Schulze Chap1

8/8/2019 Schulze Chap1

1/33

Introductory Concepts Final Draft, Revised 1/18/00 Page 1

An Introduction to Soil Mineralogy

Darrell G. Schulze, Purdue University, West Lafayette

Table of Contents

I. CHEMICAL AND STRUCTURAL CLASSIFICATION OF

MINERALS ................................................................................................... 2

A. Composition of the Earths Crust .....................................................2

B. Definition of a Mineral ..................................................................... 2C. Mineral Classification.......................................................................2

1. Halide, Sulfate, and Carbonate Minerals .....................................3

2. Sulfides.........................................................................................3

3. Oxides, Hydroxides and Oxyhydroxides......................................3

4. Silicates ........................................................................................4

II. PHYLLOSILICATE MINERALS IN SOILS....................................... 4

A. Basic Structural Concepts.................................................................4

1. Closest Packing of Spheres ..........................................................4

2. Tetrahedra and Octahedra............................................................4

3. Representing Tetrahedra and Octahedra......................................5

4. Ionic Radii and Radius Ratios......................................................5

5. Tetrahedral and Octahedral Sheets...............................................6

B. Phyllosilicate Minerals Common in Soils.........................................7

1. The 1:1 Type Minerals................................................................. 7

2. The 2:1 Type Minerals................................................................. 7

C. Structural Details of Phyllosilicates..................................................9III. OTHER ALUMINOSILICATE MINERALS COMMON IN SOILS 10

A. Zeolites ...........................................................................................10

B. Allophane and Imogolite ................................................................10

IV. SOME CRYSTALLOGRAPHIC CONCEPTS..................................11

A. Periodicity in Crystals.....................................................................11

B. The Unit Cell ..................................................................................11

C. Miller Indices..................................................................................12

D. X-ray Diffraction ............................................................................12

1. Bragg's Law................................................................................12

2. The Scherrer Equation................................................................13

V. SUMMARY........................................................................................13

VI. SUGGESTED READING ..................................................................13

VII. QUESTIONS AND EXERCISES ......................................................13

VIII. ACKNOWLEDGMENTS................................................................... 14

IX. REFERENCES................................................................................... 14

X. LIST OF FIGURES............................................................................. 14

XI. LIST OF TABLES.............................................................................. 15


2/33


An Introduction to Soil Mineralogy

Darrell G. Schulze, Purdue University, West Lafayette, Indiana

Minerals make up about one-half of the volume of most soils. Theyprovide physical support for plants and create the water- and air-filled poresthat make plant growth possible. Mineral weathering releases plant nutrientsthat are retained by other minerals through adsorption, cation exchange, andprecipitation. Minerals are indicators of the amount of weathering that hastaken place and the presence or absence of particular minerals gives clues tohow soils formed. The physical and chemical characteristics of soil mineralsare important considerations in planning, constructing, and maintainingbuildings, roads, and airports. Soil minerals can adsorb many organic andinorganic environmental pollutants, promoting their degradation to nontoxicforms, attenuating their movement through the soil, or preventing theiruptake by plants and their introduction into the food chain. Some mineralsare themselves pollutants and can cause serious environmental problemswhen they are exposed to weathering at the soil surface by mans activities.An understanding of soil mineralogy is central to understanding virtually allfacets of mans use and misuse of soils and is often the key to solvingspecific environmental problems.

This chapter develops a core of concepts and terminology needed forunderstanding soil minerals. The chemical composition of the earth's crust isdiscussed first to show that the most abundant elements in the crust are, notsurprisingly, the ones most likely to be encountered in soil minerals. Thenthe chemical and structural classification of minerals is discussed and themajor minerals represented in soils are mentioned. The phyllosilicateminerals are covered separately because of their major role in soils. Basicstructural concepts common to all minerals are covered at this point toprovide the background for the discussion of the phyllosilicate structures.The overall structural theme of the phyllosilicates is then presented alongwith a very brief summary of their most important properties. The chapterconcludes with a treatment of some crystallographic and x-ray diffractionconcepts used throughout the book.

I. CHEMICAL AND STRUCTURAL CLASSIFICATION OFMINERALS

A. Composition of the Earths Crust

Most of the weight and volume of the earth's crust is made up by only afew elements (Table 1-1). Oxygen and Si make up most of the weight, whileoxygen alone accounts for more than 90% of the volume. In general, the

larger O atoms are in an approximately close-packed arrangement heldtogether by smaller metal atoms in the interstitial space. Most of theelements in the crust and in soils occur in minerals. Thus, the elements listedin Table 1-1 are major constituents of the minerals discussed in this book.

B. Definition of a Mineral

Klein and Hurlbut (1993) define a mineral as follows: "A mineral is anaturally occurring homogeneous solid with a definite (but not generallyfixed) chemical composition and a highly ordered atomic arrangement. It isusually formed by inorganic processes." Both chemical composition andcrystal structure (ordered atomic arrangement) are important parts of thisdefinition. Neither alone is sufficient to explain the properties of minerals.Minerals with similar chemical composition but different crystal structures or, conversely, minerals with similar crystal structures but differentchemical compositions can be quite different from one another despitetheir chemical or structural similarities.

C. Mineral Classification

Although different classification schemes could be used, mineralogistshave determined that first separating minerals into groups based on theirchemical composition gives classes with the greatest similarities in manyother properties. Thus, minerals are first divided into classes depending uponthe dominant anion or anionic group. The classes include: (1) nativeelements, (2) sulfides, (3) sulfosalts, (4) oxides and hydroxides, (5) halides,(6) carbonates, (7) nitrates, (8) borates, (9) phosphates, (10) sulfates,(11) tungstates, and (12) silicates (Klein and Hurlbut, 1993). These classesare then subdivided based on chemical and structural similarities. This samegeneral classification is followed in this book, but with some exceptions.

Mineral classes such as the native elements, sulfosalts, nitrates, borates, andtungstates that occur only rarely in soils, are covered only briefly or not atall. Commonly occurring mineral classes, particularly the silicates, arecovered in detail.

Soil minerals are also referred to as either primary or secondaryminerals. Primary minerals form at elevated temperatures. They are usuallyderived from igneous or metamorphic rocks, but they can be inherited fromsedimentary rocks as well. Secondary minerals are formed by low-temperature reactions and are either inherited from sedimentary rocks orformed in soils by weathering (Jackson, 1964). The separation of mineralsinto primary and secondary mineral classes is not mutually exclusive andsome minerals can occur in both. The concept is useful however, andappears widely in the soil science literature. The major mineral classesrepresented in soils are discussed below.


3/33


1. Halide, Sulfate, and Carbonate Minerals

The major soil minerals of this group are halite (NaCl), gypsum(CaSO42H2O), calcite (CaCO3), and dolomite [CaMg(CO3)2] (Table 1-2).This group is characterized by minerals with relatively simple structures.The halite structure is one of the simplest of all minerals. It consists ofalternating Na

+and Cl

-ions arranged in cubic closest packing (Fig. 1-1). The

other minerals in this group have similar structures with cations such asCa

2+, Mg

2+, or Fe

2+alternating with anions such as S2

2-, SO4

2-, or CO3

2-. The

bonds between the cations and anions are predominantly ionic. Theseminerals are among the most soluble and the softest of all soil minerals andthey are easily broken down by physical and chemical weathering processes.They occur mainly in soils of arid regions or in youthful soils in more humidregions, where weathering has been minimal.

Halite is the most soluble of this group and accumulates in only the mostarid environments. It is one of the minerals present in the salic horizon ofAridisols. Gypsum is about 100 times less soluble than halite, but it too isabundant only in arid regions. Gypsum is a major mineral in the gypsichorizon of Aridisols. Halite and gypsum could also occur in soils that havebeen contaminated by salt water, or by soluble salts leaching from industrialstock piles or waste piles.

Calcite and dolomite are common carbonate minerals that occur in avariety of soils. These minerals precipitate in the soil profile in arid andsemiarid climates (aridic and some ustic soil moisture regimes). Calcic and petrocalcic horizons form if the accumulation is great enough. Carbonateminerals in many soils are inherited from limestone or other calcareousparent materials. Carbonates are usually stable and can be found throughoutthe soil profile under aridic and ustic soil moisture regimes (arid tosubhumid climates), but are leached from the soil profile and are generallypresent only in the C horizons under the udic soil moisture regime (humid

climates). Calcite and dolomite can be introduced into soils originally free ofthese minerals via the limestone aggregate used for road construction insome areas.

