Prediction of Mineralogy of Some Soils of Ganges Floodplain Under Different Cropping Patterns

7/28/2019 Prediction of Mineralogy of Some Soils of Ganges Floodplain Under Different Cropping Patterns

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PREDICTION OF MINERALOGY OF SOME SOILS OF GANGES

FLOODPLAIN UNDER DIFFERENT CROPPING PATTERNS

HUMAIRA HASANStudent ID: 081321

Session: 2010-11

______________________________________

Soil Science Discipline

Khulna University

Khulna

April, 2012

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Course Title: Project Thesis

Course No: SS-4106

This project thesis paper has been prepared and submitted to Soil Science

Discipline, Khulna University, for the partial fulfillment of four years

professional B. Sc. (Hons) degree in Soil Science.

Submitted By

________________________________HUMAIRA HASAN

Student ID: 081321

Session: 2010-11

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Approved As To Style And Content By

...

KHANDAKAR QUDRATA KIBRIA

Associate ProfessorHead & Chairman of Examination Committee

Soil Science DisciplineKhulna University, Khulna

Bangladesh

___________________________


Khulna University

Khulna

April, 2012

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DECLARATION

This project thesis paper is entirely the candidates own

investigation.

Supervisor

Md. Sadiqul Amin

Assistant ProfessorSoil Science Discipline

Khulna University, Khulna

Candidate

HUMAIRA HASAN

Student ID: 081321

_______________________________________


Khulna University

Khulna

April, 2012

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ACKNOWLEDGEMENT

At first I wish to express all my devotion and reverence to almighty Allah, who has

enable me to complete the thesis paper for the fulfillment of the degree of Bachelor of

Science in Soil Science.

I would like to express my heartfelt gratitude, sincere appreciation and profound

regards to Md. Sadiqul Amin, Assistant professor, Soil Science Discipline, Khulna

University, under whose careful supervision and scholastic guidance this thesis has

been conducted.

I am sincerely thankful to Khandakar Qudrata Kibria, Associate professor and

Head and Chairman of the Examination Committee, Soil Science Discipline,

University of Khulna for providing all possible facilities during the course of this

thesis paper and all other teachers of Soil Science Discipline for their cordial co-

operation at different stages for accomplishing this study.

I convey my respect and gratefulness to Md. Sanaul Islam,Associate professor, Soil

Science Discipline, University of Khulna for his cordial co-operation during this

study.

Special appreciation goes to my friend Falguni Akter for her excellent co-operationduring this study.

I am also thankful to my senior Sobuj Ahmed and junior Ali Reza, for their cordial

co-operation during my work.

Finally, I respectfully remember my parents and my younger sister for their support

for my higher education and without whose blessings this work would not have been

possible.

April, 2012. The

Author

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DEDICATION

To my beloved family

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CONTENT

Titles Page No.

CONTENT I

LIST OF TABLES IV

LIST OF FIGURES V

Chapter one: Introduction

1. Introduction 1

Chapter two: Literature review

2. Literature review

2.1. Primary soil minerals2.2. Secondary soil minerals

2.3. Formation of clay minerals

2.3.1. Formation of Vermiculite

2.3.1.1. Release of Potassium

2.3.1.2. Oxidation of Iron

2.3.1.3. Hydroxyl Orientation

2.3.2. Formation of Smectite

2.3.2.1. Transformation from Mica

2.3.2.1.1. Environment that promote thetransformation of mica

2.3.2.2. Transformation from Chlorite

2.3.2.3. Formation from Solution

2.3.3. Formation of Kaolinite

2.3.3.1. Equilibrium environment and conditions forsynthesis of Kaolinite

2.3.3.2. Kaolinite formation from Hydroxy-AlInterlayered Montmorillonite

2.4. Structure of clays minerals

2.4.1. The 1:1 Type Minerals

3

56

7

8

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9

9

10

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10

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Titles Page No.

