The prices per ton range from $90 to as much as $800 for very fine highly refined material.

Palygorskite and sepiolite are surface mined in open pits, similar to bentonite. The processing involves crushing, drying, pulverization, classifying, bagging and loading. Palygorskite and sepiolite are both elongate minerals, which have caused some health organizations to concern because of the problems with asbestos, which is an elongate mineral. However, numerous tests have shown that neither of these minerals are carcinogens. Again these clays are environmentally safe as long as dust abatement procedures are taken. The open pits are reclaimed back to their original conformation so that crops or trees can be grown and harvested.

Table 15. Industrial uses of palygorskite (attapulgite) and sepiolite. Drilling fluids Adhesives Cat box absorbents Paint Pharmaceuticals Suspension fertilizer Paper Catalyst supports Agricultural carrier Ceramics Animal feed Industrial floor absorbents Asphalt emulsions Petroleum refining Mineral and vegetable oil refining Cosmetics Anti-caking agent Tape-joint compounds Sealants Reinforcing fillers Environmental absorbent

Palygorskite and sepiolite, particularly those classified as gel clays, are traded in a global market. Absorbent grades are produced and used in regional and local markets. No large multinational company is involved in the production of these clays. Palygorskite produced in Senegal, Turkey, China, and Somalia contribute a livelihood for many people who are involved in the mining, production, transportation, and marketing of these clays.

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Palygorskite and sepiolite are excellent absorbents and are essential for use as carriers for fertilizer, insecticides, and pesticides in agricultural applications when blended with bentonite; the material is superior barrier clay for use in landfills and toxic waste repositories.

6.4 Clay application for the future The future will see further growth in the use of clays in the applications they now serve and will see new utilization, particularly in industrial and environmental applications. The growth of the current industries will create expanded markets for paper, ceramics, plastics, etc. This will cause a demand for exploration for new deposits and improved processing to upgrade marginal quality deposits so that marketable products can be made.

Processing techniques will be improved and new equipment will be available for improving clay mineral products. In Table (16) a list of processes that are continually being improved and modified to produce higher quality and new products.

Table 16. Special processes applicable to clays.

Acid activation Fine pulverization Air classification Flotation Calcination Granular sizing Centrifuging High-heat drying Chemical leaching Magnetic separation Delamination Organ cladding Dewatering Oxidation Dispersion Selective flocculation Extrusion Surface treatments

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CHAPTER SEVEN

7 CLAYS IN FINLAND

The bedrock of Finland is deeply eroded and composed of metamorphic rocks or plutonic rocks, in which clay minerals occur only in fracture zones as products of hydrothermal alteration or weathering.

Hydrothermal clays are typically smectite in fractures of crystalline bedrock and they are common especially in southern Finland, which were exceptionally clay-rich (over 10% clay fraction < 2 µm) compared to the more normal Finnish sandy tills (2-3% < 2µm). Kaolin deposits of Virtasalmi, Kainuu, Kuusamo, Salla, Sodankylä, and Kittilä were developed by in situ weathering in warm and humid climate. Around these areas the Finnish tills contain kaolinite and smectites among more than usual clay minerals such as illite and chlorites. The unmetamorphosed sedimentary rocks contains clay minerals usually illite in mudstone units of Satakunta and Muhos fluvial deposits, developed in intracratonic rifts basins during Mesoproterozoic time (1600-1300 Ma) and preserved in tectonic depressions and graben structures (Kohonen and Rämö 2005). Satakunta Formation exhibits grain sizes varying from con- glomerate to siltstone and mudstone, so that the sandstone forms the bulk of these strata. The clastic components of the sedimentary rocks are, in decreasing, order quartz (45-60 vol %), microcline, and minor plagioclase (20-40 %), lithic fragments, and detrital muscovite. Among the heavy minerals used for such study, magnetite and hematite are abundant in these sediments, followed by zircon, garnet, epidote, tourmaline, apatite, monazite, and some rather exotic, in part digenetic phases (Marttila 1968). X-ray diffraction analysis revealed the presence of illites, mixed layer minerals, muscovite, and chlorite. The Neoproterozoic Hailuoto Formation contains also minor amount of kaolinite in mudstones strata. The remains of sedimentary rocks occurring in impact structures such as Lappajärvi, Sääksjärvi, Söderfjärden, Naakkima and Saarijärvi.

