Chapter 1

An Introduction to Soil Mineralogy

OARRELL G SCHULZE, Purdue University, West Lafayette. Indiana

MineraIs make up about one-halfofthe volume ofmost soils. They provide phys­ical support for plants and create the water- and air-filled pores that make plantgrowth possible. Mineral weathering releases plant nutrients that are retained byother minerais through adsorption, cation exchange, and precipitation. Minerais areindicators ofthe amount ofweathering lhat has taken place and the presence or ab­sence ofparticular minerais gives c1ues to how soils fonned. The physical and chem­ical characteristics ofsoil minerais are important considerations in planning, con­structing, and maintaining buildings, roads, and airports. Soil mineraIs can adsorbmany organic and inorganic environmental pollutants, promoting their degradationto nontoxic fonns, attenuating their movement through the soil, or preventing theiruptake by plants and their introduction into lhe food chain. Some minerais are them­selves pollutanls and can cause serious environmental problems when they are ex­posed to weathering at the soil surface by human activities. An understanding ofsoil mineralogy is central lo understanding virtually ali facets ofhumanity's use andmisuse ofsoils and is ofien the key to solving specific environmental problems.

This chapter develops a core ofconcepts and tenninology needed for under­standing soil mineraIs. The chemical composition ofthe Earth's crust is discussedfirst to show that the most abundant elements in the crust are, not surprisingly, theones most likely to be encounlered in soil minerais. Then the chemical and struc­tural c1assification ofminerals is discussed, and the major mineraIs represented insoils are mentioned. The phyllosilicate mineraIs are covered separately because oftheir major role in soils. Basic structural concepts common to ali minerais are cov­ered at lhis point to provide the background for the discussion ofthe phyllosilicatestructures. The overall structural theme ofthe phyllosilicates is then presented, alongwith a very briefsummary oftheir most importanl properties. The chapter concludeswith a treatment ofsome crystallographic and x-ray diffraction concepts importantin identifYing and characterizing soil minerais.

I. CHEMICAL AND STRUCTURAL CLASSIFICATION OF MINERALS

A. Composition ofthe Earth's Crust

Most ofthe weight and volume oflhe Earth's crusl is made up ofonly a fewelements (Table 1-\). Oxygen and Si make up mosl of lhe weight, while oxygen

Copyright 2002 Cd Soil Science Society ofAmerica, 677 South Segoe Road. Madison. WI 53711, USA.Soil Mineralogy with Environmental Applications. SSSA Book Series. no. 7.

1

2 SCHULZE

Table I-I. lhe 12 most eommon ehemieal e1ements in the Earth's erust (after Klein & Hurlbut. 1993).

Element Crustal average Mole fraetion lonie radiust Volume

gkg- I nm %

O 466.0 0.6057 0.136 (3) 92.88Si 2772 0.2052 0.026 (4) 0.22AI 81.3 0.0626 0.039 (4) 0.23Fe 50.0 0.0186 0.078 (6) 0.54Ca 36.3 0.0188 0.100(6) 1.15Na 28.3 0.0256 0.\02(6) 1.66K 25.9 0.0138 0.151 (8) 2.89Mg 20.9 0.0179 0.072 (6) 0.41Ti 4.4 0.0019 0.061 (6) 0.03H 1.4 0.0289 ~ ~P 1.0 0.0007 0.017 (4) <0.01Mn 0.9 0.0003 0.083 (6) 0.01

t Numbers in parentheses refer to coordination number. Radii for Fe and Mn are for lhe redueed (2+)fonn.

~ lonie radius and volume ofW are negligible eompared wilh ()l-.

alone accounts for more than 90% of the volume. ln general, the larger O atomsare in an approximately close-packed arrangement held together by smaller metalatoms in the interstitial space. Most ofthe elements in the crust and in soils occurin minerais. Thus, the elements listed in Table I-I are major constituents ofthe min­erais discussed in this book.

B. Definition or a Mineral

Klein and Hurlbut (1993) define a mineral as follows: "A mineral is a natu­rally occurring homogeneous solid with a definite (but not generally fixed) chem­ical composition and a highly ordered atomic arrangement. It is usually forrned byinorganic processes." 80th chemical composition and crystal structure (orderedatomic arrangement) are important parts of this definition. Neither alone is suffi­cient to explain the properties of minerais. Minerais with similar chemical com­position but different crystal structures-or, conversely, minerais with similar crys­tal structures but different chemical compositions-<:an be quite different from oneanother despi te their chemical or structural similarities.

C. Mineral Classification

Although different classification schemes could be used, mineralogists havedeterrnined that first separating mineraIs into groups on the basis oftheir chemicalcomposition gives classes with the greatest similarities in many other properties.Thus, minerais are first divided into classes depending upon the dominant anion oranionic group. The classes include: (i) native elements. (ii) sulfides, (iii) sulfosalts,(iv) oxides and hydroxides. (v) halides, (vi) carbonates, (vii) nitrates, (viii) borates,(ix) phosphates, (x) sulfates, (xi) tungstates, and (xii) silicates (Klein & Hurlbut,1993). These classes are then subdivided based on chemical and structural simi­larities. This sarne general classification is followed in this book, but with some ex­ceptions. Mineral classes such as the native elements, sulfosalts. nitrates. borates,

INTRODUCTION TO SOIL MINERALOGY

Table 1-2. Common nonsilicate minerais in soils.

s

Mineral c1ass

HalidesSulfates

SulfidesCarbonates

Oxides and hydroxidesAluminumlron

Manganese

Titanium

Mineral

HaliteGypsumJarositePyriteCalciteDolomiteNahcoliteTronaSoda

GibbsiteHematiteGoethiteLepidocrociteMaghemiteFerrihydriteMagnetiteBimessiteLithiophoriteHollanditeTodorokiteRutileAnataseOmenite

Chemical fonnulat

NaCICaSO.0 2H20Kfe)(SO.h(OH>6Fe~

CaCO)CaMg(CO)hNaHCO]Na2CO)oNaHCO)02H20Na2CO)0IOH20

AI(OHhFe20]FeOOHFeOOHFe20 ]FejÜ-7(OH)04H20fe)O.(Na,Ca,Mn2+) Mn,o.02.8 H20LiAI2Mn2·+Mn]+ 06(OH)6Ba(Mn·+,Mn]+)80 16 .(Na,Ca,K>O.]-ll.s(Mn·+,Mn)+)6 0 1203.SH20Ti02TiO,Fe2+-TiO,

t After Klein & lIurlbun (1993), Kãmpf et aI. (1999).

and tungstates that occur only rarely in soils, are covered only brietly or not at ali.Commonly occurring mineral classes, particularly the silicates, are covered in de­tail.

Soil minerais are also referred to as eitherprimary or secondary minerais. Pri­mary minerais form at elevated temperatures. They are usually derived from igneousor metamorphic rocks, but they can be inherited from sedimentary rocks as well.Secondary minerais are formed by low-temperature reactions and are either inher­ited from sedimentary rocks or formed in soils by weathering (Jackson, 1964). Theseparation ofminerals into primary and secondary mineral classes is not mutuallyexclusive; some minerais can occur in both. The concept is useful however, and ap­pears widely in the soil science literature. The major mineral classes representedin soils are discussed below.

lo Halide, Sulfate, and Carbonate Minerais

The major soil minerais of this group are halite (NaCI), gypsum(CaS04·2H20), calcite (CaC03), and dolomite [CaMg(C03hl (Table 1-2). Thisgroup is characterized by minerais with relatively simple structures. The halite struc­ture is one ofthe simplest ofali minerais. It consists ofaltemating Na+ and CI- ionsarranged in cubic c10sest packing (Plate I-I). The other minerais in this group havesimilar structures, with cations such as Ca2+, Mg2+, or Fe2+altemating with anionssuch as CI-, S01-, or COI-o The bonds between the cations and anions are pre­dominantly ionic. These mineraIs are among the most soluble and the softest of ali

4 SCHULZE

soil minerais, and they are easily broken down by physical and chemical weather­ing processes. They occur mainly in soils ofarid regions or in youthful soils in morehumid regions, where weathering has been minimal.

Halite is the most soluble ofthis group and accumulates in only the most aridenvironments. It is one ofthe minerais present in the salic horizon ofAridisols. Gyp­sum is about 100 times less soluble than halite, but it too is abundant only in aridregions. Gypsum is a major mineral in the gypsic horizon of Aridisols. Halite andgypsum could also occur in soils that have been contarninated by salt water, or bysoluble salts leaching from industrial stock piles or waste piles.

Calcite and doJomite are common carbonate mineraIs that occur in a varietyofsoils. These minerais precipitate in the soil profile in arid and semiarid climates(aridic and some ustic soiJ moisture regimes). Calcic and petrocalcic horizonsfonn ifthe accumulation is great enough. Carbonate minerais in many soils are in­herited from Jimestone or other calcareous parent materiais. Carbonates are usu­ally stable and can be found throughout the soil profile under aridic and ustic soilmoisture regimes (arid to subhumid cJimates). but are leached from the soil profileand are generally present only in the C horizons under the udic soil moistureregime (humid climates). Calcite and dolomite can be introduced into soils origi­nally free ofthese minerais via lhe limestone aggregate used for road constructionin some areas.

2. Sulfides

Pyrite, FeS2' the most common mineral in this group. does not occur exten­sively in soils. but when it is present it causes some unique problems. Pyrite pre­cipitates in soils on wet tidal tlats and river deltas of some coastal areas and alsooccurs in some geologic fonnations originally deposited in similar environments.Thus. pyrite ofien occurs in cJose association with coaI. Pyrite is unstable under ox­idizing conditions and weathers quickly when pyritic soils are drained or when min­ing leaves pyritic material on the surface. The weathering products incJude the min­erais jarosite, KFe3(S04h(OH)6' and gypsum and sulfuric acid, H2S04. The largeamount of acidity causes problems in using pyrite-containing soils and in revege­tating mined areas and is a serious and costly environrnental problem.