2. Sulfides

Pyrite, FeS2, the most common mineral in this group, does not occurextensively in soils, but when it is present it causes some unique problems.Pyrite precipitates in soils on wet tidal flats and river deltas of some coastalareas and also occurs in some geologic formations originally deposited insimilar environments. Thus, pyrite often occurs in close association withcoal. Pyrite is unstable under oxidizing conditions and weathers quicklywhen pyritic soils are drained or when mining leaves pyritic material on thesurface. The weathering products include the minerals jarosite,KFe3(SO4)2(OH)6, and gypsum and sulfuric acid, H2SO4. The large amount

of acidity causes problems in utilizing pyrite-containing soils and inrevegetating mined areas and is a serious and costly environmental problem.

3. Oxides, Hydroxides and Oxyhydroxides

Primary minerals break down during weathering and release cations andanions that recombine to form other more stable minerals. Several elements,in particular Al, Fe, and Mn, form oxide, hydroxide, or oxyhydroxideminerals that are stable in the soil weathering environment. Representativemineral species are given in Table 1-2.

a.AluminumGibbsite, [Al(OH)3] is the most common Al hydroxide mineral in soils.

It is generally associated with the latter stages of weathering when leachingof silica has progressed to the point that phyllosilicate minerals no longerform. Gibbsite is common in highly weathered Oxisols of tropical regions. Ithas a very low cation exchange capacity and contributes to the low nativefertility of most Oxisols. Gibbsite is also commonly found at the weatheringinterface between igneous rock and saprolite and in Andisols formed fromvolcanic ash. Despite the high solution silica content in these environments,gibbsite precipitates as a metastable phase that eventually dissolves andgives way to more silica-rich minerals (Sposito, 1989, p. 101).

b.IronIron oxide minerals form from Fe released from primary minerals. Iron

oxides are strong pigments and small amounts of these minerals account formost of the brown and red colors of soils. Goethite (FeOOH) is the mostcommon mineral of this group. It accounts for the brownish to yellowishcolor of many soils, although it may be present in only small quantities.Hematite (Fe2O3) is only slightly less common than goethite and usually

occurs in association with it. Hematite is usually bright red and isresponsible for the red color of many soils. Goethite and hematite are stableminerals in an oxidizing environment. Large amounts of these two mineralsin well-aerated soils, usually in association with gibbsite and kaolinite,usually indicate an advanced stage of weathering. In soils that are saturatedwith water for at least some time during the year, the very insoluble Fe

3+in

goethite, hematite, and other iron oxides can be reduced to very soluble Fe2+

.The Fe

2+can easily move with the soil water to other parts of the soil profile,

or even to other associated soils in the landscape, where it can then reoxidizeto Fe

3+and reprecipitate as goethite, lepidocrocite, or ferrihydrite when

oxidizing conditions return. Repeated cycles of oxidation and reduction giverise to mottles and nodules that reflect an inhomogeneous distribution ofiron oxide minerals within a soil. Soil scientists use these inhomogeneous ormottled color patterns to estimate the depth to a soil water table and thelength of time the soil is saturated during the year. Thus, when delineatingwetlands, determining the best location for a septic system leach field, or


4/33


mapping soils, soil scientists rely on predictable soil color patterns that resultfrom iron oxide mineralogy and distribution.

c.ManganeseManganese oxides and hydroxides (Table 1-2) are commonly found in

soils as brown or black nodules or as thin coatings on the faces of soilstructural units. They are often associated with Fe oxides. Manganese occursfrequently as birnessite or lithiophorite in soils. Manganese can be oxidizedand reduced in the soil environment similar to iron. Thus, theinhomogeneous distribution of Mn into nodules is an indicator of reductionand oxidation as a result of periodic water saturation.

d. TitaniumRutile and ilmenite (Table 1-2) occur in soils mainly as primary

minerals inherited from igneous rocks. Anatase is less common and isgenerally considered a secondary mineral. Although frequently found insoils, these minerals do not occur in sufficient quantity to impact soilphysical or chemical properties.

4. Silicates

The silicate mineral class is an extremely large and important group ofminerals. Nearly 40% of the common minerals are silicates, as are mostminerals in igneous rocks. Silicates constitute well over 90% of the earth'scrust (Klein and Hurlbut, 1993) and comprise the bulk of most soils as well.Silicates occur as both primary minerals inherited from igneous ormetamorphic rocks and as secondary minerals formed from the weatheringproducts of primary minerals.

As explained in more detail below, the fundamental unit of all silicatestructures is the SiO4 tetrahedron. It consists of four O

2-ions at the apices of

a regular tetrahedron coordinated to one Si4+ at the center. The individualtetrahedra are linked together by sharing O

2-ions to form more complex

structures. Several different arrangements of the SiO4 tetrahedra occur,partly accounting for the large number of silicate minerals and providing the basis for their classification. The tetrahedra may be present as singletetrahedra (nesosilicates), double tetrahedra (sorosilicates), rings(cyclosilicates), single or double chains (inosilicates), sheets (phyllosilicates)or three-dimensional frameworks (tectosilicates) (Table 1-3). In most ofthese arrangements adjacent SiO4 tetrahedra share corners, that is, they sharea common O

2-. The more common minerals from each silicate class likely to

be found in soils are given in Table 1-3.The most common minerals in igneous rocks are the olivines,

pyroxenes, amphiboles, micas, feldspars, and quartz. These primary mineralspredominate in the sand and silt size fractions of soils. The clay (


5/33


voids of the lower plane of spheres, are called tetrahedral sites because thefour O

2-ions surrounding this site form the apices of a regular tetrahedron.

The "B" sites, formed over the "B" voids of the lower plane of spheres, arecalled octahedral sites because the six O

2-ions surrounding each of these

sites form the apices of a regular octrahedron (Fig. 1-3). This illustrates that both octahedral and tetrahedral sites are a consequence of the closestpacking of two planes of spheres.

A cation in a tetrahedral site is in 4-fold (or tetrahedral) coordination because it is bonded to four O

2-ions whose centers define the apices

(corners) of a tetrahedron. Likewise, a cation in an octahedral site is in6-fold (or octahedral) coordination because it is bonded to six O

2-ions

whose centers define the apices of an octahedron. Tetrahedrally andoctahedrally coordinated cationstetrahedra and octahedra for shortarecommon structural elements in many mineral structures. The approach ofdescribing mineral structures as assemblages of polyhedra is used widelyand, once mastered, provides an efficient way to understand even complexmineral structures.

3. Representing Tetrahedra and Octahedra

Octahedra and tetrahedra are commonly represented using threedifferent types of models. Each model portrays the same concept, butdifferent representations are necessary to show various structural features.Fig. 1-3 shows two views of a tetrahedron and of an octahedron drawn inthese three different ways.

The sphere-packing model gives an impression of the space occupied bythe atoms in a structure. The sphere-packing model is particularly useful forvisualizing the shape of the external surface of a mineral, but it does notallow one to easily see the interior of the structure. The ball-and-stick modelshows the bonds within the crystal since the bonds are represented as

"sticks" connecting the balls. The tetrahedra and octahedra are more difficultto visualize, but the interior of the structure is easier to see because theatoms are drawn much smaller than in the close-packed model. Thepolyhedral models give the best impression of tetrahedra and octahedra sincethe atoms are represented only as points in space or as small spheres. Thecenters of the O

2-ions are connected by lines to form the edges of tetrahedra

and octahedra. Thus, each apex represents the position of an O2-

ion, while acation resides in the center of the tetrahedron or octahedron. In general, onlythose tetrahedra or octahedra actually occupied by cations are represented bya solid polyhedron, sites that are not occupied by cations are not representedby solid polyhedra. One should use care in interpreting polyhedral models,since it is easy to get the impression that a structure contains a large amountof open space or pores. The same structure drawn as a close-packed modelwill show that much of the space is actually occupied by the large oxygenatoms. Other ways of representing mineral structures are encountered aswell. Sometimes, polyhedral models are drawn without spheres to represent

the oxygens and hydroxyls at the corners of the polyhedra (see, for example,Fig. 1-9 later in this chapter), while other models show only sticks torepresent bonds, but no spheres to represent the atoms.

There is some ambiguity associated with each type of representation,and it is important to understand the correspondence between, andlimitations of, each of the different structural representations. Thus, it isimportant to utilize as many different learning aids as possible. Computergraphics programs that allow the viewer to rotate and manipulate structuralmodels on the computer screen offer significant advantages over staticdrawings. Physical models should be studied as well, if they are available.Far from being childs play, building your own models is one of the bestways to learn about mineral structures. Models can be built with inexpensive balls and some glue or with paper polyhdera [see Moore and Reynolds(1997) for templates].