The clay reserves of Finland were mostly deposited from the latest stage of deglaciation until present days on the bottoms of lakes and seas. The location, quantity and quality of clay bodies were controlled by smelting of the ice and the stages of Baltic Sea.

7.1 Clay minerals in fracture zones of the Finnish bedrock

According Uusinoka (1975) the most common minerals in faults and joints are quartz, feldspar, mica (biotite, sericite, illite), chlorite, and the clay minerals with expanding lattice or swelling clays (smectites and the mixed-layer minerals smectite-illite and smectite-chlorite). Kaolinite, talc, and calcite are also fairly common, but vermiculite, epidote, and zeolites are rare or sometimes found. According to the sampled material, smectite-bearing clay is present in at least about 40 per cent of the clayey veins and weathering zones in the crystalline bedrock of Finland. The formation and occurrence of smectite as a secondary mineral depends on the intensity of leaching rather than on the petrographic character of the bedrock (Fig. 40).

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Figure 40. Smectite-bearing clays in Finland (Uusinoka 1975).

Hydrothermal alteration has been found from fracture zones of Olkiluoto. Blomqvist et al. (1992) have studied the fracture infillings zone at Olkiluoto. Based on mineral association fracture infillings were classified into five groups according to decreasing temperature and age. Two oldest groups are hydrothermal (T <300°C). The first one is associated with the emplacement of the Laitila rapakivi batholith. The hydrothermal fluids from the batholith caused alteration of the mica gneiss, including silicification and formation of albite and muscovite. The fracture infill-

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ings are characterised by muscovite-greisens fractures, silicified microbreccias, albite veins and quartz veins. The younger hydrothermal group is characterised by clay minerals, especially illite, crystallised on chlorite shear planes and fractures. Other typical fracture infillings include pyr- rhotite, baryte, laumontite-leonhardite, analsime, adular and fluorite. On the basis of three illite samples dated at 1031 Ma, 1353 Ma and 1365 Ma, the group is presumably associated with hydrothermal activity before or during the Postjotnian magmatism, represented by olivine dia- base dykes. This hydrothermal activity was concentrated along fracture zones.

7.2 Kaolin occurrences in Finland

Kaolin clays are widely used as a pigment by the paper industry both as filler and a coating. For that reason GTK has been considerate kaolin exploration is one of the most interest especially for Finnish paper industry and every year increasing amounts of kaolins are imported to our country (Fig. 41).

CONSUMPTION OF PAPER PIGMENTS IN FINLAND, 1990-2006

3500000

3000000

gypsum 2500000 talc 2000000 chalk PCC 1500000 GCC

(dry metrictonnes) kaolin 1000000

500000

0 1990 1991 1992 1993 1994 1995 1996 1997 1998 1999 2000 2001 2002 2003 2004 2005 2006

(year)

Figure 41. Consumption of paper pigments is increasing in Finland.

7.2.1 Virtasalmi kaolins

The first indications of kaolin at Virtasalmi were discovered in existing drill cores taken during the preliminary phase of kaolin investigation in 1986 by the Geological Survey of Finland (GTK). The actual kaolin project continued 1986-1992. Six main and three minor kaolin deposits were discovered at Virtasalmi and two minor deposits in Joroinen municipality, Mikkeli province, southeastern Finland (Fig. 42). The deposits at Virtasalmi are Litmanen, Eteläkylä, Vuori- joki, Ukonkangas, Montila, Niittylampi, Kahdeksaisiensuo, Hyväjärvi and Montilanlampi and they are located in a NW-trending, 20 km long and 5 km wide zone (the Virtasalmi kaolin area). Two small deposits near each others, namely Tervajoensuo and Eteläperä in Joroinen, are the only deposits discovered so far outside the Virtasalmi area. The nearest paper mill, which con- sumes kaolin for coated paper, is at Varkaus, located 60 km NE of the Litmanen deposit. Kaolin