2.4.1.1. Kaolinite

2.4.1.2. Halloysite

2.4.2. The 2:1 type Minerals

2.4.2.1. Smectite (Montmorillonite)

2.4.2.2. Vermiculite

2.4.2.3. Mica

2.4.2.4. Illite

2.5.2.5. Chlorites

2.4.3. Interstratified Clay Minerals

2.4.3.1. Montmorillonite-Vermiculite-ChloriteIntergrade

2.4.3.2. Swelling 2:1 to 2:2 intergrades

2.5. Soil mineral weathering sequences

2.6. Pridiction of clay mineralogy

2.7. Ganges Floodplain soils of Bangladesh

2.8. The agricultural landuse of Ganges Floodplain soils

2.9. Clay mineralogy of the soils of Bangladesh

17

19

20

20

21

23

24

25

26

27

27

28

35

35

36

36

Chapter three: Materials and Methods

3. Materials And Methods

3.1. Study area

3.2. Collection of soil samples

3.3. Processing of soil samples

3.4. Soil analyses

3.4.1. Particle size analysis3.4.2. pH

3.4.3. ECe

3.4.4. Organic carbon

41

41

42

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Titles Page No.

3.4.5. Cation exchange capacity

3.4.6. Exchangeable K+

3.4.7. Water soluble K+

3.4.8.Water soluble Mg2+

42

43

43

43

Chapter four: Results and Discussion

4. Results and Discussion

4.1. Physical Characteristics

4.2. Chemical Characteristics

4.2.1. pH

4.2.3. CEC

4.2.4. K+ content

4.2.5. Mg2+ content

4.2.6. Organic C and humus

4.3. Prediction of clay minerals

4.4. Weathering stability of clay minerals

44

44

44

44

44

44

46

46

46

46

Chapter five: Summary and conclusion 49

Chapter six: References 50

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LIST OF TABLES

Titles Page No.

Table 2.1 Common primary mineral of soil 6

Table 2.2 Common secondary Minerals in Soils 7

Table 2.3 Types of minerals as indicators for relative degree of soildevelopment (the higher the number, the higher the degreeof development

30

Table 2.4 The typical ratio of CEC to clay for some classes of claymineralogy

35

Table 3.1 General information of sampling sites 41

Table 4.1 Different properties of studied soils 45

Table 4.2 The clay mineralogy of the soil samples 46

Table 4.3 Pedological indices for studied soil samples 47

Table 4.4 Pedological indices of different clay minerals underequilibrium condition

47

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LISTS OF FIGURES

Titles Page No.

Fig 2.1 General conditions for the formation of the various layersilicate clays and oxides of iron and aluminum

8

Fig 2.2 Composition of solutions equilibrated withmontmorillonites

12

Fig 2.3 Tetrahedron, octahedron and classification of silicatemineral

16

Fig 2.4 Models of a 1:1 layer structure 17

Fig 2.5 Schematic structure of Kaolinite 18

Fig 2.6 Schematic structure of Halloysite 19

Fig 2.7 Models of a 2:1 layer structure 20

Fig 2.8 Schematic structure of montmorillonite 21

Fig 2.9 Schematic structure Vermiculite 22

Fig 2.10 Schematic structure of Mica 24

Fig 2.11 Schematic structure of Chlorite 26

Fig 2.12 Mineral stability sequence. Stability of minerals 29

Fig 2.13 Relative solubility of a few clay minerals 32

Fig 2.14 Relative solubility of selected primary and secondaryminerals at pH 7.0

34

Fig 2.15 A clay mineralogical map of Bangladesh soils 40

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1. Introduction

Bangladesh is the largest deltaic floodplain in the world with a total area of 147570

km2 of which 88892 km2 is occupied by major rivers and estuaries. The great delta is

flat throughout and sketches from near the foothills of the Himalayas Mountain in thenorth to a southern irregular deltaic coastline that faces the Bay of Bengal. The major

portion of the Bengal basin of the Sindhu-Ganges depression lies in Bangladesh

(Saheed, 2010).