4. Ionic Radii and Radius Ratios

Whether or not a given atom can occupy a tetrahedral or an octahedralsite depends both on the atoms charge and size. The size of atoms and ionsvaries depending on their atomic number, ionization state, and coordinationnumber (Table 1-4).

The central void within a tetrahedron is smaller than the one within anoctahedron. In other words, the largest sphere that can be placed in atetrahedral site without pushing the O

2-ions apart is smaller than the sphere

that can be placed in an octahedral site without pushing. Simple calculations(see, for example, Klein and Hurlbut, 1993) illustrate that the maximumradius of a sphere that will just fit in the tetrahedral site is 0.225 times theradii of the 4 surrounding O

2-ions. For an octahedral site, the maximum

radius is 0.414 times the radii of the 6 surrounding O2-

ions. The next largertype of site (not illustrated) is a cubic site consisting of 8 O

2-ions arranged at

the corners of a cube. For a cubic site, the maximum radius is 0.732 timesthe radii of the 8 surrounding O

2-ions. Assuming a radius of 0.140 nm for

the O2-

ions, the limiting radius for a tetrahedral site is 0.032 nm (0.140 nm 0.225 = 0.032 nm), for an octahedral site it is 0.058 nm, and for a cubicsite it is 0.102 nm. The Si

4+ion has a radius >0.032 so Mg

2+is unlikely to occur in tetrahedral

coordination, but since its radius is between 0.058 and 0.102 nm, it fitseasily in an octahedral site. The Al

3+ion, with a radius near to the limiting

radius of 0.032 nm, can fit in both tetrahedral and octahedral sites. Cationswith radii >~ 0.102 nm tend to occur in cubic, dodecahedral (12-fold) orhigher coordination sites rather than in octahedral or tetrahedral sites(Table 1-4).


6/33


5. Tetrahedral and Octahedral Sheets

Tetrahedra arranged into sheets are common to the structures of allphyllosilicates. Octahedra arranged into sheets are present in the structuresof phyllosilicates and in some hydroxide minerals as well. The structure oftetrahedral and octahedral sheets will be discussed first, then it will beshown how different combinations of tetrahedral and octahedral sheets giverise to the different clay mineral structures. Most of the diagrams in thischapter show ideal crystal structures based on the approximate closestpacking of spheres.

a. The Octahedral SheetConsider two planes of spheres representing hydroxyl (OH

-) ions in

hexagonal closest packing. (The H+ takes up very little space and the OH-

can be considered a sphere of roughly the same size as an O2-

ion). There aretwo ways of filling the octahedral sites depending on the valence of thecation. A divalent cation, such as Mg

2+, can be placed into each octahedral

site to obtain the network of octahedra illustrated in Fig. 1-4a. Thisarrangement is called trioctahedral because 3 of every 3 octahedral sites areoccupied by a cation. This gives a formula of Mg3(OH)6 or Mg(OH)2 for thewhole sheet; each Mg

2+is surrounded by 6 OH

-'s, but since each OH

-is

shared equally among 3 different Mg2+

atoms, each OH-contributes only 1/3

of its negative charge to each Mg2+

, giving the formula Mg(OH)2. The sheetis electrically neutral because the charge is balanced within the sheet.

The other possibility is to place a trivalent cation, such as Al3+

, into theoctahedral sites. To preserve charge neutrality, only 2 trivalent cations areneeded compared to 3 divalent cations, giving the formula Al 2(OH)6 orAl(OH)3. Thus, only 2 of every 3 possible octahedral sites are filled. Thisarrangement is called dioctahedral and is illustrated in Fig. 1-4b. The unitformula for the dioctahedral sheet can be deduced by following the same

reasoning used above. Each Al3+ is surrounded by six OH-'s, but since eachOH

-is shared equally between 2 different Al

3+ions (not 3 as for Mg

2+), each

OH-

contributes 1/2 of its negative charge to each Al3+

, giving the formulaAl(OH)3. Again the charge is balanced and the sheet is electrically neutral.

Note how the trioctahedral and dioctahedral sheets appear in thedifferent types of representations. The sphere-packing model shows that both the trioctahedral and dioctahedral sheets contain the same densepacking of OH

-'s. The pattern of unoccupied octahedral sites is more evident

in the polyhedral model, where unoccupied octahedra appear as open spaces,and in the ball-and-stick model.

Octahedral sheets, stacked one on top of the other and held together byweak residual bonds, make up the structures of gibbsite (Fig. 1-7) andbrucite [Mg(OH)2]. Gibbsite and brucite differ in that gibbsite contains Al inthe octahedral sites and is dioctahedral, while brucite contains Mg in theoctahedral sites and is trioctahedral. The structures of gibbsite and bruciteare the simplest in a series of structures containing octahedral sheets.

The discussion above followed a sphere-packing approach in which weshowed that the structure of the octahedral sheets could be derived assumingthat the large oxygen atoms were hard spheres of identical size closelypacked in three dimensions. Thus, we assumed that the oxygen atoms wereour basic structural units and that the cations merely served to balance thenegative charge of the oxygens. This approach works for some of the oxideand hydroxide minerals, but in many mineral structures the oxygen atomsare not as closely packed as strict geometric packing would dictate. A moregeneral way of describing mineral structures is called the polyhedralapproach. In the polyhedral approach, the coordination polyhedra, primarilyoctahedra and tetrahedra in the phyllosilicates, are considered the basicstructural building blocks. Thus, we can consider the octahedral sheetsillustrated in Fig. 1-4 as an assemblage of octahedra in which adjacentoctahedra share two oxygens with one another. In other words, the octahedrashare edges with one another. In general, coordination polyhedra can sharecorners (one shared oxygen), edges (two shared oxygens), or faces (three ormore shared oxygens) with adjacent polyhedra. The octahedral sheet couldbe described using either the sphere packing or polyhedral approaches. Thetetrahedral sheet described below can only be describe using the polyhedralapproach because the oxygens in the tetrahedral sheet are not close-packedas they are in the octahedral sheet.

b. The Tetrahedral SheetThe tetrahedral sheet consists of SiO4 tetrahedra arranged such that

three O2-

ions of each tetrahedron are shared with the three nearest neighbortetrahedra (Fig. 1-5). These shared O

2-ions are all in the same plane and are

referred to as basal oxygens. Note that two adjacent tetrahedra share onlyone O

2-between them. (The tetrahedra share apices or corners.) The fourth

O2-

ion of each tetrahedron is not shared with another SiO 4 tetrahedron andis free to bond to other polyhedral elements. These unshared O

2-ions are

referred to as apical oxygens. Since each basal oxygen contributes a formalcharge of -1 to each Si

4+ion, the addition of H

+ions to the apical oxygens to

form hydroxyls should result in an electrically neutral tetrahedral sheet. Suchindividual tetrahedral sheets, although sometimes postulated to occur astransient weathering products in aqueous solution, do not stack to formstable mineral structures on their own. Thus, tetrahedral sheets only occur incombination with octahedral sheets as described below.

The upper half of Fig. 1-5 shows all of the apical oxygens pointing inthe same direction, namely, out of the plane of the paper towards the reader.Structures in which the apical oxygens of a single sheet all point in the samedirection are the most common, but structures also occur in which the apicaloxygens point alternately in opposite directions. Minerals containing this


7/33


8/33


giving ideal formulas of Mg3Si4O10(OH)2 for talc and Al2Si4O10(OH)2 for pyrophyllite. In each case the charge is balanced within the 2:1 layer,making it electrically neutral. Adjacent 2:1 layers are held together only byweak van der Waals forces.

Talc and pyrophyllite occur only rarely in soils, usually only when theyare inherited from low-grade metamorphic rocks. Talc and pyrophyllite areused industrially as ingredients in paints, ceramics, plastics, paper, andcosmetics. The occurrence of talc in river and estuarial sediments when thereis no geological source is a reflection of industrial activity in the watershed.

c.Micas.Mica minerals have the 2:1 layer structure described for talc and

pyrophyllite but with two important differences. First, instead of having onlySi4+ in the tetrahedral sites, one-fourth of the tetrahedral sites are occupied by Al

3+. Because of this substitution, there is an excess of one negative

charge per formula unit in the 2:1 layer. Second, this excess negative chargeis balanced by monovalent cations, commonly K

+, that occupy interlayer

sites between two 2:1 layers. This gives an ideal formula ofKAl2(AlSi3)O10(OH)2 for a mica mineral with Al in the octahedral sites.