3. Oxides, Hydroxides, and Oxybydroxides

Primary minerais break down during weathering and release cations and an­ions that recombine to fonn other more stable minerais. Several elements, in par­ticular AI. Fe. and Mn. fonn oxide. hydroxide. or oxyhydroxide minerais that arestable in the soil weathering environment. Representative mineral species are givenin Table 1-2.

a. Aluminum. Gibbsite, [AI(OHhl is the most common AI hydroxide min­eraI in soils. It is generally associated with the latter stages ofweathering when leach­ing ofsilica has progressed to the point that phyllosilicate minerais no longer fonn.Gibbsite is common in highly weathered Oxisols oftropical regions. It has a verylow cation exchange capacity and contributes to the low native fertility ofsome Ox­isols. Gibbsite is also commonly found at the weathering interface between igneousrock and saprolite and in Andisols fonned from volcanic ash.

INTRODUCTION TO SOIL MINERALOGY 5

b. lron. lron oxide minerais fonn from Fe released from primary mineraIs.Iron oxides are strong pigments and small amounts of these minerais account formost ofthe brown and red colors ofsoils. Goethite (FeOOH) is the most commonmineral ofthis group. It accounts for the brownish to yellowish color ofmany soils,although it may be present in only small quantities. Hematite (Fe203) is onlyslightly less common than goethite and usually occurs in association with it.Hematite is usually bright red and is responsible for the red color of many soils.Goethite and hematite are stable minerais in an oxidizing environment. Largeamounts ofthese two minerais in well-aerated soils, usually in association with gibb­site and kaolinite, usually indicate an advanced stage of weathering. ln soils thatare saturated with water for at least some time during the year, the very insolubleFe3+ in goethite, hematite, and other Fe oxides can be reduced to very soluble Fe2+.

The Fe2- can easily move with the soil water to other parts ofthe soil profile, oreven to other associated soils in the landscape, where it can then reoxidize to Fe3+

and reprecipitate as goethite, lepidocrocite, or ferrihydrite when oxidizing condi­tions retum. Repeated cycles of oxidation and reduction give rise to mottJes andnodules that ref1ect an inhomogeneous distribution of Fe oxide minerais within asoil. Soil scientists use these inhomogeneous or mottled color patterns to estimatethe depth to a soil water table and the length oftime the soil is saturated during theyear. Thus, when delineating wetlands, detennining the best location for a septicsystem leach field, or mapping soils, soil scientists rely on predictable soil colorpatterns that result from Fe oxide mineralogy and distribution.

e. Manganese. Manganese oxides and hydroxides (Table 1-2) are com­monly found in soils as brown or black nodules or as thin coatings on the faces ofsoil structural units. They are ofien associated with Fe oxides. Manganese occursfrequentlyas birnessite or lithiophorite in soiJs. Manganese can be oxidized and re­duced in the soil environment similar to Fe. Thus, the inhomogeneous distributionofMn into nodules is an indicator ofreduction and oxidation as a result of periodicwater saturation.

d. Titanium. Rutile and ilmenite (Table 1-2) occur in soils mainly as primaryminerais inherited from igneous rocks. Anatase is less common and is generally con­sidered a secondary mineral. Although frequently found in soils, these minerais gen­erally do not impact soil physical or chemical properties.

4. Sitieates

The silicate mineral c1ass is an extremely large and important group ofmin­erals. Nearly 40% ofthe common minerais are silicates, as are most minerais in ig­neous rocks. Silicates constitute well over 90% ofthe Earth's crust (Klein & Hurl­but, 1993) and comprise the bulk of most soils as well. Silicates occur as bothprimary minerais inherited from igneous or metamorphic rocks and as secondaryminerais formed from the weathering products ofprimary minerais.

As explained in more detail below, the fundamental unit ofall silicate struc­tures is the Si04 tetrahedron. It consists of four 0 2- ions at the apices of a regulartetrahedron coordinated to one Si4+ at the center. The individual tetrahedra are linkedtogether by sharing 0 2- ioos to fonn more complex structures. Several different

6

Table )-3. Classification ofsilicate minerais.

SCHULZE

Silicate c1ass, unit composition,arrangement ofSi04 tetrahedrat

Nesosilicates (Si04)4-

Sorosilicates (Si20 7)1t-

AvCyclosilicates (Si60 18)12-

*Inosilicates(single chains) (Si03)2-

Mineral

OlivineForsteriteFayaJite

ZirconSpheneTopazGamets

AndalusiteSillimaniteKyaniteStaurolite

Epidote

BerylTounnaline

PyroxenesAugiteEnstatiteHyperstheneDiopsideHedenbergite

PyroxenoidsWollastoniteRhodonite

Ideal fonnulat

(Mg.FehSiO.Mg2SiO.Fe2SiO.

ZrSiO.CaTiÜ(Si04)

AI2SiO.(F,OHhX)Y2(SiO.h. X = Ca,Mg,Fe2+,

Mn2+. Y= AI,Fe)+'Crl+AI2SiOsAI2SiOsAI2SiOsFe2Alq06(SiO.).(O,OHh

Ca2(AI,Fe)AI2Ü(SiO.)(Si20 7)(OH)

Be)AI2(Si60 18)

(Na,Ca)(Li,Mg.AI)(AI,Fe,Mn)6(BO)h(Si60 18)(OH).

(Ca,Na)(Mg.Fe,AI)(Si,Alh06MgSiO)(Mg.Fe)SiO)CaMgSi20 6CaFeSi20 6

Inosilicates(double chains) (Si.OI\)1t-

AmphibolesHomblende (Ca,Nah.)(Mg.Fe,AI)sSi6

(Si,AI}z°22(OHhTremolite Ca2MgsSis022(OHhActinolite Ca2(Mg,Fe)sSisOn(OHhCummingtonite (Mg.Fe),Sis0 22(OHhGrunerite Fe7SisOn(OHh

(continued on next page)

arrangements ofthe Si04 tetrahedra occur, partly accounting for the large numberof silicate minerais and providing the basis for their c1assification. The tetrahedramay be present as single tetrahedra (nesosilicates), double tetrahedra (sorosilicates),rings (cyclosilicates), single or double chains (inosilicates), sheets (phyl1osili­cates), or three-dimensional frameworks (tectosi Iicates) (Table 1-3). ln most ofthesearrangements adjacent Si04 tetrahedra share comers; that is, they share a common0 2-. The more common minerais from each silicate c1ass likely to be found in soilsare given in Table 1-3.

The most common mineraIs in igneous rocks are the olivines, pyroxenes, am­phiboles, micas. fe1dspars, and quartzo These primary mineraIs make up most ofthesand and silt size fractions of soils. The clay «2 ~m) fraction of most soils con­sists primarily ofphyl1osilicate mineraIs. Because oftheir importance in soils andsediments, phyl10silicates are discussed in detail below and in other chapters ofthisbook.

INTRODUCTION TO SOIL MINERALOGY 7

Table 1-3. Continued.

Silicate c1ass, unit composition,arrangement ofSi04 tetrahedrat Mineral Ideal formulat

PhyJlosilicates (Si~05)~- Micas

*Muscovite KAI2(AISi)OIO)(OHhBiotite K(Mg,Feh(AISi)OIO)(OHhPhlogopite KMg)(AISi)OIO)(OHh

Chlorites (Mg,Feh(Si.AI)401O(OHh(Mg,Feh(OH)(,

Clay minerais(seJected examples)

Tale Mg)Si4O IO(OHhPyrophyJlite AI2Si4O IO(OHhKaolinite AI2Si2O;(OH)4Smectite M+o)AI2(Alo.)Sí3.7l0IO(OHb

M+= Ca2+, Mi+' K+, etc.Vermiculite M+o7AI2(Alo.7Siu)OIO(OHh.

M+= Ca2+, Mg2+, K+, etc.Serpentines

Antigotite Mg)Si20 5(OH)4Chrysotile Mg)Si2O;(OH)4

Tectosilicates (Si02)o Feldspars

*Orthoclase KAISi)O.Albite NaAISi)O.Anorthite CaAI2Si2Os

Si02GroupQuartz Si02Tridymite Si02Cristobalite Si02

ZeolitesAnaleime NaAISi20 6-H2O

FeldspathoidsNephelene (Na.K IAISi04

t After Allen & Fanning (1983).t Klein & Hurlbut (1993).

II. PHYLLOSILlCATE MINERALS lN SOILS

Phyllosilicate minerais strongly inf1uence both the chemical and physicalproperties of soils because of their generally small particle sizes. high surfaceareas, and unique cation exchange properties. A clear understanding ofthe phyl­losilicate minerais is central to understanding soil c1ay mineralogy and many en­vironmental processes.

The discovery in the early 1920s that clays are crystalline was the key to un­derstanding many ofthe properties ofsoils and c1ays. The crystal structures ofthephyllosilicates define the different mineral species and account for many oftheirunique properties. The structural schemes of these minerais are discussed below.The purpose is to give an overview ofthe most commonly occurring structures andthe major properties ofeach major mineral group. Subsequent chapters will add de­tails, ifnecessary. Emphasis will be placed on developing an understanding ofthevarious structural representations encoontered in the literature.

8

A. Basic Structural Concepts

SCHULZE

The following discussion will treat atoms as rigid spheres. This is an over­simplified but convenient model for developing the key concepts ofmineral struc­tures.

I. Closest Packing ofSpheres

The most efficient way of packing spheres in one plane is called hexagonalc10sest packing (Plate 1-2). Each sphere touches six of its neighbors and a hexa­gon results ifthe centers ofthese six neighbors are connected. There are two typesof voids between the spheres: A voids (where the triangular outline points down­ward in the illustration) and B voids (where the triangular outline points upward).Ifa second plane ofspheres is placed on top ofthe first (Plate 1-3), and ifthe firstsphere ofthe second plane is placed in the dimple formed by the three spheres sur­rounding an A void, then ali the remaining spheres of the second plane must beplaced over A voids as well. The spheres illustrated in Plate )-3 represent 0 2-, OH-,or F- in the phyllosilicate structures, while positively charged cations occupy thespace between these negatively charged anions.