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deposits were found by gravity and electromagnetic methods and drilling. The volume and quality of the deposits were assessed by drilling and laboratory tests (Fig. 43)

The Virtasalmi area of southeastern Finland belongs to the Svecofennian domain on the SW-side of the Raahe-Ladoga Zone. The deposits occur in valleys, 100-110 m above sea level, between drumlin ridges in the middle of an undulating plain, and usually belong to the assemblage of quartz-feldspar gneiss, amphibolites and carbonate rock. The occurrence of the deposits is also controlled by a NW-SE-trending fracture zone. The deposits are lenticular in shape, generally less than a few hundred metres wide and from a half to two kilometres long (Fig. 44). The thickness of kaolin is usually 20-40 m; in some places the thickness reaches upto100 m. In a vertical profile kaolin gradually changes via partially altered rock, into fresh parent rock. There is a gradual transition zone between kaolin and its parent rock and the kaolin contains well-preserved primary features seen in the parent rock. The transition zone and the primary features show the residual origin of kaolin. White kaolin is generally derived from quartz-feldspar gneiss or tonali- ties and coloured kaolin from amphibolites or mica gneiss, which contain more mafic minerals than the former rocks.

Figure 42. The relief of the Virtasalmi kaolin area and surrounding areas (on the right); the gravity map showing location of the Virtasalmi kaolin deposits in gravity lows (on the left), (after Sarapää1996).

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Figure 43. Aeroborne magnetic map and geological map of Virtasalmi showing the kaolin deposits and Iso-Naakkima sedimentary deposit (modified after Pekkarinen and Hyvärinen (1984) and Vorma (1971)).

The kaolin derived from quartz-feldspar gneiss is mostly white and composed of kaolinite and quartz. Kaolin derived from mica gneiss is sometime white, grey or coloured and the typical mineral assemblages are kaolinite-quartz, kaolinite-quartz-mica-(pyrite-graphite) and kaolinite- quartz-goethite (rarely haematite) respectively. Kaolin developed on amphibolites varies in colour depending on mineralogical composition; kaolinite, kaolinite-goethite and kaolinite-smectite in the less altered basal part. Kaolin on tonali- ties is mostly white and characterized by the assemblage of kaolinite-quartz. Small amounts of ilmenite, anatase and pyrite are often present in all kaolin types. The mineralogical composition is typical of a residual kaolin deposit.

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Figure 44. Litmanen kaolin deposit (Sarapää 1996)

White kaolin contains 40-75% kaolinite, 20-30% quartz, <5% feldspar and mica, and addition of hematite and goethite causes kaolin colors to change from gray to pink. Ilmenite and graphite are other staining components in both kaolin types change colour to dark gray. The grain size fractions of <20 μm and <2 μm are almost entirely composed of kaolinite with rare traces of quartz, mica (illite) and smectite. Particle size distribution shows that the proportion of the <2 μm fraction is 30% and that of the <20 μm fraction 60%. The natural brightness of white kaolin in the <20 μm fraction is on average 70% (60-85%) see Table (17).

Table 17. Brightness, yellowness, mineral and chemical composition of Litmanen kaolin (after Sarapää, 1996).

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After magnetic and chemical refining white kaolin reaches the brightness of paper coating grade (> 85%), or at least of filler grade (> 76%). The viscosity of processed kaolin varies widely, but the reasons are not fully understood. The abrasiveness fulfils the requirements of paper kaolin. There are an estimated 18 million tons of probable white kaolin resources, suitable for paper manufacturing, in the deposits at Litmanen, Vuorijoki, Ukonkangas, Eteläkylä, Kahdeksaisien- suo and Montola-Niittylampi. Resources of coloured kaolin, which could be used in ceramics, are 16 million tons (Table 18).