Just as in talc and pyrophyllite, the octahedral sheet can be eitherdioctahedral (Fig. 1-7) or trioctahedral. There are several different micaspecies because Fe

2+and Fe

3+can substitute for Mg

2+and Al

3+in the

octahedral sheet and Na+

and Ca2+

can substitute for K+

in the interlayer.Mica in soils is usually inherited from the parent rock and is likely to

occur in soils derived from various igneous and metamorphic rocks, as wellas from sediments derived from them. Muscovite, biotite, and phlogopite arethe three most common mica group minerals in rocks, and consequently insoils. All three contain K in the interlayer (Table 1-3), but they differ in thecomposition of the octahedral sheet and whether they are di- or trioctahedral.

Mica in the clay fraction of soils and sediments differs somewhat fromthe macroscopic muscovite mica it most closely resembles. This clay-sizemica is often referred to as illite. Glauconite is another mica mineral that issimilar to illite, but it contains more Fe and less Al in its octahedral sheetthan illite.

Micas weather to other minerals, particularly to vermiculites andsmectites and the K

+released during weathering is an important source of K

for plants. As a rule, the dioctahedral micas, such as muscovite, are moreresistant to weathering than trioctahedral micas. Thus, muscovite tends to bethe most common mica mineral found in soils. Micas are used commerciallyin paints and cosmetics.

d. Vermiculites.Vermiculite has a 2:1 layer structure as described for mica, but instead

of having a layer charge of ~1 per formula unit and K+

in interlayer positions, vermiculite has a layer charge of 0.9 to 0.6 and contains

exchangeable cations, primarily Ca and Mg, in the interlayer (Fig. 1-7). Thehigh charge per formula unit gives vermiculites a high CEC and causes themto have a high affinity for weakly hydrated cations such as K

+, NH4

+and

Cs+. Fixation of K

+ by vermiculites can be significant in soils high in

vermiculite.Vermiculites in soils are believed to form almost exclusively from the

weathering of micas and chlorites. The weathering of micas to vermiculite(or smectite) is believed to occur by replacement of K

+in the interlayer sites

with hydrated exchangeable cations. The integrity of the 2:1 layer ispreserved but there is a reduction in the layer charge. Vermiculite does notswell as extensively as smectite and this is illustrated in Fig. 1-7 by thepresence of only two planes of water molecules surrounding the hydratedcations in the interlayer space. Some commercial uses of vermiculitesinclude horticultural potting media and thermal insulating materials.

e. Smectites.The smectite group consists of minerals with the 2:1 structure already

discussed for mica and vermiculite, but with a still lower charge per formulaweight, namely 0.6 to 0.25. As in vermiculite, the interlayer containsexchangeable cations (Fig. 1-7).

As for the micas and vermiculites, dioctahedral smectites are morecommon in soils than trioctahedral smectites. The most common smectiteminerals range in composition between the three end-members:montmorillonite, beidellite, and nontronite. All are dioctahedral, but theydiffer in the composition of the tetrahedral and octahedral sheets. Smectitesdo not fix K

+as readily as do vermiculites because smectites have a lower

layer charge, but smectites swell more extensively than vermiculite. This isillustrated in Fig. 1-7 by the larger spacing between the 2:1 layers.

Smectites are important minerals in temperate region soils because oftheir high surface area and their adsorptive properties. Smectites shrink upondrying and swell upon wetting. This shrink-swell behavior is mostpronounced in the Vertisol order and in vertic subgroups of other soil orders.The shrink-swell properties lead to cracking and shifting problems whenhouses, roads, and other structures are built on smectitic soils. Smectites areused widely in industry as catalysts, adsorbents for spills, as sealants forponds and abandoned wells, in drilling fluids for oil wells, and in liners forlandfills.

f. Chlorites.Like mica, chlorite minerals have a 2:1 layer structure with an excess of

negative charge. In contrast to mica, however, the excess charge is balancedby a positively charged interlayer hydroxide sheet (Fig. 1-7), rather than K

+.

The interlayer hydroxide sheet is an octahedral sheet as illustrated in Fig. 1-4and can be either di- or trioctahedral. Instead of being electrically neutral asin brucite or gibbsite, the hydroxide sheet has a positive charge cause by


9/33


substitution of higher valence cations for lower valence ones, for example,Mg2Al(OH)6

+. Either octahedral sheet the one that is part of the 2:1 layer

or the interlayer hydroxide sheet can be di- or trioctahedral, and cancontain Mg

2+, Fe

2+, Mn

2+, Ni

2+, Al

3+, Fe

3+, and Cr

3+, giving a large number

of different mineral species.Chlorite minerals in soils may be primary minerals inherited from either

metamorphic or igneous rocks. They may also be inherited from sedimentaryrocks such as shales, or from hydrothermally altered sediments. Chlorites arerather infrequent minerals in soils and when they do occur they generallymake up only a small amount of the soil. Chlorite weathers to formvermiculite and smectite and the ease with which they break down makesthem sensitive indicators of weathering.

g.Hydroxy-interlayered vermiculite and smectite.Hydroxy-interlayered vermiculite and smectite can be considered a solid

solution with vermiculite or smectite as one end member and chlorite as theother. Hydroxy-interlayered minerals form as Al

3+released during

weathering hydrolyzes and polymerizes to form large polycations with a postulated formula of Al6(OH)15

3+(or similar) in the interlayers of

vermiculite and smectite. These polycations balance some of the charge ofthe 2:1 layer. The combination of a 2:1 layer with hydroxy Al in theinterlayer gives a structure similar to that of chlorite (Fig. 1-7). Thus, theseminerals are also called secondary chlorites. The degree of filling of theinterlayer with hydroxy Al can vary from none to almost complete, withproperties of the clay varying accordingly. The interlayer hydroxy Al is notexchangeable, therefore it lowers the CEC of smectite or vermiculite almostlinearly as a function of the amount of Al adsorbed in the interlayer.

Interlayer hydroxy Al prevents smectite from shrinking and swelling asit normally would. In vermiculite, it reduces K

+fixation by lowering the

exchange capacity and by preventing the interlayer from collapsing aroundthe K+. The positively charged hydroxy interlayers also provide potentialsites for anion adsorption.

Hydroxy-interlayered vermiculite and smectite are most common inAlfisols and Ultisols. Within a given profile they tend to be most abundantnear the soil surface. Hydroxy-interlayered smectites produced industriallyare used as catalysts in the chemical industry.

h.Interstratification in Phyllosilicates.Phyllosilicates in soils do not always occur as discrete particles of mica,

vermiculite, smectite, chlorite, or kaolinite. For example, instead of beingmade up of a stack of identical 2:1 vermiculite layers, one physically discreteparticle may consist of a mixture of both mica and vermiculite layers instead.Such minerals are referred to as mixed-layerorinterstratified minerals.

Different types of interstratified minerals have been identified. Two-component systems include: mica-vermiculite, mica-smectite, mica-chlorite,

chlorite-vermiculite, chlorite-smectite, chlorite-swelling chlorite, andkaolinite-smectite. Three-component mixed layer systems can also occur.The sequence of layers can be either regular or random. A regularlyinterstratified mineral consisting of two types of layers denoted by A and Bcould have a sequence like ABABAB, or ABBABBABB, or any otherrepeating sequence. In a randomly interstratified mineral the sequence oflayers is random for example, ABBABAABBAAA. Randominterstratification of layer-silicates is more common in soils than regularinterstratification, though regular interstratification, especially in weatheringmicas, is not rare.

Partial removal of interlayer K from micas or of interlayer hydroxidefrom chlorite is one way that interstratified minerals can form in soils. Otherpossibilities include (i) fixation of adsorbed K

+by some vermiculite layers

to give mica-like layers, and (ii) the formation of hydroxide interlayers toproduce chlorite-like layers.

i. Palygorskite and sepiolite.Palygorskite and sepiolite are considered phyllosilicates, but are distinct

structurally from the typical 1:1 and 2:1 layer structures. Both minerals havecontinuous tetrahedral sheets, but adjacent bands of tetrahedra within onetetrahedral sheet point in two opposite directions rather than in one directionas in the 1:1 and 2:1 structures. The result is a structure that can be describedas ribbons of 2:1 layers joined at their edges as illustrated in Fig. 1-8. Watermolecules occur in the spaces between the ribbons. The 2:1 ribbons arewider in sepiolite than in palygorskite.

Palygorskite and sepiolite are often found in soils of arid and semiaridenvironments. Both minerals have a fibrous morphology in contrast to the platy morphology of most 1:1 and 2:1 minerals. Because of their fibrousmorphology, suspensions of these clays can form thick gels even at low solidconcentrations. Thus, palygorskite and sepiolite are used industrially asgelling agents to keep other solids in suspension.