2. Tetrahedra and Octahedra

Another examination ofPlate )-3 shows that there are two different types ofsites into which cations can be placed. The A sites, formed over the A voids ofthelower plane of spheres, are called tetrahedral sites because the four 0 2- ions sur­rounding this site form the apices ofa regular tetrahedron. The B sites, formed overthe B voids of the lower plane of spheres, are called octahedral sites because thesix 0 2- ions surrounding each ofthese sites form the apices ofa regular octahedron(Plate \-3). This illustrates that both octahedral and tetrahedral sites are a conse­quence ofthe closest packing oftwo planes ofspheres.

A cation in a tetrahedral site is in fourfold (or tetrahedral) coordination be­cause it is bonded to four 0 2- ions whose centers define the apices (corners) of atetrahedron. Likewise, a cation in an octahedral site is in sixfold (or octahedral) co­ordination because it is bonded to six 0 2- ions whose centers define the apices ofan octahedron. Tetrahedrally and octahedrally coordinated cations-tetrahedra andoctahedra for short-are common structural elements in many mineral structures.The approach ofdescribing mineral structures as assemblages ofpolyhedra is usedwidely and, once mastered, provides an efficient way to understand even complexmineral structures.

3. Representing Tetrahedra and Octabedra

Octahedra and tetrahedra are commonly represented using three differenttypes ofmodels. Each model portrays the sarne concept, but different representa­tions are necessary to show various structural features. Plate )-3 shows two viewsofa tetrahedron and ofan octahedron drawn in these three different ways.

The sphere-packing model gives an impression ofthe space occupied by theatoms in a structure. The sphere-packing model is particularly useful for visualiz-

INTRODUCTION TO SOIL MINERALOGY

ing the shape ofthe external surface ofa mineral, but it does not allow one to eas­ily see the interior ofthe structure. The ball-and-stick model shows the bonds withinthe crystal since the bonds are represented as sticks connecting the balls. The tetra­hedra and octahedra are more difficult to visualize, but the interior ofthe structureis easier to see because the atoms are drawn much smaller than in the close-packedmodel. The polyhedral models give the best impression oftetrahedra and octahe­dra since the atoms are represented only as points in space or as small spheres. Thecenters ofthe 0 2- ions are connected by lines to form the edges oftetrahedra andoctahedra. Thus, each apex represents the position ofan 0 2- ion, while a cation re­sides in the center ofthe tetrahedron or octahedron. ln general, only those tetrahe­dra or octahedra actually occupied by cations are represented by a solid polyhedron;sites that are not occupied by cations are not represented by solid polyhedra. Oneshould use care in interpreting polyhedral models, since it is easy to get the im­pression that a structure contains a large amount ofopen space or pores. The sarnestructure drawn as a c1ose-packed model will show that much ofthe space is actu­ally occupied by the large O atoms. Other ways ofrepresenting mineral structuresare encountered as well. Sometimes, polyhedral models are drawn without spheresto represent the oxygens and hydroxyls at the corners ofthe polyhedra (see, for ex­ample, Plate 1-9later in this chapter), while other models show only sticks to rep­resent bonds, but no spheres to represent the atoms.

There is some ambiguity associated with each type of representation, and itis important to understand the correspondence between, and limitations of, each ofthe different structural representations. Thus, it is important to use as many differ­ent leaming aids as possible. Computer graphics programs that allow the viewer torotate and manipulate structural models on the computer screen offer significantadvantages over static drawings. Physical models should be studied as well, iftheyare available. Far from being child's play, building your own models is one ofthebest ways to learn about mineral structures. Models can be built with inexpensiveballs and some glue, or with paper polyhedra (see Moore & Reynolds [1997] fortemplates).

4. Ionic Radii and Radius Ratios

Whether or not a given atom can occupy a tetrahedral or an octahedral sitedepends both on the atom's charge and size. The size ofatoms and ions varies de­pending on their atomic number, ionization state, and coordination number (TableI~).

The central void within a tetrahedron is smaller than the one within an octa­hedron. ln other words, the largest sphere that can be placed into a tetrahedral sitewithout pushing the 0 2- ions apart is smaller than the sphere that can be placed intoan octahedral site without pushing. Simple calculations (see, for example, Klein &Hurlbut, 1993) illustrate that the maximum radius of a sphere that will just fit inthe tetrahedral site is 0.225 times the radii ofthe four surrounding 0 2- ions. For anoctahedral site, the maximum radius is 0.414 times the radíi ofthe six surrounding0 2- ions. The next larger type of site (not illustrated) is a cubic site consisting ofeight 0 2- ions arranged at the corners ofa cube. For a cubic site, the maximum ra­dius is 0.732 times the radii ofthe eight surrounding 0 2- ions. Assurning a radius

"Table 1--4. lonic radii and coordination ofcommon elements in phyllosilicate minerais.

SCHULZE

lon Jonic radiust

nm

fP-- 0.140(6)F"" 0.133(6)a- 0.181(6)Si'" 0.026(4)A13+ 0.054(6)Fe3+ 0.065(6)

~+ 0.072(6)0.061(6)

Fe2+ 0.078(6)Mn2+ 0.083(6)Na+ 0.118(8)cr+ 0.112(8)K+ 0.151(8)&2+ 0.142(8)Rb' 0.161(8)

Typical coordination

TetrahedralTetrahedral and octahedralTetrahedral and octahedralOctahedralOctahedralOctahedralOctahedralCubic and dodecahedralCubic and dodecahedralCubic and dodecahedralCubic and dodecahedralCubic and dodecahedral

t lonic radii vary with coordination number. Radii are for ions wilh lhe coordination nurnbers listed inparentheses (fram Klein & Hurlbut. 1993. p. 188).

of 0.140 nm for the 0 2- ions, the limiting radius for a tetrahedral site is 0.032 nm(0.140 nm x 0.225 = 0.032 nm), for an octahedral site it is 0.058 nm, and for a cubicsite it is 0.102 nm. The Si4+ion has a radius <0.032 nm (Table 1--4). Thus, Si4

+ al­ways occurs in tetrahedral rather than octahedral sites. The Mg2

T ion has a radiusthat is >0.032 so Mg2

T is unlikely to occur in tetrahedral coordination, but since itsradius is between 0.058 and 0.102 nm, it fits easily in an octahedral site. The A13+ion, with a radius near to the limiting radius ofO.032 nm, can fit in both tetrahe­dral and octahedral sites. Cations with radii >-0.102 nm tend to occur in cubic, do­decahedral (12-fold) or higher coordination sites rather than in octahedral or tetra­hedral sites (Table 1--4).

5. Tetrabedral and Octabedral Sbeets

Tetrahedra arranged into sheets are common to the structures of ali phyl­losilicates. Octahedra arranged into sheets are present in the structures of phyl­losilicates and in some hydroxide minerais as well. The structure oftetrahedral andoctahedral sheets will be discussed first. Then it will be shown how different com­binations of tetrahedral and octahedral sheets give rise to the different clay min­eral structures. Most ofthe diagrams in this chapter show ideal crystal structuresbased on the approximate c10sest packing of spheres.

a. Tbe Octabedral Sbeet. Consider two planes of spheres representing hy­droxyl (OH-) ions in hexagonal closest packing. (The W takes up very little spaceand the OH- can be considered a sphere of roughly the sarne size as an 0 2- ion).There are two ways offilling the octahedral sites depending on the valence ofthecation. A divalent cation, such as Mg2" can be placed into each octahedral site to

INTRODUCTlON TO SOIL MINERALOGY 11

obtain the network of octahedra illustrated in Plate l-4a. This arrangement iscalled trioctahedral because three of every three octahedral sites are occupied bya cation. This gives a formula ofMg3(OH)6 or Mg(OHh for the whole sheet; eachMg2+is surrounded by six OH-s, but since each OH- is shared equally among threedifferent Mg2+atoms, each OH- contributes only one-third of its negative chargeto each Mg2+, giving the formula Mg(OHh- The sheet is electrically neutral becausethe charge is balanced within the sheet.

The other possibility is to place a trivalent cation, such as AJ3+, into the oc­tahedral sites. To preserve charge neutrality, only two trivalent cations are neededcompared with three divalent cations, giving the formula AI2(OH)6 or AI(OHh.Thus, only two ofevery three possible octahedral sites are filled. This arrangementis called dioctahedral and is illustrated in Plate l-4b. The unit formula for the dioc­tahedral sheet can be deduced by following the sarne reasoning used above. EachAJ3+ is surrounded by six OH-s, but since each OH- is shared equally between twodifferent A13+ ions (not three as for Mg2+), each OH- contributes one-halfof its neg­ative charge to each AP+, giving the formula AI(OHh. Again the charge is balancedand the sheet is electrically neutral.