Table 18. Mineral and chemical composition of the most common paper kaolins used in Finland compared with the average values of refined kaolins from Litmanen. The paper kaolins are random samples from the Kotka plant of ECC International.

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SPS SUPER C DINKIE GEORGIA OPACITEX LITMANEN

Brightness 85 84 83 85 79 88 Kaolinite 95 95 95 100 amorf 100 Quartz tr tr tr 0 0 0 K-feldspar tr tr tr 0 0 0 Plagioclase 0 tr 0 0 0 0 Muscovite tr tr tr 0 0 0

SiO2 48.22 49.08 48.36 45.96 51.09 44.98

TiO2 0.09 0.05 0.04 0.93 2.87 0.29

Al2O3 37.24 36.13 36.44 38.67 42.83 37.70

Fe2O3 0.66 0.75 0.73 0.64 1.11 0.35 Mno 0.00 0.01 0.00 0.00 0.00 0.00 MgO 0.26 0.25 0.28 0.10 0.12 0.10 CaO 0.04 0.05 0.05 0.04 0.19 0.11

Na2O 0.10 0.13 0.14 0.18 0.23 0.02

K2O 1.53 2.58 2.23 0.11 0.23 0.25

P2O5 0.06 0.04 0.06 0.06 0.07 0.22 C <0.05 <0.05 <0.05 <0.05 <0.09 0.09 S 0.02 0.02 0.03 0.03 0.01 0.03

H2O 13.40 11.90 12.51 12.51 1.03 15.27 TOTAL 101.64 100.98 100.86 100.83 99.85 99.45

The behaviour of major and trace elements in Virtasalmi kaolins, especially the REE elements, is similar to that of recent tropical weathering crusts, but differs from hydrothermal deposits. Mass balance calculations show that Mg, Ca, Na and K have almost completely dissolved during kaolinization. The parent rock has lost silica (> 50%) and iron (> 70%). The volume reduction is 10- 30%. Based on the silica loss established in the Eteläkylä profile, the development of this kind kaolin deposit takes at least 16 million years in warm and humid climate.

The Virtasalmi kaolin deposits are thicker and richer in kaolinite than residual deposits in Cen- tral Europe and elsewhere in Fennoscandia. In a warm humid climate near the Equator, oxidation of sulphide minerals, and a high content of CO2 in the atmosphere maintained the groundwater acidity causing kaolinization of aluminium silicates. A high content of relatively unstable plagioclase and small amounts of K-feldspar and mafic minerals in quartz-feldspar gneiss were favourable for the formation of good quality and white kaolin. Fractured gneisses, steep schistosity and easily soluble carbonate rocks have offered a good drainage system for intense chemical weathering.

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The age of the Litmanen kaolin, dated by the K/Ar-method from authigenic illite, is 1180 million years. The Mesoproterozoic kaolin was presumably covered by the Neoproterozoic kaolinite- bearing sediments of the Iso-Naakkima sequence, which protected kaolin from erosion in the down faulted block. Mesozoic kaolinization had probably no influence on the kaolins buried under the sediments. Later most kaolins except those in deep fractures, were removed by erosion (Fig. 45)

Figure 45. Iso-Naakkima impact structure in gravity map and Iso-Naakkima sedimentary sequence. (Elo et al. 1996).