Bangladesh has a wide range of mineral soils developed in parent materials derived

from various sources ranging from fresh alluvial deposits of recent origin to strongly

weathered old alluvium and sedimentary rocks of tertiary formation. The soil formed

under varied topographic and hydrological conditions include poorly to very poorlydrained, strongly gleyed hydromorphic soils of the young alluvial lands on one hand

and well drained, red-brown, deeply weathered and leached latosolic soils of the

upland terrace and hills of the other. These soils are tentatively correlated in the

USDA Soil Taxonomy with Entisols, Inceptisols, Mollisols, Alfisols, Ultisols and

also may be with some Vertisols, encompassing about 512 soil series altogether. At

the sub-order level they are: Aquents, Psamments, Orthents, Aquepts, Ochrepts,

Umbrepts, Aquolls, Udalfs, Ustalfs, Udults and Ustults and also may be some Uderts.Among them the Aquepts are by far the most dominant soil covering about 60% of

the area (Saheed, 1985).

In Bangladesh, some data on both sand and clay mineralogy are available but they are

not adequate. After the reconnaissance soil survey, it has become apparent that for a

more efficient use of the soil survey information a further comprehensive grouping of

these soils series is needed at the family level. For this a systematic mineralogical

study is urgently required (Saheed, 1985). Recently formed delta and alluvial plains ofthe Ganges, Brahmaputra andMeghna rivers are the most extensive area of cropping

practices. Pedological indices for determination of the weathering stability of soil

mineral could be influenced by different type cropping pattern.

The objectives of this study is

To predict the mineralogy of some soils of Ganges floodplain of

Bangladesh under different cropping pattern.

To describe the weathering stability of the clay minerals.

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2. Literature Review

The solid phase of soil is composed inorganic and organic materials. Of these the

inorganic materials, or minerals, predominate in virtually all soils except Histosols

(Essington, 2004). A mineral is a naturally occurring, inorganic, solid element orcompound with a definite chemical composition and a regular internal crystal

structure (Montgomery, 2000). Minerals are generally classified as primary minerals

and secondary minerals (Essington, 2004). A primary mineral is one that has not been

altered chemically since its deposition and crystallization from molten lava or magma.

Quartz and feldspar are common primary minerals in the soil. A secondary mineral is

one resulting from the weathering of a primary mineral; either by an alteration in the

structure or from reprecipitation of the products of weathering (dissolution) of aprimary mineral. Common secondary minerals in soils are the aluminosilicate or

phyllosilicate minerals such as kaolinite and montmorillonite, oxides such as gibbsite,

goethite, amorphous materials such as imogolite and allophane (Sparks, 1995). In soil,

a naturally occurring material composed primarily of fine-grained minerals (


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The kind and amount of clay influence plant growth by affecting available water

capacity, water intake rate, aeration, leaching, erodibility etc. By slightly slowing the

rate of water movement, it can reduce the rate of nutrient loss through leaching. It also

determines the swelling capacity of the soil (Brady and Weil, 2002). Mineral

wreathing releases plant nutrients that are retained by other minerals through

adsorption, cation exchange, and precipitation. Clay content can be increase or

decrease in the soil. Up to a certain point, an increase in the amount of clay in the

subsoil is desirable. The amount of clay accumulation and its location in the profile

provide clues for the soil scientist about soil genesis. Irregular clay distribution as

related to depth may indicate lithologic discontinuities, especially if accompanied by

irregular sand distribution (Dixon and Weed, 1989).

According to Jenny (1941) soil formation depends upon some factors. These are

isolated as the climatic factor, the biotic factor, the topographic factor, the parent

material and the time factor (Churchman, 2000). The formation of clay minerals in

soils generally results from the combination and addition of ions and molecules from

the soil solution to the solid phase. Formerly, soil minerals were thought to form by

differential migration of ions into and out of existing silicate structures. The diffusion

of A13+ or Mg2+ out of the lattice was supposedly balanced by the inward diffusion of

other ions. Clay mineral formation occurs by ion substitution between soil solution

and an existing solid (Bhon et al., 1979). The synthesis of the clay minerals at

elevated temperatures and pressures from oxides and hydroxides and from various

minerals, particularly the feldspars, in the presence of acids and alkalies has been

studied in considerable detail. The environmental conditions favorable for the

formation of the clay minerals can be decided from the available data from synthesis

experiments (Grim, 1986). At low temperatures and pressures, acid conditions

apparently favor the formation of the kaolinite type of mineral, whereas alkaline

conditions favor the formation of smectite or mica, if potassium is the alkali metal and

if it is present in concentration above a certain level. The presence of magnesium

particularly favors the formation of montmorillonite. At temperatures somewhat

above about 350C and with moderate pressures, pyrophyllite forms in place of

kaolinite, with an excess Al2O3 forming boehmite. At more elevated temperatures and

pressures, other alumina phases develop. Mica can form under acid conditions, and