Note how the trioctahedral and dioctahedral sheets appear in the differenttypes ofrepresentations. The sphere-packing model shows that both the trioctahe­dral and dioctahedral sheets contain the sarne dense packing ofOH-s. The pattemofunoccupied 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 by Hbonds. make up the structures of gibbsite (see Plate 1-7 below) and brucite[Mg(OHhl. Gibbsite and brucite differ in that gibbsite contains AI in the octahe­dral sites and is dioctahedral, while brucite contains Mg in the octahedral sites andis trioctahedral. The structures of gibbsite and brucite are the simplest in a seriesof 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 assuming thatthe large O atoms were hard spheres of identical size c10sely packed in three di­mensions. Thus, we assumed that the O atoms were our basic structural units andthat the cations merely served to balance the negative charge ofthe oxygens. Thisapproach works for some of the oxide and hydroxide minerais, but in many min­eraI structures the O atoms are not as c10sely packed as strict geometric packingwould dictate. A more general way ofdescribing mineral structures is the polyhe­dral approach. ln the polyhedral approach. the coordination polyhedra. primarilyoctahedra and tetrahedra in the phyllosilicates, are considered the basic structuralbuilding blocks. Thus. we can consider the octahedral sheets illustrated in Plates1-4a and 1-4b as an assemblage of octahedra in which adjacent octahedra sharetwo oxygens with one another. ln other words, the octahedra share edges with oneanother. ln general, coordination po1yhedra can share comers (one shared oxygen),edges (two shared oxygens), or faces (three or more shared oxygens) with adjacentpolyhedra. The octahedral sheet could be described using either the sphere pack­ing or polyhedral approaches. The tetrahedral sheet described below can only bedescribed using the polyhedral approach because the oxygens in the tetrahedral sheetare not c1ose-packed as they are in the octahedral sheet.

12 SCHULZE

b. The Tetrahedral Sheet. The tetrahedral sheet consists ofSi04 tetrahedraarranged such that three O~- ions ofeach tetrahedron are shared with the three near­est neighbor tetrahedra (Plate 1-5). These shared O~- ions are ali in the sarne planeand are referred to as basaloxygens. Note that two adjacent tetrahedra share onlyone O~- between them. (The tetrahedra share apices or comers.) The fourth O~- ionofeach tetrahedron is not shared with another Si04 tetrahedron and is free to bondto other polyhedral elements. These unshared O~- ions are referred to as apical oxy­gens. Since each basal oxygen contributes a formal charge of -I to each Si4+ ion,the addition ofW ions to the apical oxygens to form hydroxyls should result in anelectrically neutral tetrahedral sheet. Such individual tetrahedral sheets. althoughsometimes postulated to occur as transient weathering products in aqueous solu­tion, do not stack to form stable mineral structures on their own. Thus. tetrahedralsheets only occur in combination with octahedral sheets as described below.

The upper portion of Plate 1-5 shows ali of the apical oxygens pointing inthe sarne direction, namely. out ofthe plane ofthe paper towards the reader. Struc­tures in which the apical oxygens of a single sheet ali point in the sarne directionare the most common. but structures also occur in which the apical oxygens pointaltemately in opposite directions. MineraIs containing this sheet-like arrangementofSi04 tetrahedra are called phy/losilicates', sheet silicates, or layer silicates.

B. Phyllosilicate MineraIs Common in Soils

Phyllosilicates are divided into two groups, I: I and 2: I type minerais. on thebasis ofthe number oftetrahedral and octahedral sheets in the layer structure.

1. The I: I Type MineraIs

a. The I: I Layer Structure. The I: I layer structure consists ofa unit madeup ofone octahedral and one tetrahedral sheet, with the apical O~- ions ofthe tetra­hedral sheets being shared with (and part of) the octahedral sheet (Plate 1-6a and1-6b). There are three planes ofanions (Plate 1-6b). One plane consists ofthe basal0 2- ions of the tetrahedral sheet, the second consists of O~- ions common to boththe tetrahedral and octahedral sheets plus OH- belonging to the octahedral sheet,and the third consists only ofOH- belonging to the octahedral sheet.~

The edge view of the I: I layer in Plate 1-6b is one of two possible depic­tions often encountered in the literature. The other view is obtained by rotating thestructure around an axis normal to the sheet and viewing the structure along a dif­ferent crystallographic axis (Plate 1-6c).

1 The prefix phy/lo-. derived from the Greek word phy/loll. meaning leaf. should not be eonfusedwilh the prefixphy/o-. derived from the Greek wordphy/olI. meaning tribe. Phy/lo- is pronouneed 'fiI-õ.while phylo- is pronouneed 'fi-lã. Thus. the proper pronuation for phyllosilieate is 'fil-õ-sil-a-kãl.

1 The usage ofthe terms plane. sheet. and /ayer follows lhe recommendations ofthe AIPEA Nomen­c1ature Committee. They refer to inereasingly thieker arrangements sueh as a single p/alie of atoms. atetrahedral or oetahedral sheet. and aI: I or 2: I /ayer. A sheet is a eombination of planes. and a layer isa eombination of sheets. Layers ean be separated by various illler/ayer materiais. sueh as eations. hy­drated eations. organie moleeules. and hydroxide octahedral groups and sheets. The total assembly ofa layer plus interlayer material is ealled a unil structure (AIPEA Nomenclature Committee, 1980).

Ii'iTROO eTlo TO OILMI ERALOGY 13

Plme I I. The strueturc ofhalite ( ael).

"A" voids "B" voids

Plate I ' Ilexagonal close!>l packing of sphcre!> in a plane.

Face Edge

"A" site

Ball-and-StickModel

Sphere-PackingModel

PolyhedralModel

Tetrahedron Octahedron Tetrahedron Octahedron

Plate I 3. Octahedra and tetrahedra as a con!>equence oft\\o planes ofclose-packed spheres and threeways of repre!>enling octahedra and telrahedra.

I~

a The Octahedral Sheet (trioctahedral)

H LZE

b The Octahedral Sheet (dioctahedral)

Plate I . The octahedral heet. (a) A tnoctahedral heet. (b) A dioctahedral sheet.

I TROD T10N TO O/L .\J1. ERALOGY

16

a

b

HULZE

PLANES OF 10NS SHEETS, LAYERS

basal O's =-oÃ,...........",...".-.,.,.foooII""""~~.,.."""" ..........",...~ } 1tetrahedral cations r tetrahedralOH's &apical O's _ } sheet 1'1