7.2.2 Kainuu kaolins

In the Kainuu region, northeastern Finland (Fig. 46), there is an occurrence dozens of kaolinite but most are small and only a few meters in thickness and without any economic potential (Al- Ani and Niemelä 2005). In major occurrences the kaolinite content is highest in the topmost part and decreases downwards and finally kaolinite disappears just before the change into a hard rock. It is obvious that these generally rather thin occurrences; often low in kaolinite are "roots" of thicker layers, whose topmost part was moved away by glacial erosion. The formation of

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weathered material was mainly caused by chemical weathering especially in fractured bedrock. The quality of kaolin has originally been influenced by the amount of feldspar and dark minerals of the parent rock and later on by the topographic position, because lower on the slopes the kaolins seem to be more coloured. The parent rock of white kaolin occurrences is feldspar-bearing serisitic quartzite. Accordingly the weathered material has high quartz and sericite content and low kaolinite content, mostly fewer than 10% and often it looks like micaceous sand. The biggest kaolin occurrence occurs in Kerkkä. The deposit is 1500 m long and 150-200 m wide and contains more than 10 Mt of white kaolin. Kaolin samples of Honkavaara, Kerkkä, Kurikkavaara, Pihlajavaara, Poskimäki and Törmänmäki occurrences in the Kainuu region (north eastern Finland) were investigated for their geochemistry, mineralogy and economic importance in the ceramic industry. The grain size analyses show that the studied deposits are dominated by sand and silty-sand and they clearly contain less fine fraction. XRD, SEM and XRF methodologies preformed chemical and mineralogical analyses were studied on representative samples of each occurrence. The chemical analyses show slightly high silica (SiO2=23 to 49.8%) and low alumina (Al2O3=12.8 to 36.9%) contents in <2 µm fractions. The studied occurrences exhibit high iron contents ranging between (Fe2O3=0.65 to 32.3%). The high percentage of iron is sometimes detected (especially on samples Kurikka K1 and Törmänmäki T3); they may be ascribed mainly to the presence of hematite, making this clay material unsuitable for ceramic applications. Whereas the Pihlajavaara and Kurikkavaara occurrences exhibit low iron content and high alumina content, may be making these clay occurrences suitable for ceramic allocations. Alkaline and alkaline- earth metals are fairly low (less than 1%) such as Na2O, CaO and MnO, but the K2O and MgO content are slightly high especially for the samples of Törmänmäki (>4% K2O), may attributed to the presence of muscovite and illite see Table (19).

Table 19. Normative mineralogical composition of the bulk material samples in Kainuu area. Occurrence Sample Quartz Kaolinite Chlorite Mica Hematite K-feldspar Plagioclase Talc Pihlajavaara PV1 60 30 10 Kurikkavaara KV1 15 70 15 Kerkkä KE3 90 5 5 Kurikka K1 10 60 5 25 Honkavaara H1 75 5 5 15 H4 65 5 20 5 5 H7 70 5 20 5 PoskimäkiP42555 1010 Törmänmäki T3 45 10 30 10 5 T4 50 30 5 10 5

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Figure 46. Map showing Kainuu region kaolins.

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7.2.3 Viittajänkä kaolin deposit

The Viittajänkä kaolin deposit is located in the northeastern Finland, in flat area with abundant wet swamps. Kaolin deposit its covered by 10-25 m thick glacial overburden and it is lenticular in shape, two km long, 500 m wide and according to borehole data 20-50 m thick (Lintinen and Al-Ani 2005). Geologically the Viittajänkä deposit is located in the southeastern extension of the volcanic serisitic and arcosic quartzite of the meta-sedimentary Matovaara formation (Fig. 47). The Viittajänkä kaolin is of primary or residual origin, formed in situ by weathering of silicate minerals (mainly feldspar and mica) of parent rocks in a warm and humid climate to kaolinite and other clay minerals (Fig. 48). Particle size distributions of the feed materials show that the amounts of the <2 µm fractions vary between 8.7 - 25.7 % with an average of 16 %, and that of the <20 µm fraction 27.7 - 73.0 % with an average of 49 %. According to the XRD analyses, the <2 µm fractions of the final products contain 85 – 95 % kaolinite, 5 – 10 % quartz and <5 % mica, while the main chemical constituents are SiO2 (50 %), Al2O3 (31 %), Fe2O3 (1.5 %) and TiO2 (0.4 %). Comparing the chemical composition of the investigated kaolin with the typical compositions of kaolin filler and coating, it appears that the Vitajänkä kaolin has higher contents of SiO2 and Fe2O3, but lower contents of Al2O3 and TiO2. This indicates that the desilication was not complete. The brightness values of the final fractions vary between 74 – 84 %, with an average of 79.5 %. The yellowness values vary between 3.5 – 9.6 % with an average of 4.3 %. The low brightness and high yellowness values mainly depend on the occurrence of iron oxides. The chemical and physical properties of the Viittajänkä kaolin were compared with the typical speci- fications for kaolin filler and coating grades (Table 20). Despite refinement, the Viittajänkä kaolin still contained considerable amounts of quartz and muscovite and the brightness remained below acceptable levels for kaolin pigment and only partly fulfilled the level for filler grade kaolin.