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kaolinite or pyrophyllite can develop in the presence of an excess of K2O, depending

on the temperature, the amount of K2O and the Al2O3:SiO2 ratio (Grim, 1986).

2.1. Primary soil minerals

The mineral which has not been altered chemically since its deposition and

crystallization from molten lava is called primary mineral (Brady and Weil, 2002).

Most common primary minerals in soils are quartz and feldspar. Other primary

minerals that are found in soils in smaller quantities include pyroxenes, micas,

amphiboles and olivines. Primary minerals occur primarily in the sand and silt

fractions of soils but may be found in slightly weathered clay-sized fractions (Sparks,

1995). The most common primary minerals found in soil are presented in Table 2.1.

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Table 2.1

Common primary mineral of soil

Name Chemical formula

Quartz SiO2

Muscovite KAl2 (AlSi3 O10) (OH)2

Biotite K(Mg, Fe)3(AlSi3 O10) (OH)2

Feldspar

Orthoclase

Microcline

Albite

KAlSi3O8KAISi3O8

NaAlSi3O8

Amphiboles

Tremolite Ca2Mg5Si8O22(OH)2

Pyroxenes

EnstatiteDiopside

Rhodonite

MgSiO3CaMg(Si2O6)

MnSiO3

Olivine (Mg, Fe)2SiO4

Epidote Ca2(Al, Fe)3Si3 O12(OH)

Tourmaline (Na, Ca) (Al, Fe3+, Li, Mg)3 Al6 (BO3)3 (Si6O18) (OH)4

Zircon ZrSiO4

Rutile TiO2

Source: Sparks, 1995

2.2. Secondary soil minerals

The secondary mineral is one resulting from the weathering of a primary mineral

(Table 2.2) either by an alteration in the structure or from reprecipitation of the

products of weathering (dissolution) of a primary mineral. Common secondary

minerals in soils are the aluminosilicate minerals such as kaolinite and

montmorillonite, oxides such as gibbsite, goethite, and birnessite, amorphous

materials such as imogolite and allophane, and sulfur and carbonate minerals. The

secondary minerals are primarily found in the clay fraction of the soil but can also be

found in the silt fraction (Sparks, 1995).

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Table 2.2

Common secondary Minerals in Soils

Name Chemical formula

Clay minerals

KaoliniteMontmorillonite

Vermiculite

Chlorite

Si4Al4O10( OH)8MX (Al, Fe2+, Mg)4Si8O20 (OH)4 [M = interlayer metal

cation]

(Al, Mg, Fe3+)4 (Si, Al)8O20(OH)4[M Al (OH)6](Al, Mg)4(Si, Al)8 O20(OH, F)4

Allophane Si3Al4O12 . nH2O

Imogolite Si2Al4O10 . 5H2O

Goethite FeOOH

Hematite -Fe2O3

Maghemite -Fe2O3

Ferrihydrite Fe10O15 . 9H2O

Bohemite -AlOOH

Gibbsite Al (OH)3

Pyrolusite -MnO2

Gypsum CaSO4 . 2H2O

Source: Sparks, 1995

2.3. Formation of clay minerals

The silicate clays are developed from the weathering of a wide variety of minerals by

at least two distinct processes: (1) a slight physical and chemical alteration of certain

primary minerals, and (2) a decomposition of primary minerals with the subsequent

recrystallization of certain of their products into the silicate clays (Brady and Weil,

2002). The simplified framework of mineral weathering is presented in Fig. 2.1.

