GEOLOGIAN TUTKUSKESKUS M19/3232/2008/41 Espoo
30.6.2008
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 FINLAND Authors
DOCUMENTATION PAGE Date / Rec. no. 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
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94
english
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M19/3232/2008/41 Confidentiality
Contents
Documentation page
CHAPTER ONE
1
1
1 1 1 1 2 4
INTRODUCTION TO CLAY MINERALOGY 1.1 Course Objectives 1.1.1 Preface 1.1.2 Clay 1.2 Identity of clay minerals 1.3 Occurrences of clay
CHAPTER TWO 2
STRUCTURE AND COMPOSITION OF THE CLAY MINERALS 2.1 Coordination Polyhedral 2.1.1 Tetrahedron 2.1.2 Octahedron 2.1.3 Layer types 2.2 Clay mineral Classification 2.2.1 Allophane 2.2.2 Kaolin Minerals 2.2.3 Smectite Minerals 2.2.4 Vermiculite 2.2.5 Illite minerals 2.2.6 Chlorite Minerals 2.2.7 Palygorskite and sepiolite 2.2.8 Mixed-layer clay minerals
CHAPTER THREE 3
PHYSICAL AND GEOCHEMICAL PROPERTIES OF CLAYS AND CLAY MINERALS 3.1 The main properties of particular clay minerals 3.1.1 Kaolin 3.1.2 Smectite (Bentonite) 3.1.3 Illite 3.1.4 Other clays 3.2 The Geochemistry of clay minerals 3.2.1 Equilibrium adsorption and Ion exchange 3.2.2 Surface charge properties 3.3 Measurement of Cation Exchange Capacity 3.4 Swelling Properties of smectite
CHAPTER FOUR
5 5 5 6 7 8 10 12 13 16 18 21 23 25 28 29 29 29 29 30 31 31 32 32 34 37 39 40
4
METHODS OF CLAY MINERALS INVESTIGATION 4.1 X-Ray Diffraction Experiment 4.2 Clay preparation for XRD 4.2.1 Procedure of Separation clay fraction from bulk sediments 4.2.2 Glycolation 4.2.3 Heating 4.2.4 Differentiation between Clays 4.3 Semi-quantitative analysis of clay minerals 4.4 Infrared spectroscopy 4.5 The Scanning Electron Microscope (SEM) 4.6 Transmission electron microscopy (TEM)
40 40 41 41 43 43 43 45 47 49 50
CHAPTER FIVE
51
5
51 51 52 53
FORMATION AND ALTERATION OF CLAY MATERIALS 5.1 Introduction 5.2 The clay cycle 5.3 Geological origin of clay minerals
CHAPTER SIX
57
6
57 57 60 62 62 64 64 65 67
INDUSTRIAL USES OF CLAYS 6.1 Kaolin 6.1.1 Paper industry 6.1.2 Other Applications 6.2 Smectite (Bentonite) 6.2.1 Bentonite barriers in sealing nuclear waste 6.2.2 Other uses of Smectite (bentonite) 6.3 Palygorskite – Sepiolite 6.4 Clay application for the future
CHAPTER SEVEN
68
7
CLAYS IN FINLAND 7.1 Clay minerals in fracture zones of the Finnish bedrock 7.2 Kaolin occurrences in Finland 7.2.1 Virtasalmi kaolins 7.2.2 Kainuu kaolins 7.2.3 Viittajänkä kaolin deposit 7.2.4 Other kaolins 7.2.5 Glacial and Postglacial clay deposits in Finland 7.2.6 Clay minerals in tills
68 68 70 70 76 79 811 81 82
8
REFERENCES
844
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CHAPTER ONE 1
INTRODUCTION TO CLAY MINERALOGY
1.1 1.1.1
Course Objectives 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 Standardization) 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. Because 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 basements. 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. Sodium 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. Furthermore, 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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4
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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7
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
b 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. Compositional 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 Nomenclature 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 magnitude 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. Consequently, 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 consists 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 designate 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 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 recrystallization 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
18
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 octahedrally-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 salvation 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 Reynolds, 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 Mgfilled, 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).
24
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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 clinochlore for Mg-dominant [end member = (Mg5Al)(Si3Al)O10(OH)8], chamosite for Fe2+dominant [end member = (Fe52+Al)(Si3Al)O10(OH)8], nimite for Ni-dominant [end member = (Ni5Al)(Si3Al)O10(OH)8], and pennantite for Mn2+-dominant [end member = (Mn52+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)
8
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, greyishwhite, or slightly coloured. It is made up of tiny, thin, pseudohexagonal, flexible sheets of triclinic 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 disintegrates the crystal lattice and produces a variety of phases, such as mullite, cristobalite, and cordierite, 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
31
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 controversial. 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 investigators, 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 3.2.1
The Geochemistry of clay minerals Equilibrium adsorption and Ion exchange
The typically small grain size ( 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 Gibbsite Hematite Goethite Na-feldspar Kaolinite Montmorillonite Quartz
pHZPC 10 4.2 - 6.9 5.9 - 6.7 6.8 2 - 4.6 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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V = 2/9(ρg - ρw) g r2/η Where: V = the settling velocity ρg = density of the mineral grain (2.6 - 2.8 g/cm3 for clay minerals) ρw = density of water (1g/cm3) 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 Nahexametaphosphate, it obviously attacked carbonate minerals (calcite peak was disappeared). Silt fraction (2-63µm) and clay fraction (