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Figure 47. Geological map showing the location of the Viittajänkä kaolin (Lintinen and Al-Ani 2005)

Figure 48. Location of drilling profiles in the gravity and cross sections of drilling profiles.

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Table 20. Mean brightness and yellowness values and respective mineralogical and chemical compositions (%) for different size fractions from the Viittajänkä kaolin deposit. N = number of analyses. Raw Kaolin <20 µm fractions <2µm fractions Composition White Colored White Colored White N=59 N=37 N=59 N=37 N=10

Brightness%   72.2 50.1 79.5 Min   60 21.7 74 Max   84.6 59.9 84.1 Yellowness   13.4 29.9 5.9

Kaolinite 30 30 66 56 92 Min 7 0 15 5 85 Max 70 80 90 95 95 Quartz 49 35 9 11 6 Feldspar 6 16 10 15 0 Muscovite 8 4 13 6 4

SiO2 76.2 69.9 51.7 52.2 52.3 Al2O3 13.5 14.7 27.9 25.6 31.4 TiO2 0.32 0.52 0.57 0.64 0.4 Fe2O3 2 4.4 2.9 4.1 1.5 MgO 1.2 3.6 1.9 3.7 0.84 CaO 0.03 0.23 0.01 0.1 0.03 Na2O 0.07 1.8 0.2 1.4  K2O 3.2 1.8 6.2 3.1 2.6

7.2.4 Other kaolins

Most kaolin deposits in Finland probably have formed in situ by weathering (Sarapää 1996). The kaolin occurrences of mid-eastern and northern Finland are generally smaller and have lower kaolinite content than the deposits of Virtasalmi. This is due to low alumina content in feldspathic quartzite, which is a typical parent rock of these kaolins. In Kainuu, mid-eastern Finland, the kaolin occurrences on hillsides are covered by a glacial over burdens only a few metres thick. The kaolin has derived from Lower Proterozoic Jatulian rocks, feldspathic sericite- quartzite and sericite schists. Kaolin is variable in colour, up to 35 m thick, and usually it contains 5-20% kaolinite. The kaolinite content is highest in the 5-10 m thick upper part of the weathering profile, but rapidly decreases downwards while the content of sericite, K-feldspar and goethite increases. The kaolin has been preserved from erosion in fractures or on the distal side of the hills in relation to glacial drift. These deposits represent a deeper section of a weathering profile than those at Virtasalmi. This is indicated by the frequent occurrence of sericite, fine quartz and K-feldspar in the kaolin. Further, the proportion of poorly kaolinized rock in relation to kaolin is greater. In northern Finland the greyish kaolin of the Sirunmaa deposit is derived from Lower Proterozoic Lapponian arkoses and arkosic quartzite. The Sirunmaa deposit resembles the deposits in Kainuu although its kaolinite content (30-50%) is clearly higher, but much lower than at Virtasalmi. The weathering reaches the depth of over 90 m, but the transition zone, which contains only a few percent kaolinites, forms one third of the total thickness. In Savukoski and Kuusamo kaolin is generally only a few m thick, variables in colour, low in kaolinite, and

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mostly derived from sericite-quartzite. Although the thickness of kaolin in Savukoski is small, the gravity measurement shows that bedrock is deeply disintegrated.