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Fig. 2.1 General conditions for the formation of the various layer silicate clays and oxides of iron andaluminum (Brady and Weil, 2002)

2.3.1. Formation of Vermiculite

Vermiculites will not form from solidification of a magma, and it is most universally

assumed that the mineral can be formed only by alteration of a micaceous mineral.

Borchardt et al. (1966) showed that trioctahedral vermiculite pseudomorphs after

mica are common in many soils. Trioctahedral vermiculite in soils may be formed by

alteration of biotite,phlogopite, or chlorite either in the soil or in the parent material

(Johnson, 1964). The alteration of micas to vermiculites may be subdivided into

separate steps, orfactors affecting alteration: (i) release of K, (ii) oxidation of Fe2+and

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(iii) OH orientation. These reactions may be mitigated by the presence of organic

acids (Robert et al., 1979).

2.3.1.1. Release of Potassium

The rate of release of K from interlayer positions in mica and replacement by other

ions may be considered a diffusion process, with diffusion of K+ out and diffusion of

the counter ion to the vacated spot (Chute and Quirk, 1967). A biotite disk, when

placed in an appropriate salt solution, will alter so that in a short time a vermiculite

halo will appear to encircle the biotite core. Alteration proceeds parallel to the X and

Y directions, with little apparent alteration parallel to the Z direction in the crystal.

Potassium may diffuse from all interlayers, with mica going directly to vermiculite. In

other cases K may diffuse along a specific (001) plans but not from adjacent (001)planes. The latter causes a wedge-shaped K+ depleted zone in the mica (Rich, 1964).

This process results in the widely reported (Bassett, 1959; Roux et al., 1963) mica-

vermiculite interstratified minerals. Very small amounts of K in solution will prevent

the replacement of K+ by such cations as Mg2+. The reaction biotite to vermiculite will

proceed only when leaching is effective removes K (Bassett, 1959).

2.3.1.2. Oxidation of Iron

Most of the iron in biotite is in the reduced state, and most of the iron trioctahedral

vermiculite is in the oxidized state (Weaver and Pollard, 1973). Some of the

octahedralposition are lost during the alteration process and some Al migratesfrom

tetrahedral sites to octahedral sites (Vicente et al., 1977), resulting in a larger

proportion of more stable dioctahedrel sites in the weathering product. This increase

in number of dioctahedrel sites also causes an increase in the rotation of tetrahedral

and thus decreases the length of the b axis (Farmeret al., 1971). Divalent iron may be

dissolved and lost via the soil solution or oxidation and precipitation. The

simultaneous oxidation and expulsion of iron from a mica may result in the formation

of either hydroxy-Fe vermiculite or a crystalline external phase of-FeOOH. Under

acidic conditions the oxidation of Fe is enhanced by a need by a decrease in total

charge of the octahedral sheet through loss of octahedral Fe and Mg; the decrease in

surface charge is minimal and vermiculite is formed (Farmeret al., 1971).

2.3.1.3. Hydroxyl Orientation

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In trioctahedral micas OH- ions sit in a hole above three divalent (Fe 2+ or Mg2+) ions.

This causes a uniform charge distribution below the OH and repels the H portion of

the OH. In trioctahedral micas the H end of the OH is pointed directly toward the K.

In dioctahedral micas the cations in the octahedral sheet are primarily trivalent (Al3+),

with only two-thirds of the available octahedral positions occupied; each OH sits

above two cations and a structure vacancy. The octahedral cations orient the OH- ion

so that the H- ion points in the general direction of the interlayer K+, but is tilted

toward the vacant octahedral site. Thus, in trioctahedral micas the K+ sits in a volume

strongly affected by the repulsive forces (polarity) of the H portion of the OH (Dixon

and Weed, 1989).

2.3.2. Formation of Smectite

In the environment smectite can form by the transformation of mica, chlorite and

solution.

2.3.2.1. Transformation from Mica

Similarities in their sheet structures have long led to the conclusion that smectites

could be derived from micas by depotassication (Crawford et al., 1983). Smectites

weathered from micas are likely to be tetrahedrally substituted, approaching the

beidellite end member in chemical composition (Ozkan and Ross, 1979).