C. Structural Details of Phyllosilicates

Figures 1-1 through 1-8 illustrate ideal mineral structures based on theregular close packing of spheres. As a consequence, all of the octahedra inFig. 1-4 are exactly the same size and all are regular (undistorted) solids.Likewise, in Fig. 1-5, all of the tetrahedra are regular solids, the 6-membertetrahedral rings form a regular hexagon, and the basal oxygens all lie inexactly the same plane. Crystallographic structure refinements of realphyllosilicate minerals show distortions from the ideal structures illustratedup until now in this chapter. These distortions are similar for all of thephyllosilicate minerals, although they vary somewhat in magnitude.

There are two reasons why real minerals deviate from these idealstructures. First, atoms are not hard spheres whose closest approach isdetermined solely by the sum of their radii. Second, if one does try to join


10/33


octahedral and tetrahedral sheets built of hard spheres like those illustratedin Fig. 1-4 and 1-5, geometry alone dictates that they will not fit together ifthe spheres representing the oxygen atoms touch each other in both theoctahedral and tetrahedral sheets. The reason becomes apparent byexamining the sphere-packing model of the tetrahedral sheet in Fig. 5. Sincethe apical oxygens are common to both the tetrahedral and octahedral sheets,two adjacent apical oxygens in the tetrahedral sheet also define the edge ofan octahedron in the octahedral sheet (Fig. 1-6a). Note that, although basaloxygens touch neighboring basal oxygens, apical oxygens do not touchneighboring apical oxygens (Fig. 5). Thus, the octahedral sheets illustratedin Fig. 1-4 cannot share apical oxygens with the tetrahedral sheet illustratedin Fig. 1-5 because the tetrahedral sheet is larger than the octahedral sheet.This problem was overcome in Fig. 1-6 to 1-8 by increasing the oxygen-oxygen distance in the octahedral sheet to match the oxygen-oxygendistance of the apical oxygens of the tetrahedral sheet. Thus, the octahedraremain undistorted but they are larger than they would be if the spheresrepresenting oxygen touched.

In the real phyllosilicate structures, this mismatch between theoctahedral and tetrahedral sheets is compensated, not by an expansion of theoctahedral sheet, but primarily by distortions in the tetrahedral sheet. The neteffect is that, the tetrahedra are rotated from the ideal arrangement shown inFig. 1-5 such that the hexagonal shapes outlined by the 6-memberedtetrahedral rings in Fig. 1-5 become the ditrigonal shapes illustrated in Fig.1-9. In addition, the tetrahedra are tilted such that the basal oxygens nolonger lie in the same plane, resulting in a corrugated basal plane in whichsome rows of atoms lie slightly below and some lie slightly above theaverage plane of the basal oxygens. These corrugations are best seen in thelowermost edge view in Fig. 1-9 where spheres representing the oxygenatoms have not been drawn. The implications of these distortions are thatmolecules interacting with phyllosilicate surfaces see, not the ideal surfaceillustrated in Fig. 1-6a, but the more distorted surface illustrated in Fig. 1-9.

The octahedral sheet is distorted from the ideal structure as well. First, because of the attraction of each central octahedral cation to its 6surrounding oxygens, octahedra occupied by cations are smaller thanoctahedra in which the central sites are unoccupied (Fig. 1-9). In addition,when two octahedra share two adjacent oxygens (the octahedra share edges),the oxygen-oxygen distance along this shared edge is shorter than theoxygen-oxygen distance along an edge that is not shared with an adjacentoccupied octahedron (Fig. 1-9). Thus, the octahedra are distorted from theideal Euclidean solids in very predictable ways. In general, trioctahedralphyllosilicates show less structural distortion and more closely resemble theideal structures than do the dioctahedral phyllosilicates.

The structural distortions just discussed probably account for relativelyminor differences in the properties of the phyllosilicates, with the majordifferences determined mainly by whether the mineral has a 1:1 or 2:1 layerstructure and by the layer charge. On the other hand, since dioctahedral

phyllosilicates are more common in soils than trioctahedral phyllosilicates,water, ions and molecules interact predominately with surfaces like thoseillustrated in Fig. 1-9, rather than surfaces with the more regulararrangement illustrated in Fig. 1-4 and 1-6a.

III. OTHER ALUMINOSILICATE MINERALS COMMON INSOILS

A. Zeolites

Zeolites are of a large group of minerals that consist structurally of SiO4tetrahedra arranged in ways that result in large amounts of pore space within

the crystals (Fig. 1-8). Aluminum substitutes for Si in the tetrahedral sitesand, as a result, the (Si,Al)O4 framework has a net negative charge. Thecharge is balanced by cations that reside in the channels and pores, alongwith water molecules. Because the cations are exchangeable, zeolites havecation exchange properties similar to the phyllosilicates, but because thetetrahedral framework of the zeolites is rigid and the size of the pores isfixed, small cations can move into and out of the pores freely, while largercations are excluded. Thus, zeolites are often referred to as molecularsieves because of their very selective cation exchange properties.

Zeolites are relatively rare in soils because they weather easily in humidregions, but they do occur in some soils in arid regions. Zeolites are widelyused in industry and agriculture, for example, as catalysts and adsorbents inthe chemical industry, as additives in animal feeds, and as ingredients inlaundry detergents. Thus, zeolites may be introduced into soils andsediments in unexpected ways.

B. Allophane and Imogolite

The aluminosilicate minerals discussed above have three-dimensionalcrystal structures with atoms packed together in a more or less regularmanner over relatively long distances (10's of nm's). They exhibit long rangeorder. Two other aluminosilicates, allophane and imogolite, exhibit short-range (orlocal) order. Structures with short-range order exhibit order overseveral nms, but on a larger scale the structure is disordered.

Allophane is a material consisting chemically of variable amounts ofO

2-, OH

-, Al

3+, and Si

4+, and characterized by short-range order and a

predominance of Si-O-Al bonds. It consists of small (3.5-5.0 nm) spheres,the structure of which has not been determined. The spheres clump togetherto form irregular aggregates.

Imogolite consists of tubes several m long with an outer diameter of2.3 to 2.7 nm and an inner diameter of ~l.0 nm. The tubes consist of a singledioctahedral sheet with the inner surface OH replaced by SiO 3OH groups(Fig. 1-8; Farmer, et al., 1983). Several individual tubes are arranged in


11/33


bundles 10-30 nm across to give thread-like particles several micrometerslong.

Allophane and imogolite usually occur as weathering products ofvolcanic ash and are important minerals in the Andisol soil order. Imogolitehas also been identified in the Bs horizons of Spodosols. Allophane andimogolite can specifically adsorb many inorganic and organic compounds.Andisols, for example, usually fix large amounts of phosphate, making itunavailable to plants, and the large amounts of organic matter common inAndisols may be due, in part, to adsorption of organic molecules byallophane and imogolite. Soils containing large amounts of allophane andimogolite usually have unique physical properties such as a low bulkdensity, high water holding capacity, high liquid and plasticity limits, and athixotropic consistence.

IV. SOME CRYSTALLOGRAPHIC CONCEPTS

In 1912 Max von Laue conducted an experiment to see if the "x-rays"discovered 17 years earlier by Wilhelm Conrad Rntgen could be diffracted by a crystal. This one experiment had two very important results: (1) itproved that x-rays behaved as waves, and (2) it showed that crystals weremade up of a regular array of atoms in space, and could therefore serve as agrating to diffract the x-rays. Within a few years of von Laue's discovery,much of the mathematical theory describing the diffraction of x-rays bycrystals had been developed and the atomic structures of some simplecrystals had been determined (Azroff, 1968; Klug and Alexander, 1974).As early as 1923 and 1924 several clays were shown to be crystalline by x-ray diffraction, but it was not until 1930 that some of the first structuredeterminations of clay minerals were made. About the same time, twoindependent papers, one by S. B. Hendricks and W. H. Fry in 1930 and theother by W. P. Kelly, W. H. Dore, and S. M. Brown in 1931, gave the first

x-ray diffraction evidence that soil materials, even those in the finestfractions, were crystalline (Marshall, 1949; Grim, 1968).

X-ray diffraction remains the most important method for identifying soilminerals. A working knowledge of certain concepts of crystallography andx-ray diffraction is necessary for understanding clay mineralogy. A briefintroduction to some of these concepts follows.

A. Periodicity in Crystals

In a crystal, a particular pattern or arrangement of atoms is repeated overand over in three dimensions. Repeating patterns occur not only in crystals but also in many familiar places such as wallpaper, tile floors, and brickwalls. Concepts that apply to crystals can be developed by starting withthese familiar examples.