octahedral cations _ octahedral laYersheet

OH's-

basalO's } }tetrahedral cations f"" letrahedral

~~~~~f~;~~: } oct:~:~1 2:1OH's &apical O's _ } sheel layer

tetrahedral cations 't.. tetrahedralbasal O's _ rl~....~"'"",,~--I'~""'~~r...~""''''>''l'''''H'''''.... sheet

cAlternateedge view

Plate I 6. (a) Oblique view of the I: I layer slrueture illuslrating the relationshlp belween the tetrahe­dral and oelahedral sheet. ote that adjaeent apieal oxygens ofthe tetrahedral sheet (arrows a) alsodefine edge;, of oelahedra in the oetahedral sheet. Arrows b point to Olls lhat lie direetly in the een­ler ofthc hexagonal rings oflelrahedra. although they appear off-eenter in this obliquc \ iew. (b) Edgevie\\ ofthe I: I and 2: I layer struclures illu;,lrating phyllosilicate nomenclature. (c) An aItemate edge\ ic,," that arises \\ hen the 2: I structure above is rotated n mlal to lhe plane of the layer.

I 'TRODUCTIO 'TO OIL I\IINERALOGY 17

Gibbsite

0.46nm [(4.6 A)

Kaolinite

0.72 nm [ ~~~~~~~~(7.2 A) ~1Itl1J'll1llAl~~'-A~IIIAIP_ Si~+

_AI3+

CWltAlIfliI"~~IIAIPwl1JP _ Si~+

_AI3+

Pyrophyllite""'""""'V''V''........................"'V'v-..,.-".

'1Il1lflillAlPt..~1Ifl~AIP _ Si~+

_AI3+

',;; A~ [~~~~~~~~~ '1Ií<liIPI......AI~..~

~1rlW1~íttI~1rlW1~t>_ Si~+,AI3+

_AI3+

Mica (muscovite)""~"'V''V'' .....................''''V"'V"''........

1.00 nm(10.oA)

Plate I 7. Structural schemc of soilmincrals bascd on octahcdral and tetrahcdral shects.

18

>1.8 nm(>18A)

CH LZE

VermiculiteQ8W'-.AlAiAII"'~IIIIÍltIIA~ Si4+,AI3+

___ A13+

Y 'Y 'Y. 'Y. ~~~ l{;]l •• ;]l ••.... .... -... .... ... ...

Smectite

• --- H20

exch. Ca2+,

Mg2+, ele.

Hydroxy-interlayeredVermiculite and Smectite

•• A13+ poly..

• ;- exch. Ca2+,

Mg2+, etc.

Chlorite

1.4 nml(14 A)

(11 r 1 r 1 1 1 I ] I )L 11M ~[ .lI II ~1 JI( I j, I( ~I ~ l( jJ j

( )

:.a :aI I II I .I:

( )j .ilI .liI.UI 1 l'l..:a .iII lT li

:I! ilL,,),. }, jJ j( ." ". ,) ,j }

Plate I 7. ontinued.

.:.:'. :. ' .:.:.:.:,.:

INTROD eTlo TO OIL 11NERALOGY

Palygorskite

19

f] IJ ~~,iI I r] 1J.1 III I II ~r I.1 li j II .( }(Atl

I ] r- I l:~

ti .l1 MI 11.[ }( j II,I 11

r1 I

.li ti .M .lf.ll

.1 ~(Ji .r }I }.11;.r,J ,1d~ 1 T LiJ TJ

'1It.,::l~~~~~~AlAIIP__ Si4+

__ Mg2+,AI3+'tl"IIII'tI"!:l.

H20

_-M.,---H20(not shown)

~1AiIA~~~;:wtIpfI.

Clinoptilolite (a zeolite)fl"'1I~~~~"""~~

Imogolite

1.0 nm(10 A)

H20(nol shown)

Plate 1-8. lructural models ofrepresentatives ofthree other alumino ilicate mineral groups thal occurfrequently in soils. Ali three are drawn to the same scale.

Structural Details Common to Phyllosilicate Minerais

Tetrahedral Sheet Octahedral Sheet

N00

occupiedoctahedra

smaller lhanunoccupied

shared edgesshorter lhan

unshared

r--corrugatedbasal plane

L-Plate I 9. Struetural detaib of phyllosilieates as illuslralcd by lhe octahcdral and tctrahcdral shects of muscovite. Figures \Vcre prcparcd from the single-cryslal structural

refinel11cnt data of ROlhbaucr ( 1971).

n:tc::r­NC"'l

INTRODUCTlON TO SOIL MINERALOGY 21

b. Kaolinite. The structure of kaolinite consists of I: I layers stacked oneabove the other. Kaolinite is dioctahedral and contains AP+ in the octahedral sitesand Si4+in the tetrahedral sites (Plate 1-7). The I: I layer is electrical1y neutral andadjacent layers are held together by hydrogen bonding between the basal oxygensof the tetrahedral sheet and the hydroxyls of the surface plane of the adjacent oc­tahedral sheet.

Kaolinite is a common mineral in soils and is the most common member ofthis subgroup. It tends to be particularly abundant in more weathered soils such asUltisols and Oxisols. Kaolinites have very little isomorphous substitution in eitherthe tetrahedral or octahedral sheets and most kaolinites are close to the ideal for­mula AI~Si~05(OH)4' The I: 1 layer has little or no permanent charge because ofthelow amount ofsubstitution. Consequently. cation exchange capacities and surfaceareas are typical1y low. Thus. soils high in kaolinite are generally less fertile thansoils in which 2: I clay minerais dominate.

Kaolinite can form in soils from AI and Si released by the weathering ofpri­mary and other secondary minerais. For example. feldspars often weather to kaoli­nite in soils formed from igneous rocks. Kaolinite can also be inherited from thesoil parent material (i.e.• clayey sediments). Kaolinite mined from geologic depositsis used widely in a variety of industrial applications such as fil1ers for plastics, inceramics, and as a fil1er and coating for paper.

e. Halloysite. Hal10ysite has ai: 1 layer structure similar to kaolinite exceptthat the I: 1 layers are separated by a layer of H~O molecules when fully hydrated(Plate 1-7). Most clay silicates occur as thin plates, but halloysite often occurs astubular or spherical particles.

Hal10ysite is usually found in soils formed from voleanic deposits. particu­larly voleanic ash and glass. It is a common clay mineral in the Andisol soil order.Hal10ysite forms early in the weathering process, but it is generally less stable thankaolinite and gives way to kaolinite with time.

2. The 2: 1 Type Minerais

ln contrast to the I: I minerais, which are mostly represented in soils by ontytwo major minerais, the 2: I minerais are structurally more diverse and are repre­sented by severaI mineral species.

a. The 2: 1 Layer Strueture. The 2: I layer structure consists oftwo tetrahe­dral sheets with one bound to each side ofan octahedral sheet (Plate 1-6b). Thereare four planes of anions. The outer two consist of the basal oxygens (Os) of thetwo tetrahedral sheets, while the two inner planes consist ofoxygens common tolhe octahedral sheet and one ofthe tetrahedral sheets, plus the hydroxyls (OHs) be­longing to the octahedral sheet.

b. Tale and Pyrophyllite. The simple structures oftale and pyrophyllite aregood starting points for discussing 2: 1 structures. 80th minerais consist of2: I lay­ers stacked one above the other. Tale has Mg~+ in the octahedral sites and is trioc­tahedral. while pyrophyllite (Plate 1-7) has AIJ+ in the octahedral sites and is dioc­tahedral. The tetrahedral sheets of both minerais contain only Si4+. giving idealformulas ofMgJSi40 IO(OHh for tale and AI~Si401O(OHh for pyrophyl1ite. ln each

22 SCHULZE

case the charge is balanced within the 2: I layer, making the layer electrically neu­tral. Adjacent 2: I layers are held together only by weak van der Waals forces.

Tale and pyrophyllite occur only rarely in soils, usually only when they areinherited from low-grade metamorphic rocks. Tale and pyrophyllite are used in­dustrially as ingredients in paints, ceramics, plastics, paper, and cosmetics. The oc­currence oftale in river and estuarial sediments when there is no geological sourceis a reflection of industrial activity in the watershed.

c. Micas. Mica minerais have the 2: I layer structure described for tale andpyrophyllite but with two important differences. First, instead ofhaving only Si4+in the tetrahedral sites, one-fourth ofthe tetrahedral sites are occupied by AI3+. Be­cause ofthis substitution, there is an excess ofone negative charge per formula unitin the 2: I layer. Second, this excess negative charge is balanced by monovalentcations, commonly K+, that occupy interlayer sites between two 2: I layers. Thisgives an ideal formula ofKAI2(AISi3)OIO(OHh for a mica mineral with AI in theoctahedral sites.

Just as in tale and pyrophyllite, the octahedral sheet can be either dioctahe­dral (Plate 1-7) or trioctahedral. There are several different mica species becauseFe2+and Fe3+can substitute for Mg2+and AI'+ in the octahedral sheet and Na+ andCa2+can substitute for K+ in the interlayer.

Mica in soils is usually inherited from the parent rock and is likely to occurin soils derived from various igneous and metamorphic rocks, as well as from sed­iments derived from these rocks. Muscovite, biotite, and phlogopite are the threemost common mica group minerais in rocks, and consequently in soils. Allthreecontain K in the interlayer (Table 1-3), but they differ in the composition ofthe oc­tahedral sheet and whether they are di- or trioctahedral.

Mica in the clay fraction of soils and sediments differs somewhat from themacroscopic muscovite mica it most closely resembles, This clay-size mica isoften referred to as i/lire. Glallconite is another mica mineral that is similar to il­lite, but it contains more Fe and less AI in its octahedral sheet than illite.

Micas weather to other minerais, particularly to vermiculites and smectitesand the K+ released during weathering is an important source of K for plants. As arule, the dioctahedral micas, such as muscovite, are more resistant to weatheringthan trioctahedral micas. Thus, muscovite tends to be the most common mica min­erai found in soils. Micas are used commercially in paints and cosmetics.

d. Vermiculites. Vermiculite has a 2: 1 layer structure as described for mica,but instead ofhaving a layer charge of-I per formula unit and K+ in interlayer po­sitions, vermiculite has a layer charge of 0.9 to 0.6 and contains exchangeablecations, primarily Ca and Mg, in the interlayer (Plate 1-7). The high charge per for­mula unit gives vermiculites a high cation exchange capacity and causes them tohave 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 ofmicas and chlorites. The weathering ofmicas to vermiculite (or smectite)is believed to occur by replacement of K+ in the interlayer sites with hydrated ex­changeable cations. The integrity ofthe 2: I layer is preserved, but there is a reductionin the layer charge. Vermiculite does not swell as extensively as smectite, and this

INTRODUCTION TO SOIL MINERALOGV 13

is illustrated in Plate 1-7 by the presence of only two planes ofwater moleculessurrounding the hydrated cations in the interlayer space. Some commercial uses ofvermiculites include horticultural poning media and thermal insulating materiais.

e. Smectites. The smectite group consists ofminerals with the 2: I structurealready discussed for mica and vermiculite, but with a stilllower charge per for­mula weight, namely 0.6 to 0.25. As in vermiculite, the interlayer contains ex­changeable cations (Plate 1-7).

As for the micas and vermiculites, dioctahedral smectites are more commonin soils than trioctahedral smectites. The most common smectite minerais range incomposition between the three endmembers: montmorillonite, beidellite, and non­tronite. Ali are dioctahedral, but they differ in the composition of the tetrahedraland octahedral sheets. Smectites do not fix KTas readily as do vermiculites becausesmectites have a lower layer charge, but smectites swell more extensively than ver­miculites. This is illustrated in Plate 1-7 by the larger spacing between the 2: I lay­ers.