Trioctahedral micas are likely to produce unstable trioctahedral smectites that tend to

lose Mg and Fe from the octahedral sheet during further weathering. The term

transformation smectite has been proposed for smectites derived from micaceous

minerals (Robert, 1973). The following equation describes the weathering of

phlogopite mica to saponite, a trioctahedral smectite: (Robert, 1973)

(Si3Al)(Mg3)O10(OH)2K + 0.5Si4+ + 0.25Ca2+

(Si 3.5Al0.5)(Mg3)O10(OH)2Ca0.25 + 0.5Al3+ + K+

The equation illustrates the essential changes necessary for smectite transformation

from a mica. They are (Komarneni et al., 1985)

1. Depotassication as in vermiculite formation

2. Dealumination of the tetrahedral sheet followed by

3. Silication of the tetrahedral sheet

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Without these extensive changes, micas do not have the low-charge characteristics

necessary for them to exhibit the properties of smectites.

2.3.2.1. Environment that promote the transformation of mica

First, temperatures and pressures must be low enough to destabilize the tetrahedral Al

that distinguishes weatherable micas (Jackson, 1963). Second, the concentration of K

must be low, such as in rainwater and in soil solutions of heavily cropped soils. Third,

Si(OH)4 concentrations must be high as provided by mafic minerals with Si potentials

higher than that of smectite (Huang, 1966). Fourth, Al concentrations must be low, as

they are in soils with pHs above 6 or 7. Thus, soils with low K+ and Al3+ and high

Si(OH)4 and Ca2+ or Mg2+ activities and pH above 6.5 are likely to contain

transformation smectites. At a soil pH < 6, mica weathers to vermiculite, whichsubsequently tends to form kaolinite (Ismail, 1970).

2.3.2.2. Transformation from Chlorite

Mafic chlorite is highly unstable in most soils. It would he expected to lose its

hydroxide interlayer at pH levels < 6 and under severe leaching and oxidizing

conditions, thus forming a smectite that is, in turn, relatively unstable due to its

trioctahedral nature. These are not conditions that would be expected to preserve

smectite indefinitely. The laboratory weathering of trioctahedral, Fe-rich chlorite can

be accomplished in a saturated Br2 solution at 100C (Senkayi et al., 1981). The

reaction proceeds with the production of regularly interstratified chlorite-vermiculite

and chlorite-smectite intermediates occurring in the first 2d. Within 2 weeks, the

chlorite disappears entirely (Adams et al., 1971).

2.3.2.3. Formation from Solution

Smectites that precipitate directly from soil or matrix solutions may be calledneogenetic smectites. Vast regions of the world have simectitic soils formed primarily

in young glacial and alluvial materials, a fact suggesting that smectite neogenesis

occurs rapidly if it occurs at all (Dixon and Weed, 1989).

2.3.3. Formation of Kaolinite

Formation of kaolinite depends upon the Equilibrium environment and conditions.

There are also possibilities to form kaolinite from Hydroxy-Al Interlayered

Montmorillonite.

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2.3.3.1. Equilibrium environment and conditions for synthesis of Kaolinite

Formation of kaolinite in the presence of smectite at 25C provides an example of

mineral precipitation under conditions much like those in field soils. Several samples

of three reference smectites were acidified with HCl and adjusted with solutions of Siand Al to provide a range of conditions prior to incubation (Kittrick, 1970). Analyses

were made periodically to determine solution compositions. After 3 to 4 year, the

formation of kaolinitic was shown in some samples. A stability line for kaolinite

derived from data obtained at equilibrium and an equation for the precipitation of

kaolinite are shown in Fig. 2.2.

Fig. 2.2. Composition of solutions equilibrated with montmorillonites from Belle Fourche, South

Dakota ( ); Otay, California (); and Aberdeen, Mississippi () after 3 to 4 year of equilibration.

Solid symbols indicate kaolinite formation, and open symbols indicate no detectable kaolinite. Size of

the symbols indicates analysis precision. Arrows indicate the direction of sample equilibration as

shown by previous analyses. Dashed solubility lines indicate metastable area (Kittrick, 1970)

.