Fig. 1-10a illustrates a pattern made up of leaves much like the patternon a piece of wallpaper. The periodicity of this pattern can be characterized

with three parameters. Along theXaxis the pattern is repeated at intervals ofa. We also say that the pattern is translated (moved) alongXat increments ofa. The translation interval along the Y axis is designated b and the anglebetween theXand Yaxes is designated by the angle . The array of leaves inFig. 1-10a consists of real objects, namely, leaves. We can represent the periodic nature of this pattern by an array of points in space, each pointhaving identical surroundings (Fig. 1-10b.). Such an array is called a lattice.A lattice of points is imaginary since each point is an imaginary infinitesimalspot in space. The concept of a lattice is useful because it allows us torepresent the periodic nature of a pattern regardless of the actual object orconfiguration of objects at each lattice point. The lattice depicted in Fig. 1-10b is called aplane lattice because it has only two dimensions. The conceptcan be extended to three dimensions to give a space lattice. A space lattice isa regular and unlimited distribution of points in space (Fig. 1-11). Theconcept of a space lattice is used in describing crystal structures becausecrystals are three dimensional objects

3. The directions of the crystallographic

axes in three dimensions are designated by X, Yand Z, and a, b, and c areused for the repeat distances along these axes (AIPEA NomenclatureCommittee, 1980). The angles between the axes are designated between YandZ, betweenXandZand betweenXand Y.

B. The Unit Cell

Joining the points of a space lattice produces a series of parallel-sidedunit cells (Fig. 1-11). Each unit cell contains a complete unit of the crystalpattern because the complete pattern is reproduced at each lattice point. Notethat there are alternative ways of outlining the unit cell. The choice of the"best" unit cell for a mineral is made when the structure is determined. Aunit cell is chosen for convenience in visualizing the symmetry and inmaking mathematical calculations. Different researchers will occasionally

choose different unit cells to describe the same mineral structure.There are 14 unique ways of arranging points in three-dimensional

space. These are known as the 14 Bravais lattices. They form thegeometrical basis for all possible unit cells. Each Bravais lattice belongs to 1of 6 crystal systems depending on the symmetry of the lattice (Table 1-5).The cubic system is most symmetric. All axial translations are equal (a = b= c) as are all axial angles ( = = ). The triclinic system is most general

3 The terms lattice and structure are often misused by authors and speakers. The twoterms are not synonymous.Lattice is the mathematical concept of an infinite uniformdistribution of points in space (e.g., the 14 Bravais lattices). Structure refers to the actual

physical assemblage of atoms (e.g., the 2:1 structure) (AIPEA Nomenclature Committee,1980). The concept of a lattice is used to describe the periodicity in crystal structures.


12/33


(and least symmetric). All axial translations and axial angles are different.The unit cell dimensions and the axial angles are important parameters indescribing the crystal structure of a mineral.

C. Miller Indices

A consistent and concise notation is used when planes of atoms in acrystal are discussed. Different planes are generally referred to by theirMiller indices, designated as (hkl). For example, we may speak of (001)planes, (110) planes and (431) planes. The concept of Miller indices followsfrom the lattice concept just discussed. Consider the lattice of points in Fig.1-12a and a plane that is parallel to c and passes through any two latticepoints (represented in an edge view by the heavy line). An identical planemust pass through each and every lattice point (light lines) because thelattice is periodic and each point is the same as any other. This family ofparallel planes cuts the a dimension of the unit cell into 2 parts and the bdimension into 3 parts. The c dimension is not cut at all (zero parts) becausethe planes are parallel to Z. The Miller indices are thus (230). In terms of(hkl), the planes cut the a dimension of the unit cell into h parts, theb dimension into k parts, and the c dimension into l parts. Additionalexamples of Miller indices are given in Fig. 1-12b.

D. X-ray Diffraction

X-ray diffraction (XRD) is the main analytical technique used both toidentify unknown mineral phases and to determine crystal structures. Soilminerals are usually studied using the powder diffraction method. Apowdered sample, with particles typically


13/33


program. Tables for determining d-values from 2 data are also given inBrindley and Brown (1980) for several common x-ray wavelengths.

The distances between the diffracting planes in the crystal, din eq. [7],are fixed by the structure of the crystal. However, x-ray wavelengths ('s)used for diffraction experiments are not always the same. If in eq. [7]changes, must change because d/n is a constant fixed by the crystalstructure of the sample. X-ray line positions expressed in 2 can only becompared when the x-ray wavelength is the same. Calculating the d-valuesfor the diffraction lines allows comparison of d-values regardless of thewavelength.

The d-values are directly related to the unit cell dimensions and theMiller indices. The d-value of a line with Miller indices (hkl), dhkl, in theorthorhombic system is given by the relationship:

dhkl = [(h2/a2) + (k2/b2) + (l2/c2)]-1/2, [8]

where a, b, and c are the unit cell dimensions. Relationships for the othercrystal systems are given elsewhere (for example, Klug and Alexander,1974).

The d/n values obtained from Bragg's Law can, in most cases, be relatedto actual distances between planes of atoms in the crystals. Thus, the d-values between successive phyllosilicate layers listed in Fig. 1-7 show up inx-ray diffraction patterns of soil clays and are used to identify the specificclay minerals present in unknown samples. In some cases however, in particular for randomly interstratified minerals or cases where diffractionlines are very broad due to the small size of the crystals, d/n values fromBragg's Law are not directly related to the distances between planes of atomsin the crystal. To explain these effects more inclusive diffraction theoryutilizing the concepts of the reciprocal lattice, structure factor, atomicscattering factors, etc. must be applied.

2. The Scherrer Equation

Once a mineral phase has been identified based on its diffraction peaks,additional information can often be obtained from diffraction line widths.Well-crystallized minerals of sand and silt size give sharp diffraction lineswhose widths are determined only by broadening caused by the x-raydiffractometer itself. Clay size particles show broader lines caused bydiffraction effects from the small particles. The smaller the particles are, thebroader the diffraction lines. X-ray line width and particle size are related bythe Scherrer equation:

Lhkl = (K)/( cos ), [9]where Lhkl is the mean crystallite dimension perpendicular to the diffractingplanes with Miller indices (hkl), Kis a constant equal to 0.9 if the particlesare cubes, is the x-ray wavelength, is the width (in radians) at one halfpeak height corrected for instrumental broadening, and is the diffractionline position. The application of the Scherrer equation involves aconsiderable number of assumptions (Klug and Alexander, 1974; p. 634-

642, 687-704) and an understanding of these assumptions is essential. Inaddition, Lhkl measures the size of the coherently diffracting domains withinthe crystals. If the crystals consist of several coherently diffracting domains,then Lhkl will be smaller than the actual physical size of the crystals observedby electron microscopy (Schulze and Schwertmann, 1984).

V. SUMMARY

This chapter has developed some of the terminology used later in this book. It has shown how the abundance of aluminosilicate minerals is aconsequence of the abundance of Si, Al, and O in the earth's crust. It hasintroduced the chemical and structural classification of minerals, mentionedthe major minerals that will be encountered in soils, and mentioned some ofthe environmental implications of some of the mineral. It has developed themajor structural themes of the phyllosilicate minerals, and it has explainedsome of the crystallographic and x-ray diffraction terminology used in laterchapters.

The chapters that follow will discuss in more detail the many facets ofsoil mineralogy, particularly as it relates to current environmental concerns.

VI. SUGGESTED READING

For a more information on general mineralogy, see introductory textssuch as Klein and Hurlburt (1993) or Nesse (1999). Azroff (1968) andMoore and Reynolds (1997) provide good, readable introductions tocrystallography and x-ray diffraction theory.

VII. QUESTIONS AND EXERCISES

1. Draw sketches from memory of an octahedron and a tetrahedron using

the sphere-packing, ball-and-stick, and polyhedral representations.2. Draw sketches from memory of sphere-packing models of a tetrahedral

sheet, a dioctahedral sheet, and a trioctahedral sheet. (A coin makes agood template for a circle.)

3. Draw three sketches of a unit cell and label the cell dimensions a, b, andc as shown in Fig. 1-12b. On the first cell sketch in the planes withMiller indices (002), on the second cell sketch planes with indices (111),and on the third cell sketch planes with indices (021).

4. a. Calculate the d-values of x-ray diffraction lines that occur at 8.84,12.36, and 26.68 2 for a sample x-rayed using CuK radiation (=0.15418 nm). b. Calculate the 2 positions for these same three lineswhen the sample is x-rayed using CoK radiation (= 0.179026 nm).Answers: a. 1.000, 0.716, and 0.334 nm, respectively. b. 10.27, 14.36,and 31.09 2, respectively.