Smectites are important minerais in temperate region soils because oftheirhigh surface area and their adsorptive properties. Smectites shrink upon drying andswell upon wening. This shrink-swell behavior is most pronounced in the Vertisolorder and in vertic subgroups ofother soil orders. The shrink-swell properties leadto cracking and shifting problems when houses, roads, and other structures are builton smectitic soils. Smectites mined from almost pure geologic deposits are usedwidely as catalysts, adsorbents for spills, cat liner, sealants for ponds and abandonedwells, in drilling tluids for oi! wells, and in liners for landfills.

f. Chlorites. Like mica. chlorite minerais have a 2: I layer structure with anexcess of negative charge. ln contrast to mica. however, the excess charge is bal­anced by a positively charged interlayer hydroxide sheet (Plate 1-7), rather thanK+. The interlayer hydroxide sheet is an octahedral sheet as illustrated in Plate 1-4and can be either di- or trioctahedral. Instead ofbeing e1ectrically neutral as in bruciteor gibbsite, the hydroxide sheet has a positive charge caused by substitution ofhighervalence cations for lower valence ones, for example, Mg2AI(OH)6+' Either octa­hedral sheet-the one that is part of the 2: I layer or the interlayer hydroxidesheet---.çan be di- or trioctahedral, and can contain Mg2+, Fe2+, Mn2+, Ni2+, AP+,Fe3+, and cr3+, giving a large number of different mineral species.

Chlorite minerais in soils are often primary minerais inherited from eithermetamorphic or igneous rocks. They may also be inherited from sedimentary rockssuch as shales, or from hydrothermally altered sediments. Chlorites are rather in­frequent minerais in soils and when they do occur they generally make up only asmall amount ofthe soil. Chlorites weather to form vermiculite and smectite andthe ease with which they break down makes them sensitive indicators ofweathering.

g. Hydroxy-Interlayered Vermiculite and Smectite. Hydroxy-interlay­ered vermiculite and smectite can be considered a solid solution with vermiculiteor smectite as one endmember and chlorite as the other. Hydroxy-interlayeredminerais form as AP+ released during weathering hydrolyzes and polymerizes toform large polycations with a postulated formula ofAI6(OH)fs+ (or similar) in theinterlayers of vermiculite and smectite. These polycations balance some of the

SCHULZE

charge ofthe 2: 1 layer. The combination of a 2: 1 layer with hydroxy AI in the in­terlayer gives a structure similar to that of chlorite (Plate 1-7). Thus. these miner­aIs are also called secondary ch/orites. The degree offilling ofthe interlayer withhydroxy AI can vary from none to almost complete. with properties ofthe c1ay vary­ing accordingly. The interlayer hydroxy AI is not exchangeable; therefore. it low­ers the cation exchange capacity ofsmectite or vermiculite almost Iinearly as a func­tion ofthe amount of AI adsorbed in the interlayer.

Interlayer hydroxy AI prevents smectite from shrinking and swelling as itwould normally. ln vermiculite. it reduces K+ fixation by lowering the exchangecapacity and by preventing the interlayer from collapsing around the K+. The pos­itively charged hydroxy interlayers also provide potential sites for anion adsorp­tion.

Hydroxy-interlayered vermiculite and smectite are most common in Alfisolsand Ultisols. Within a given profile they tend to be most abundant near the soil sur­face. Hydroxy-interlayered smectites produced industrially are used as catalysts inthe chemical industry.

h. InterstratificatioD in Phyllosilicates. Phyllosilicates in soils do not alwaysoccur as discrete particles ofmica. vermiculite. smectite. chlorite. or kaolinite. Forexample. instead of being made up of a stack of identical 2: I vermiculite layers.one physically discrete particle may consist of a mixture Df both mica and vermi­culite layers instead. Such minerais are referred to as mixed-/ayer or interstratifiedminera/s.

Ditferent types ofinterstratified mineraIs have been identified. Two-compo­nent systems include: mica-vermiculite. mica-smectite. mica-ehlorite. chlo­rite-vermiculite. chlorite-smectite. chlorite-swelling chlorite. and kaolinite-smec­tite. Three-component mixed layer systems can also occur. The sequence oflayerscan be either regular Dr random. A regularly interstratified mineral consisting oftwo types of layers denoted by A and B could have a sequence Iike ABABAB...,or ABBABBABB...• Dr any other repeating sequence. ln a randomly interstratifiedmineral the sequence oflayers is random. for example. ABBABAABBAAA. ... Ran­dom interstratification oflayer-silicates is more common in soils than regular in­terstratification, though regular interstratification. especially in weathering micas,is not rare.

Partial removal of interlayer K from micas or of interlayer hydroxide fromchlorite is one way that interstratified minerais can form in soils. Other possibili­ties include (i) fixation of adsorbed K+ by some vermiculite layers to give mica­like layers and (ii) the formation ofhydroxide interlayers to produce chlorite-likelayers.

i. Palygorskite and Sepiolite. Palygorskite and sepiolite are consideredphyllosilicates. but are distinct structurally from the typical I: I and 2: I layer struc­tures. 80th minerais have continuous tetrahedral sheets. but adjacent bands oftetrahedra within one tetrahedral sheet point in two opposite directions rather thanin one direction as in the I: I and 2: I structures. The result is a structure that can bedescribed as ribbons of 2: I layers joined at their edges. as illustrated in Plate 1-8.Water molecules occur in the spaces between the ribbons. The 2: I ribbons are widerin sepioJite than in palygorskite.

INTRODUCTION TO SOIL MINERALOGY

Palygorskite and sepiolite are often found in soils ofarid and semiarid envi­ronments. Both minerais have a fibrous morphology in contrast to the platy mor­phology of most I: I and 2: I minerais. Because of their fibrous morphology, sus­pensions ofthese clays can fonn thick gels even at low solid concentrations. Thus,palygorskite and sepiolite are used industrially as gelling agents to keep othersolids in suspension.

C. Structural Details of Pbyllosilicates

Plates I-I through 1-8 illustrate ideal mineral structures based on the regu­lar close packing of spheres. As a consequence, ali of the octahedra in Plate 1-4are exactly the sarne size and ali are regular (undistorted) solids. Likewise, in Plate1-5, ali ofthe tetrahedra are regular solids, the six-member tetrahedral rings fonna regular hexagon, and the basal oxygens alllie in exactly the same plane. Crys­tallographic structure refinements of real phyllosilicate minerais show distortionsfrom the ideal structures illustrated up until now in this chapter. These distortionsare similar for ali ofthe phyllosilicate minerais; although they vary somewhat inmagnitude.

There are two reasons why real minerais deviate from these ideal structures.First, atoms are not hard spheres whose closest approach is detennined solely bythe sum oftheir radii. Second, ifone does try to join octahedral and tetrahedral sheetsbuilt ofhard spheres like those illustrated in Plates 1-4 and 1-5, geometry alonedictates that they will not fit together ifthe spheres representing the O atoms toucheach other in both the octahedral and tetrahedral sheets. The reason becomes ap­parent by examining the sphere-packing model ofthe tetrahedral sheet in Plate l-S.Since the apical oxygens are common to both the tetrahedral and octahedral sheets,two adjacent apical oxygens in the tetrahedral sheet also define the edge ofan oe­tahedron in the octahedral sheet (Plate l--6a). Note that, although basal oxygenstouch neighboring basal oxygens, apical oxygens do not touch neighboring apicaloxygens (Plate I-S). Thus, the octahedral sheets illustrated in Plate 1-4 cannot shareapical oxygens with the tetrahedral sheet illustrated in Plate 1-5 because the tetra­hedral sheet is larger than the octahedral sheet. This problem was overcome in Plates1--6 to 1-8 by increasing the 0-0 distance in the octahedral sheet to match the 0-0distance ofthe apical oxygens ofthe tetrahedral sheet. Thus, the octahedra remainundistorted, but they are larger than they would be ifthe spheres representing oxy­gen touched.

ln the real phyllosilicate structures, this mismatch between the octahedral andtetrahedral sheets is compensated, not so much by an expansion ofthe octahedralsheet, but more so by distortions in the tetrahedral sheet. The net effect is that, thetetrahedra are rotated from the ideal arrangement shown in Plate I-S such that thehexagonal shapes outlined by the six-membered tetrahedral rings in Plate I-S be­come the ditrigonal shapes illustrated in Plate 1-9. ln addition, the tetrahedra aretilted such that the basal oxygens no longer lie in the sarne plane, resulting in a cor­rugated basal plane in which some rows of atoms lie slightly below and some lieslightly above the average plane ofthe basal oxygens. These corrugations are bestseen in the lowennost edge view in Plate 1-9 where spheres representing the Oatoms have not been drawn. The implications ofthese distortions are that molecules

26 SCHULZE

interacting with phyllosilicate surfaces "see", not the ideal surface iIIustrated in Plate1-6a, but the more distorted surface illustrated in Plate 1-9.

The octahedral sheet is distorted from the ideal structure as well. First, be­cause ofthe attraction ofeach central octahedral cation to its six surrounding oxy­gens, octahedra occupied by cations are smaller than octahedra in which the cen­traI sites are unoccupied (Plate 1-9). ln addition, when two octahedra share twoadjacent oxygens (the octahedra share edges), the 0-0 distance along this sharededge is shorter than the 0-0 distance along an edge that is not shared with an ad­jacent oeeupied octahedron (Plate 1-9). Thus, the oetahedra are distorted from theideal Euclidean solids in very predietable ways. ln general, trioctahedral phyl­losilieates show less struetural distortion and more closely resemble the ideal strue­tures than do the dioctahedral phyllosilieates.

The structural distortions just discussed probably aecount for relatively minordifferences in the properties ofthe phyllosilicates, with the major differences de­terrnined mainly by whether the mineral has a I: I or 2: I layer structure and by thelayer charge. On the other hand, since dioctahedral phyllosilicates are more com­mon in soils than trioctahedral phyllosilicates, water, ions, and molecules interaetpredominately with surfaces like those illustrated in Plate 1-9. rather than surfaceswith the more regular arrangement illustrated in Plates 1---4 and l--6a.

III. OTHER ALUMINOSILICATE MINERALS COMMON lN SOILS

A. Zeolites