Apart from the above, small occurrences of kaolin in situ and kaolinitic weathering crusts have been found in northern Finland at Muonio, Vuotos and Kittilä. A thin weathering crust, containing kaolinite and halloysite, occurs in Ylivieska, western Finland.

7.2.5 Glacial and Postglacial clay deposits in Finland

The clay reserves are mostly concentrated in southern side of Salpausselkä (Hyyppä 1980). When the margin of a continental ice sheet retreats and then clay and silt were deposited at the front of the ice on the bottom of the lakes or seas. Smelting of continental ice and development of Baltic Sea controlled the deposition of clay and silt. In South and SW-Finland glacial and post-glacial clay deposits contains >50% clay material and they are thick. In eastern and northern Finland the clay deposits are rare and in Mid-Finland deposits are rich in silt and usually varved deposit. The mineralogy determines the usability of clay material. Typically clays contain quartz, feldspar and mica and in finest fraction illite chlorite, vermiculate and mixed clays. Kao- linite and smectite have not been found. The smelting point of Finnish clays is 1000-1200 C. Clays are mostly used for making bricks, expanded clay and cement.

Glacial clays, which were deposited in cold Baltic ice-lake and Yoldia Sea during smelting of the continental ice, they are varved and contain only small amounts of salt, sulphur and organic matter. So they are more useful for brick making than postglacial clays of the Anculus Lake and the Litroina Sea which accumulated in more salty and warm conditions.

7.2.6 Clay minerals in tills

Pekka Pulkkinen (2004) studied the mineralogy and geochemistry of the fine and clay fractions of till in different moraine types and in different bedrock areas in northern Finland.

In the clay fraction of till quartz, plagioclase, microcline and amphibole are the primary minerals the clay minerals proper include vermiculite, chlorite, illite, and swelling-lattice vermiculite and mixed-layer clay minerals. Kaolinite occurs most abundantly in the clay fractions of till in the Kittilä, Jerisjärvi, Kaaresuvanto and Pulju areas. Kaolinite and dioctahedral illite are evidence of the mixing of the weathered bedrock material into the till matrix. In the fine fraction of till most abundant minerals are primary minerals and clay minerals are in a minor role.

In the clay fraction of till the content of primary minerals are at higher and secondary minerals are at lower level in the Granitic and Achaean gneiss areas than in the Greenstone Belt, Svecoka- relian schists and gneiss and Granulite areas. Amphibole, microcline and plagioclase occur in very low amounts or are totally destroyed by chemical weathering in the clay fraction of the till in the Kittilä area. The mineral composition of fine and clay fractions in the tills of northern Inari gives an indication that there occur much more mafic volcanites than is known today. The mineralogical compositions of fine fraction of the tills correlates quite well with the underlying bedrock in all study areas, but clay fraction does not.

Geochemical results are in accordance with the mineralogical composition of both fractions. In the fine fraction of the till Si, Ca and Na contents are higher than in the clay fraction. Clay frac-

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tion is enriched in Al, Fe, Mg, K and trace elements as compared to the fine fraction. Pekka Pulkkinen study points out that the distribution of chemical elements in the clay fraction of the till does not correlate with the composition of the underlying bedrock, but fine fraction does so with a few exceptions. The chemical composition of till in Kaarensuvanto and Inari does not fully correspond to the composition of the underlying bedrock as known today. In northern Inari and Kittilä the results give an indication that there are more mafic vulcanite and/or sulfides min- eralization occurring in these areas than is known at the present time.

The most important factors controlling the mineralogical and geochemical composition of the fine and clay fractions of the tills in northern Finland are the composition of the bedrock and the possible occurrence of an old weathering crust. The final grain size composition of the tills and consequently the quantitative proportions of the different minerals are often related to the last glacial quarrying and sorting processes; therefore the mineralogical composition of the tills is to a certain extent bound also to the respective moraine type.

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