5. A goethite sample was x-rayed with CoK radiation. The 111diffraction line was centered at 24.70 2 and its width at half-height


14/33


(corrected for instrumental broadening) was found to be 0.35 2. Usethe Scherrer equation to calculate the size of the crystals (more exactly,the size of the coherently diffracting domains) perpendicular to the(110) planes. Assume that K= 0.9. Answer: 27 nm.

VIII. ACKNOWLEDGMENTS

I thank the many students over the years that have provided feedback onthis chapter.

IX. REFERENCES

AIPEA Nomenclature committee. 1980. Summary of recommendations

of AIPEA nomenclature committee. Clays Clay Min. 28:73-78.Allen, B. L. and D. S. Fanning. 1983. Composition and soil genesis.

p. 141-192. In L. P. Wilding, N. E. Smeck, and G. F. Hall (ed.) Pedogenesisand soil taxonomy I. Concepts and interactions. Elsevier, New York.

Azroff, L. V. 1968. Elements of x-ray crystallography. McGraw-HillBook Co., New York. 610 p.

Brindley, G. W. and G. Brown (eds). 1980. Crystal structures of clayminerals and their x-ray identification. Mineralogical Society, London.495 p.

Farmer, V. C., M. J. Adams, A. R. Fraser, and F. Palmieri. 1983.Synthetic imogolite: Properties, synthesis, and possible applications. ClayMin. 18:459-472.

Grim, R. E. 1968. Clay mineralogy. 2nd Ed. McGraw-Hill Book Co.,New York. 596 p.

Jackson, M. L. 1964. Chemical composition of soils. p. 71-141. In F. E.Bear (ed.) Chemistry of the soil. Reinhold Publishing Co., New York.

Kmpf, N., A. C. Scheinost, and D. G. Schulze. 1999. Oxide minerals.

p. F125 F168. In: M. E. Sumner (ed.), Handbook of Soil Science. CRCPress, Boca Raton, FL.

Klein, C. and C. S. Hurlbut, Jr. 1993. Manual of mineralogy (afterJames D. Dana). 21st Ed. John Wiley & Sons, New York. 681 p.

Klug, H. P. and L. E. Alexander. 1974. X-ray diffraction procedures forpolycrystalline and amorphous materials. 2nd Ed. John Wiley & Sons, NewYork. 966 p.

Marshall, C. E. 1949. The colloid chemistry of the silicate minerals.Academic Press Inc., New York. 195 p.

Moore, D. M. and R. C. Reynolds, Jr. 1997. X-ray diffraction andidentification and analysis of clay minerals. 2

nded. Oxford University Press,

New York. 378 p.Nesse, William D. 1999. Introduction to mineralogy. Oxford University

Press, New York. 464 p.Rothbauer, R. 1971. Untersuchung eines 2M1-Muskovits mit

Neutronenstrahlen. N. Jahrbuch f. Mineralogie. Monatsheite 1971:143-154.

Schulze, D. G. and U. Schwertmann. 1984. The influence of aluminiumon iron oxides: X. Properties of Al-substituted goethites. Clay Min. 19:521-539.

Sposito, G. 1989. The chemistry of soils. Oxford University Press, NewYork. 277 p.

X. LIST OF FIGURES

Fig. 1-1. The structure of halite (NaCl).

Fig. 1-2. Hexagonal closest packing of spheres in a plane.

Fig. 1-3. Octahedra and tetrahedra as a consequence of two planes of close-packed spheres and three ways of representing octahedra and tetrahedra.

Fig. 1-4a. A trioctahedral sheet.

Fig. 1-4b. A dioctahedral sheet.

Fig. 1-5. The tetrahedral sheet.

Fig. 1-6. (a) Oblique view of the 1:1 layer structure illustrating the

relationship between the tetrahedral and octahedral sheets. Note that

adjacent apical oxygens of the tetrahedral sheet (arrows a) also define

edges of octahedra in the octahedral sheet. Arrows marked b point to OHs

that lie directly in the center of the hexagonal rings of tetrahedra, although

they appear off-center in this oblique view. (b) Edge view of the 1:1 and 2:1

layer structures illustrating phyllosilicate nomenclature. (c) An alternate

edge view that arises when the 2:1 structure above is rotated normal to theplane of the layer.

Fig. 1-7. Structural scheme of soil minerals based on octahedral and

tetrahedral sheets.

Fig. 1-8. Structural models of representatives of three other aluminosilicate

mineral groups that occur frequently in soils. All three are drawn to the same

scale.

Fig. 1-9. Structural details of phyllosilicates as illustrated by the octahedral

and tetrahedral sheets of muscovite. Figures were prepared from the single-

crystal structural refinement data of Rothbauer (1971).

Fig. 1-10. (a) A lattice array of leaves and (b) a plane lattice of points.


15/33


Fig. 1-11. A space lattice with several alternative unit cells outlined (after

Klug and Alexander, 1974).

Fig. 1-12. (a) Planes with Miller indices 230 (c axis is out of the plane of the

figure) and (b) planes with different Miller indices on an orthogonal lattice.

Fig. 1-13. (a) In phase and out of phase waves, and (b) geometry of the

Bragg "reflection" analogy for diffraction of x-rays by crystals (after

Azroff, 1968).

XI. LIST OF TABLES

Table 1-1. The twelve most common chemical elements in the earth's crust

(Klein and Hurlbut, 1993).

Table 1-2. Classification of common non-silicate minerals in soils.

Table 1-3. Classification of primary silicate minerals.

Table 1-4. Ionic radius, radius ratio and coordination number of common

elements in silicate minerals (Klein and Hurlbut, 1993).

Table 1-5. Axial ratios and angles between crystal axes for the six crystal

symmetry systems.

Table 1-1. The twelve most common chemical elements in the earth's crust(after Klein and Hurlbut, 1993).

Element Crustal

average

Mole fraction Ionic radius Volume

g kg-1

(nm) %

O 466.0 0.6057 0.136 (3) 92.88

Si 277.2 0.2052 0.026 (4) 0.22

Al 81.3 0.0626 0.039 (4) 0.23

Fe 50.0 0.0186 0.078 (6) 0.54

Ca 36.3 0.0188 0.100 (6) 1.15

Na 28.3 0.0256 0.102 (6) 1.66

K 25.9 0.0138 0.151 (8) 2.89

Mg 20.9 0.0179 0.072 (6) 0.41

Ti 4.4 0.0019 0.061 (6) 0.03

H 1.4 0.0289

P 1.0 0.0007 0.017 (4) 0.00

Mn 0.9 0.0003 0.083 (6) 0.01

Numbers in parentheses refer to coordination number. Radii for Fe and Mn are

for the reduced (2+) form.

Ionic radius and volume of H+

are negligible compared to O2-

.


16/33


Table 1-2. Common non-silicate minerals in soils.

Mineral class Mineral Chemical formula

Halides Halite NaCl

Sulfates Gypsum CaSO42H2O

Jarosite KFe3(SO4)2(OH)6

Sulfides Pyrite FeS2

Carbonates Calcite CaCO3

Dolomite CaMg(CO3)2

Nahcolite NaHCO3

Trona Na2CO3NaHCO32H2O

Soda Na2CO310H2O

Oxides and Hydroxides

Aluminum Gibbsite Al(OH)3

Iron Hematite Fe2O3

Goethite FeOOH

Lepidocrocite FeOOH

Maghemite Fe2O3

Ferrihydrite Fe5O7(OH)4H2O

Magnetite Fe3O4

Manganese Birnessite (Na,Ca,Mn2+

) Mn7O42.8

H2O

Lithiophorite LiAl2Mn24+

Mn3+

O6(OH)6

Hollandite Ba(Mn4+

,Mn3+

)8O16

Todorokite (Na,Ca,K)0.3-0.5(Mn4+

,Mn3+

)6

O123.5H2OTitanium Rutile TiO2

Anatase TiO2

Ilmenite Fe2+

TiO3

after Klein and Hurlburt (1993), Kmpf et al. (1999)

Table 1-3. Classification of silicate minerals.