Zeolites are a large group ofminerals that consist structurally ofSi04 tetra­hedra arranged in ways that result in large amounts of pore space within the crys­tais (Plate 1-8). Aluminum substitutes for Si in the tetrahedral sites ando as a re­sult. the (Si,AI)04 framework has a net negative charge. The charge is balanced byeations that reside in the channels and pores, along with water molecules. Becausethe cations are exchangeable, zeolites have cation exchange properties similar tothe phyllosilicates. but because the tetrahedral framework ofthe zeolites is rigidand the size ofthe pores is fixed, small cations can move into and out ofthe poresfreely. while larger cations are excluded. Thus, zeolites are often referred to as mol­ecular sieves because oftheir very selective cation exchange properties.

Zeolites are relatively rare in soils because they weather easily in humid re­gions, but they do occur in some soils in arid regions. Zeolites are widely used inindustry and agriculture. for example, as catalysts and adsorbents in the chemicalindustry. as additives in animal feeds. and as ingredients in laundry detergents. Thus,zeolites may be introduced into soils and sediments in unexpected ways.

B. Allophane and Imogolite

The aluminosilicate mineraIs discussed above have three-dimensional crys­tal structures with atoms packed together in a more or less regular manner over rel­atively long distances (tens of nanometers). They exhibit long range order. Twoother aluminosilicates. allophane and imogolite. exhibit short-range (or local)

INTRODUCTlON TO SOIL MINERALOGY 27

arder. Structures with short-range order exhibit order over several nanometers, buton a larger scale the structure is disordered.

Allophane is a material consisting chemically of variable amounts of ()2-,OH-, A13+, and Si4

+ and characterized by short-range order and a predominance ofSi-o-AI bonds. It consists of small (3.5-5.0 nm) spheres, the structure of whichhas not been determined. The spheres clump together to form irregular aggregates.

Imogolite consists oftubes several micrometers long with an outer diameterof2.3 to 2.7 nm and an inner diameter of-1.0 nm. The tubes consist ofa single dioc­tahedral sheet with the inner surface OH replaced by SiO,OH groups (Plate 1-8;Farmer et aI., 1983). Several individual tubes are arranged in bundles lOto 30 nmacross to give thread-like particles several micrometers long.

Allophane and imogolite usually occur as weathering products of volcanieash and are important minerais in the Andisol soil order. lmogolite has also beenidentified in the Bs horizons of Spodosols. Allophane and imogolite can specifi­eally adsorb many inorganic and organic compounds. Andisols, for example, usu­ally fix large amounts of phosphate, making it unavailable to plants, and the largeamounts of organic matter common in Andisols may be due, in part, to adsorptionoforganic molecules by allophane and imogolite. Soils containing large amountsof allophane and imogolite usually have unique physical properties such as a lowbulk density, high water holding eapacity, high liquid and plasticity 1imits, and athixotropic consistence.

IV. SOME CRYSTALLOGRAPHIC CONCEPTS

ln 1912 Max von Laue conducted an experiment to see ifthe x-rays diseov­ered 17 yr earlier by Wilhelm Conrad Rõntgen could be diffracted by a crystal. Thisone experiment had two very important results: (i) it proved that x-rays behave aswaves, and (ii) it showed that crystals are made up ofa regular array ofatoms inspace, and can therefore serve as a grating to diffract the x-rays. Within a few yearsofvon Laue's discovery, much ofthe mathematieal theory deseribing the diffrae­tion ofx-rays by crystals had been developed and the atomic structures ofsome sim­pie crystals had been determined (Azároff, 1968; Klug & Alexander, 1974). As earlyas 1923 and 1924, severa1clays were shown to be crystalline by x-ray diffraetion,but it was not until 1930 that some ofthe first strueture determinations ofclay min­erais were made. About the sarne time, two independent papers, one by S.B. Hen­drieks and w'H. Fry in 1930 and the other by w'P. Kelly, w'H. Dore, and S.M.Brown in 1931, gave the first x-ray diffraction evidence that soil materiaIs, eventhose in the finest fraetions, were crystalline (Marshall, 1949; Grim, 1968).

X-ray diffraction remains the most important method for identif)ting soil min­erais. A working knowledge of certain concepts of crystallography and x-ray dif­fraetion is necessary for understanding day mineralogy. A brief introduction to someofthese coneepts follows.

A. Periodicity in Crystals

ln a erystal, a particular pattem or arrangement ofatoms is repeated over andover in three dimensions. Repeating pattems oceur not only in erystals but also in

8 b

SCHULZE

.,l. ..••••........ . ..

I y••••

.t....-';Ir-...........,..-~_

Fig. I-I. (a) A lattice array af leaves and (b) a plane lattice of points.

many familiar places such as wallpaper, tile floors. and brick walls. Concepts thatapply to crystals can be developed by starting with these familiar examples.

Figure I-Ia illustrates a pattem made up ofleaves much like the pattem ona piece ofwallpaper. The periodicity ofthis pattem can be characterized with threeparameters. Along the X axis the pattem is repeated at intervals ofa. We also saythat the pattem is translated (moved) along X at increments of a. The translationinterval along the Yaxis is designated b and the angle between the X and Yaxes isdesignated by the angle y. The array ofleaves in Fig. I-Ia consists ofreal objects.namely. leaves. We can represent the periodic nature ofthis pattem by an array ofpoints in space. each point having identical surroundings (Fig. I-I b.). Such an arrayis called a lattice. A lattice of points is imaginary since each point is an imaginaryinfinitesimal spot in space. The concept of a lattice is useful because it allows usto represent the periodic nature of a pattem regardless of the actual object or con­figuration ofobjects at each lattice point. The lattice depicted in Fig. I-I b is calleda plane lattice because it has only two dimensions. The concept can be extendedto three dimensions to give a space lattice. A space lattice is a regular and unlim­ited distribution ofpoints in space (Fig. 1-2). The concept ofa space lattice is usedin describing crystal structures because crystals are three dimensional objects.3 Thedirections ofthe crystallographic axes in three dimensions are designated by X. Y.and Z, and a. b, and c are used for the repeat distances along these axes (AIPEANomenclature Committee. 1980). The angles between the axes are designated abetween Yand Z, pbetween X and Z, and y between X and Y.

B. Tbe Unit Cell

Joining the points of a space lattice produces a series of parallel-sided unitcells (Fig. 1-2). Each unit cell contains a complete unit of the crystal pattem be­cause the complete pattem is reproduced at each lattice point. Note that there arealtemative ways of outlining the unit cell. The choice of the "best" unit cell for amineral is made when the structure is deterrnined. A unit cell is chosen for conve-

] The tenns fal/ice and structure are ofien misused by authors and speakers. The two tenns are notsynonymous. Lal/ice is the mathematical concept ofan infinite unifonn distribution ofpoints in space(e.g., the 14 Bravais lattices). Structure refers to the actual physical assemblage of atoms (e.g., the 2: 1slructure) (AIPEA Nomenclature Committee, 1980). The concept ofa lattice is used to describe the pe­riodicity in crystal slructures.

INTRODUCTION TO SOIL MINERALOGY 29

Fig. )-2. A space lanice with several altemative unit cells outlined (after Klug and Alexander, 1974).

nience in visualizing the symmetry and in making mathematical calculations. Dif­ferent researchers will occasionally choose different unit cei Is to describe the sarnemineral structure.

There are 14 unique ways of arranging points in three-dimensional space.These are known as the 14 Bravais lattiees. They forro the geometrieal basis for alipossible unit eells. Each Bravais lattiee belongs to one of six erystal systems de­pending on the symmetry ofthe lattiee (Table 1-5). The eubie system is most sym­metrie. Ali axial translations are equal (a = b = c) as are ali axial angles (a. = ~ =y). The triclinie system is most general (and least symmetrie). Ali axial translationsand axial angles are different. The unit eell dimensions and the axial angles are im­portant parameters in deseribing the erystal strueture of a mineral.

C. Miller Indices

A consistent and concise notation is used when planes ofatoms in a crystalare diseussed. Different planes are generally referred to by their Miller indices, des­ignated as (hkl). Forexample, we may speak of(OOI) planes, (110) planes, and (431)planes. The eoneept of Miller indiees follows from the lattice eoneept diseussedabove. Consider the lattice ofpoints in Fig. 1-3a and a plane that is parallel to cand passes through any two lattice points (represented in an edge view by the heavyline). An identieal plane must pass through eaeh and every lattiee point (Iightlines) beeause the lattiee is periodie and eaeh point is the sarne as any other. Thisfamily ofparallel planes euts the a dimension ofthe unit cell into two parts and theb dimension into three parts. The c dimension is not eut at ali (zero parts) because

Table l-S. Axial ratios and angles between crystal axes for lhe six crystal symmetry systems (after Klugand Alexander. 1974).

System

Cubic (isometric)HexagonalTetragonalOrthorhombicMonoclinicTriclinic

Axial ratios

a=b=ca=b~c

a=b~c

a~b~c

a~b~c

a#h#c

Angles between crystal a.xes

a=~=y=90°

a = ~ = 90°. Y= 120°a=~=y=90°

a=~=y=90°

a=y=90°,~>90°

a#~#y=90°

30

a b c

SCHULZE

Fig. )-3. (a) Planes with Miller indices 230 c axis is out ofthe plane ofthe figure) and (b) planes withdifferent Miller indices on an orthogonallanice.

the planes are parallel to Z. The Miller indices are thus (230). ln tenns of(hkf), theplanes cut the a dimension ofthe unit cell into h parts, the b dimension into k parts,and the c dimension into I parts. Additional examples of Miller indices are givenin Fig. 1-3b.

D. X-ray Diffraction

X-ray diffraction (XRD) is the main analytical technique used both to iden­tify unknown mineral phases and to detennine crystal structures. Soil minerais areusually studied using the powder diffraction method. A powdered sample, with par­ticles typically <50 Ilm in diameter, is placed in a diffractometer and irradiated withx-rays. A plot ofdiffracted x-ray intensity vs. the diffraction angle (i.e., 28) is ob­tained. A simple mathematical relationship called the Bragg Law is then used to re­late the peaks on the x-ray chart with the distances between the diffracting planesofatoms within the crystals.

1. The Bragg Law

X-rays are electromagnetic waves with wavelengths on the order ofthe spac­ings between the planes ofatoms in crystals. Thus, crystals diffract x-rays. Under­standing the additive relationships ofwaves that are in phase or out ofphase is im­portant to understanding x-ray diffraction. Fig. l--4a illustrates three different sinewaves, ali ofwhich begin at the sarne point in their cycles. The waves are in phasebecause the maxima of each wave coincide. The sum ofali three waves is a muchlarger wave with the sarne period (distance between successive maxima) as the orig­inai waves but with three times the intensity. The three waves on the right are iden­tical to those on the left, except that they are out ofphase. Each wave begins at adifferent point in its cycle relative to the others. The maxima ofthe three waves nolonger coincide, and thus the sum, which takes into account the positive and neg­ative aspects ofthe three individual waves, approaches zero.

ln XRD, constructive interference occurs when many waves are in phase. Theenergy from the large summation wave is detected by the diffractometer. Destruc­tive interference occurs when the waves are out ofphase. ln this case no apprecia­ble energy is detected. The peaks ofan XRD pattem correspond to areas where con-

INTRODUCTION TO SOIL MINERALOGV

a Waves in Phase Waves out of Phase

b

31

Fig. 1-4. (a) ln phase and oul ofphase waves and (b) geometry oflhe Bragg reflection analogy for dif­fraclion ofx-rays by crystals (after AzárotT, 1968).

structive interference occurs. The Bragg Law relates the position ofthese peaks tothe distances between planes of atoms in the crystal. This information. along withthe intensities of the peaks is then used for identif)'ing mineral phases and deter­mining crystal structures. The following derivation ofthe Bragg Law will show howit works.

Consider the parallel planes of atoms with indices (hkl) separated by a dis­lance d depicted in Fig. l--4b. A parallel (or collimated) beam ofx-rays ofwave­length à impinges on the (hkl) planes at an angle e. The incident x-rays are in phasewith one another. Two waves. one traveling along OA and being retlected off theupper plane of atoms and traveling along AP. and a second wave traveling alongO'C and being retlected offthe second plane ofatoms and then traveling along Cp'may. under specific conditions. be in phase when they reach points P and p'. re­spectively. ffthe waves are ín phase at fine AB. when will they be in phase at lineAD and thus at PP'? A wave along CP' is in phase with a wave along AP only ifthedistance BC + CD is an integral number ofwave\engths. Thus

BC + CD = nÃ. where n = 1,2,3,...

Since BC = CD, substituting into Eq. [I] gives

2BC =nÀ..

[I]

[2]

32

From Fig. 1-4b,

sin e = BCld,

which can be rewritten as

BC=dsine.

Multiplying both sides by 2 gives

2BC = 2d sin e.

Substituting Eq. [2] into Eq. [5],

nÀ = 2d sin e.

This expression is the Bragg Law. Rewritten as

dln = Ã1(2 sin e),

SCHULZE

[3]

[4]

[5]

[6]

[7]

the Bragg Law relates the perpendicular distance between diffracting planes, d, tothe diffraction angle e. Since the order of diffraction, n, is not always known, n isusually set to 1and the quantity dln is referred to as the d-value ofa diffraction line.

Note that the quantity needed in Eq. [7] is in units ofe, while the data fromx-ray diffractometers is in units of2e. This is a consequence ofthe physical con­struction ofthe diffractometer and division by 2 is necessary to obtain e from 2edata before using Eq. [7] to calculate d-values. The d-values can be easily calcu­lated with a pocket calculator or a computer programo Tables for determining d-val­ues from 2e data are also given in Brindley and Brown (1980) for severaI commonx-ray wavelengths.

The distances between the diffracting planes in the crystal, din Eq. [7], arefixed by the structure ofthe crystal. However, x-ray wavelengths (Às) used for dif­fraction experiments are not always the sarne. IfÀ in Eq. [7] changes, e must changebecause dln is a constant fixed by the crystal structure ofthe sample. X-ray line po­sitions expressed in 2e can only be compared when the x-ray wavelength is lhe sarne.Calculating the d-values for the diffraction lines allows comparison ofd-values re­gardless ofthe wavelength.

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

[8]

where a, b, and c are the unit cell dimensions. Relationships for the other crystalsystems are given elsewhere (e.g., Klug & Alexander, 1974).

The dln values obtained from the Bragg Law can, in most cases, be relatedto actual distances between planes ofatoms in the crystals. Thus, the d-values be­tween successive phyllosilicate layers listed in Plate 1-7 show up in x-ray ditfrac-

INTRODUCTION TO SOIL MINERALOGY

tion pattems of soi! c1ays and are used to identify the specific c1ay minerais pre­sent in unknown samples4. ln some cases, however, in particular for randomly in­terstratified minerais or cases where diffraction lines are very broad due to the smallsize of the crystals, d/n values from the Bragg Law are not directly related to thedistances between planes ofatoms in the crystal. To explain these effects more in­clusive diffraction theory utilizing such concepts as the reciprocallattice, structurefactor, and atomic scattering factors must be applied.

2. Tbe Scberrer EquatioD

Once a mineral phase has been identified based on its diffraction peaks, ad­ditional information can ofien be obtained from diffraction line widths. Well-crys­tallized mineraIs ofsand and silt size give sharp diffraction lines whose widths aredetermined only by broadening caused by the x-ray diffractometer itself. Clay-sizeparticles show broader lines caused by diffraction effects from the small particles.The smaller the particles are, the broader the diffraction lines. X-ray line width andparticle size are related by the Scherrer equation:

Lhkl = (KÀ)/(~ cos e), [9]

where Lhkl is the mean crystallite dimension perpendicular to the diffracting planeswith Miller indices (hkl), K is a constant equal to 0.9 ifthe particles are cubes, À. isthe x-ray wavelength, ~ is the width (in radians) at one-halfpeak height correctedfor instrumental broadening, and e is the diffraction line position. The applicationofthe Scherrer equation involves a considerable number of assumptions (Klug &Alexander, 1974; p. 634-642,687-704>, and an understanding ofthese assumptionsis essential. ln addition, Lhkl measures the size ofthe coherently diffracting domainswithin the crystals. If the crystals consist of several coherently diffracting do­mains, then Lhkl will be smaller than the actual physical size ofthe crystals observedby electron microscopy (Schulze & Schwertmann, 1984).

V.SUMMARY

This chapter has developed some ofthe terminology used later in this book.lt has shown how the abundance of aluminosilicate minerais is a consequence ofthe abundance ofSi, AI, and O in the Earth's crust. lt has introduced the chemicaland structural c1assification ofminerals, mentioned the major minerais that will beencountered in soils, and mentioned some ofthe environmental implications ofsomeofthe minerais. lt has developed the major structural themes ofthe phyllosilicateminerais, and it has explained some of the crystallographic and x-ray diffractionterminology used in later chapters.

The chapters that follow will discuss in more detail the many facets of soilmineralogy, particularly as it relates to current environmental concems.

4 Strict usage ofSI units requires that d-spacings be given in nanometers (nml. Angstroms (Al, how­ever, are used in the older literature and are still used widely in lhe current crystallography literatureand in many ofthe databases used for mineral identification. There are 10 Ain [ nm (I nm = [O AJ.

34

VI. ACKNOWLEDGMENTS

SCHULZE

I thank the many students over the years who have provided feedback on thischapter.

VII. QUESTIONS AND EXERCISES

I. Draw sketches from memory of an octahedron and a tetrahedron using thesphere-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 a good templatefor a circle.)

3. Draw three sketches of a unit cell and label the cell dimensions a, b, and c asshown in Fig. 1-3b. On the first cell sketch in the planes with Miller indices (002),on the second cell sketch planes with indices (lI I ), and on the third cell sketchplanes with indices (021).

4. (a) Calculate the d-values ofx-ray diffraction lines that occur at 8.84, 12.36, and26.68° 28 for a sample x-rayed using CuKa radiation (À. = 0.15418 nm). (b) Cal­culate the 28 positions for these sarne three lines when the sarnple is x-rayed usingCoKa radiation (À. = 0.179026 nm).Answers: (a) 1.000,0.716, and 0.334 nm, respectively. (b) 10.27, 14.36, and31.09° 28, respectively.

5. A goethite sample was x-rayed with CoKa radiation. The III diffraction tinewas centered at 24.70° 28, and its width at half-height (corrected for instrumentalbroadening) was found to be 0.35° 28. Use the Scherrer equation to calculatethe size ofthe crystals (more exactly, the size ofthe coherently diffracting do­mains) perpendicular to the (110) planes. Assume that K = 0.9.Answer: 27 nm.

VIII. SUGGESTED READING

For more information on general mineralogy, see introductory texts such asKlein and Hurlburt (1993) or Nesse (1999). Azároff{ 1968) and Moore and Reynolds(1997) provide good, readable introductions to crystallography and x-ray diffrac­tion theory.

REFERENCES

AIPEA Nomenclature Comminee. 1980. Summary ofrecommendations ofAIPEA nomenclature com­minee. Clays Clay Miner. 28:73-78.

Allen, B.L., and D.S. Fanning. 1983. Composition and soil genesis. p. 141-192./n L.P. Wilding et aI.(ed.) Pedogenesis and soil taxonomy I. Concepts and interactions. Elsevier. New York, NV.

Azároff, L.v. 1968. Elements ofx-ray crystallography. McGraw-Hill Book Co., New York, NV.Brindley, CiW., and Ci Brown (ed.) 1980. Crystal structures ofclay mineraIs and their x-ray identifi­

cation. Mineralogical Society, London, UK.

INTRODUCTlON TO SOIL MINERALOGY

Fanner, v.c., M.J. Adams, A.R. Fraser, and F. Palmieri. 1983. Synthetic imogolite: Properties, synthesis,and possible applications. Clay Miner. 18:459--472.

Grim, R.E. 1968. Clay mineralogy. 2nd ed. McGraw-Hill Book Co., New York, NY.Jackson, M.L. 1964. Chemical composition ofsoils. p. 71-141./n F.E. Bear (ed.) Chemistry ofthe soi!.

Reinhold Publ., New York. NY.Kãmpf, N., A.C. Scheinost. and D.G Schulze. 1999. Oxide minerais. p. FI25-FI68./n M.E. Sumner

(ed.) Handbook ofsoil science. CRC Press, Boca Raton, FL.Klein, C., and C.S. Hurlbut, Jr. 1993. Manual ofmineralogy (after James D. Dana). 21st ed. John Wiley

& Sons, New York. NY.Klug, H.P., and L.E. Alexander. 1974. X-ray diffraclion procedures for polycrystalline and amorphous

materiais. 2nd ed. John Wiley & Sons, New York. NY.Marshall, C.E. 1949. The colloid chemistry ofthe silicale minerais. Academic Press. New York, NY.Moore, D.M., and R.C. Reynolds. Jr. 1997. X-ray diffraction and identification and analysis ofc1ay min­

erais. 2nd ed. Oxford University Press, New York. NY.Nesse, w'D. 1999. Introduclion lo mineralogy. Oxford University Press, New York, NY.Rothbauer. R. 1971. Untersuchung eines 2M 1-Muskovits mit Neutronenstrahlen. N. Jahrbuch f. Min­

eralogie. Monatshefte 1971: 143-154.Schulze, D.G, and U. Schwertmann. 1984. The influence ofaluminium on iron oxides: X. Propenies

ofAI-substituted goethites. Clay Miner. 19:521-539.