Silicate class,

unit composition,

Arrangement of

SiO4 tetrahedra Mineral Ideal Formula

Nesosilicates (SiO4)4-

Olivine (Mg,Fe)2SiO4Forsterite Mg2SiO4Fayalite Fe2SiO4

Zircon ZrSiO4Sphene CaTiO(SiO4)

Topaz Al2SiO4(F,OH)2Garnets X3Y2(SiO4)3,

X=Ca,Mg,Fe2+

,Mn2+

,

Y=Al,Fe3+

,Cr3+

Andalusite Al2SiO5Sillimanite Al2SiO5Kyanite Al2SiO5Staurolite Fe2Al9O6(SiO4)4(O,OH)2

Sorosilicates (Si2O7)6-

Epidote Ca2(Al,Fe)Al2O(SiO4)

(Si2O7)(OH)

Cyclosilicates (Si6O18)12-

Beryl Be3Al2(Si6O18)

Tourmaline (Na,Ca)(Li,Mg,Al)(Al,Fe,

Mn)6 (BO3)3(Si6O18)(OH)4

Inosilicates Pyroxenes

(single chains) (SiO3)2- Augite (Ca,Na)(Mg,Fe,Al)(Si,Al)2O6Enstatite MgSiO3Hypersthene (Mg,Fe)SiO3Diopside CaMgSi2O6Hedenbergite CaFeSi2O6

Pyroxenoids

Wollastonite CaSiO3Rhodonite MnSiO3

Inosilicates Amphiboles

(double chains) (Si4O11)6-

Hornblende (Ca,Na)2-3(Mg,Fe,Al)5Si6(Si,Al)2O22(OH)2

Tremolite Ca2Mg5Si8O22(OH)2Actinolite Ca2(Mg,Fe)5Si8O22(OH)2


17/33


Cummingtonite (Mg,Fe)7Si8O22(OH)2Grunerite Fe7Si8O22(OH)2

Phyllosilicates (Si2O5)2-

Micas

Muscovite KAl2(AlSi3O10)(OH)2Biotite K(Mg,Fe)3(AlSi3O10)(OH)2Phlogopite KMg3(AlSi3O10)(OH)2

Chlorites (Mg,Fe)3(Si,Al)4O10(OH)2..........

Clay minerals

(selected)

Talc Mg3Si4O10(OH)2

Pyrophyllite Al2Si4O10(OH)2Kaolinite Al2Si2O5(OH)4Smectite M

+0.3Al2(Al0.3Si3.7)O10(OH)2

M+= Ca

2+, Mg

2+, K

+, etc.

Vermiculite M+

0.7Al2(Al0.7Si3.3)O10(OH)2M

+= Ca

2+, Mg

2+, K

+, etc.

Serpentines

Antigorite Mg3Si2O5(OH)4Chrysotile Mg3Si2O5(OH)4

Tectosilicates (SiO2)0

Feldspars

Orthoclase KAlSi3O8Albite NaAlSi3O8Anorthite CaAl2Si2O8

SiO2 Group

Quartz SiO2Tridymite SiO2Cristobalite SiO2

Zeolites

Analcime NaAlSi2O6H2O

Feldspathoids

Nephelene (Na,K)AlSiO4 after Allen and Fanning (1983).

Klein and Hurlbut (1993).

Table 1-4. Ionic radii and coordination of common elements in phyllosilicate

minerals.

Ion Ionic radius

(nm)

O2-

0.140(6)

F-

0.133(6)

Cl

-

0.181(6)

Si4+

0.026(4)

Al3+

0.054(6)

Fe3+

0.065(6)

Mg2+

0.072(6)

Ti4+

0.061(6)

Fe2+

0.078(6)

Mn2+

0.083(6)

Na+

0.118(8)

Ca2+

0.112(8)

K+

0.151(8)

Ba2+

0.142(8)

Rb+

0.161(8)

Ionic radii vary with coordination number. Radii are for ions with the

coordination numbers listed in parentheses (from Klein and Hurlbut, 1993, p.

188).

tetrahedral sites

octahedral sites

cubic and

dodecahedral sites


18/33


Table 1-5. Axial ratios and angles between crystal axes for the six crystal

symmetry systems (after Klug and Alexander, 1974).

System Axial ratios Angles between crystal axes

Cubic (isometric) a = b = c = = = 90

Hexagonal a = b c = = 90, = 120

Tetragonal a = b c = = = 90

Orthorhombic a b c = = = 90

Monoclinic a b c = = 90, > 90

Triclinic a b c = 90


19/33

Inserts for Table 3

Nesosilicates

Sorosilicates

Inosilicates(double chains)

Inosilicates(single chains)

Cyclosilicates

Phyllosilicates

Tectosilicates


20/33

Na Cl

Fig. 1-1. The structure of halite (NaCl).


21/33

Fig. 1-2. Hexagonal closest packing of spheres in a plane.

"A" voids "B" voids


22/33

Face Edge

Tetrahedron

Fig. 1-3. Octahedra and tetrahedra as a consequence of two planes of close-packed spheres and three ways of representing octahedra and tetrahedra.

Sphere-PackingModel

Ball-and-StickModel

PolyhedralModel

Octahedron OctahedronTetrahedron

"A" site "B" site


23/33

The Octahedral Sheet (trioctahedral)

Figure 1-4a. A trioctahedral sheet. The upper three rows of hydroxyls in the face view of the sphere-packing model have been omitted for clarity.

Mg2+

OH-


24/33

The Octahedral Sheet (dioctahedral)

Figure 1-4b. A dioctahedral sheet. The upper three rows of hydroxyls in the face view of the sphere-packing model have been omitted for clarity.

Al3+

OH-


25/33

The Tetrahedral Sheet

Figure 1-5. The tetrahedral sheet.

apical oxygens

Si4+

basal oxygens


26/33

Fig. 1-6. (a) Oblique view of the 1:1 layer structure illustrating the relationship between

the tetrahedral and octahedral sheets. Note that adjacent apical oxygens of the

tetrahedral sheet (arrows a) also define edges of octahedra in the octahedral sheet.

Arrows marked b point to OHs that lie directly in the center of the hexagonal rings

of tetrahedra, although they appear off-center in this oblique view. (b) Edge view of the

1:1 and 2:1 layer structures illustrating phyllosilicate nomenclature. (c) An alternate edge

view that arises when the 2:1 structure above is rotated normal to the plane of the layer.

ab

b

a

tetrahedral cationsbasal O's

PLANES OF IONS

OH's & apical O'soctahedral cations

OH's

1:1layer

tetrahedralsheet

octahedralsheet

SHEETS, LAYERS


OH's & apical O'soctahedral cations


OH's & apical O's

2:1layer

tetrahedralsheet

octahedralsheet

tetrahedralsheet

c

b

Alternateedge view


27/33

Mica (muscovite)

1.00 nm(10.0 )

fixed K+

0.46 nm(4.6 )

Fig. 1-7. Structural scheme of soil minerals based on octahedral and tetrahedral sheets.

1.00 nm(10.0 )

0.92 nm(9.2 )

Pyrophyllite

Halloysite (hydrated)

Kaolinite

Gibbsite

0.72 nm(7.2 )

Al3+

OH-

O2-

H2O

Si4+

Al3+

Si4+

Al3+

Si4+

Al3+

Si4+,Al3+

Al3+


28/33

>1.8 nm(>18 )

Smectite

Vermiculite

Si4+,Al3+

Al3+

hydr. Ca2+,Mg2+, etc.

1.4 nm(14 )

Chlorite

Hydroxy-interlayeredVermiculite and Smectite

Fig. 1-7. Continued.

1.4 nm

(14 )

1.4 nm(14 )

exch. Ca2+,Mg2+, etc.

H2O

Al3+ poly.,exch. Ca2+,Mg2+, etc.

Mg2+, Al3+,

Fe

3+

, etc.


29/33

Palygorskite

Clinoptilolite (a zeolite)

Imogolite

1.0 nm(10 )

Si

4+

Si4+,Al3+

H2O

H2O(not shown)

exch. cations& H2O

(not shown)

Si4+Al3+

H2O(not shown)

Fig. 1-8. Sturctural models of representatives of three other aluminosilicate mineral

groups that occur frequently in soils. All three are drawn to the same scale as Fig. 1-9.

Mg2+,Al3+


30/33

Fig. 1-9. Structural details of phyllosilicates as illustrated by the octahedral and tetrahedral sheets of muscovite. Figures prepared from the single-

crystal structural refinement data of Rothbauer (1971).

Structural Details Common to Phyllosilicate Minerals

Tetrahedral Sheet Octahedral Sheet

ditrigonalhole

occupiedoctahedra

smaller thanunoccupied

corrugatedbasal plane

shared edgesshorter than

unshared


31/33

a

b g

a b

Fig. 1-10. (a) A lattice array of leaves and (b) a plane lattice of points.


32/33

Fig. 1-11. A space lattice with several alternative unit cells outlined (after Klug and Alexander, 1974).


33/33

Fig. 1-12. (a) Planes with Miller indices 230 (c axis is out of the plane of the figure) and

(b) planes with different Miller indices on an orthogonal lattice.

Fig. 1-13. (a) In phase and out of phase waves, and (b) geometry of the Bragg "reflection" analogy for diffraction of x-rays by crystals (after Azroff, 1968).

a

b

Documents

Schulze Chap1