A stability line for kaolinite is drawn from Eq. [2.1] for the precipitation of kaolinite

from Al3+ and H4SiO4

2Al3+ + 2H4SiO4 + H2O 6H+ + Si2Al2O5(OH)4(kaolinitic) . [2.1]

The equilibrium constant, K, for the reaction is

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An activity of unity is assumed for kaolinite and H2O. Taking log on both sides. [Eq.

2.2]

log K = -6pH 2log Al3+ - 2log H4Sio4 .. [2.2]

Dividing by 2 and rearranging gives [Eq.2.3]

log Al3+ + 3pH = -log H4Sio4 log K [2.3]

The slope of the kaolinite stability line in fig. 2.2 is -1, and the intercept is -1/2 log K

from Eq. [2.3]. The intercept is based on the sample with the lowest log H 4SiO4-

thepoint ( ) at the far left in fig. 2.2 The last analysis of this sample had a pH of 3.47,

log H4SiO4 of -3.53 (assuming all dissolved Si as H4SiO4) and log Al3+ of -3.78

(corrected from a molar concentration of log Al = -3.59 for Al equilibria and

influence of ionic strength). Thus, employing Eq. [2.2] gives

log K = -6(3.47) 2(-3.78) 2(-3.53)

log K = -6.20

From Eq. [2.3] the intercept of the kaolinite stability line in fig. 2.2 is - 1/2 log K, or

3.10. The stability line in fig. 2.2 represents the boundary between the areas of

supersaturation (above) and undersaturation (below) for kaolinite. Kaolinite was

detected by X-ray diffraction analysis only in solutions saturated or supersaturated

with respect to kaolinite (solid data points). None of the data points indicated

supersaturation with respect to gibbsite. All of the undersaturated solutions and all of

the supersaturated solutions, except for one unexplained example, appear to be

progressing toward the kaolinite stability line (Dixon and Weed, 1989). By

calculating the free energy of formation, G, equivalent to the stability line for the

kaolinite derived above, comparison can be made with the stability of other kaolinite

subgroup minerals from Eq. [2.1]

Gr = G (kaolinite) - 2G3+Al - 2G H4Sio4 - GH2O . [2.4]

where, Gr is the standard free energy of reaction. And = Gr = - RT In K = 1.364

pK, or 35.4 kJ mol

-1

at 25C. Hence G (kaolinite) = -3.783 MJ mol

-1

.

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The G of -3.783 MJ mol-1 suggests a slightly greater stability for the synthetic

kaolinite than does the G of -3.78 1 MJ mol-1 for a kaolinitic of reportedly near

maximum crystallinity. The G values for kaolin precipitated in soils and sediments

at 25C probably will range from 3.760 MJmol-1 for halloysite to the initial

precipitation level of -3.783 MJ mol-1 for kaolinite (Kittrick, 1970)

2.3.3.2. Kaolinite formation from Hydroxy-Al Interlayered Montmorillonite

Kaolinite formation has been induced by hydrothermal treatment of hydroxy-Al

interlayered montmorillonite at 220 C. The kaolinite was expanded to approximately

0.85 nm during the first few days of reaction. After 14d of hydrothermal treatment or

heating at 300 C following hydrothermal treatment, the kaolinite collapsed to 0.72

nm. The early product had a spacing suggestive of halloysite or a halloysite-kaoliniteinterstratified mixture (Dixon and Weed, 1989).

Efforts to synthesize kaolinite from smectite have produced kaolinite-smectite

interstratified. Both tetrahedral and octahedral sheets of smectite were ought to

dissolve, releasing H4SiO4 for reaction with Al3+ to form kaolinite according to the

following equations (2.5 and 2.6): Tetrahedral and octahedral sheets of smectite

dissolve:

(Al,Fe,Mg)2Si4O10(OH)2 + 4H2O + 6H+ 2(Al,Fe,Mg)3+ + 4 H4SiO4 [2.5]

H4SiO4 and Al3+ react to form kaolinite

2H4SiO4 + 2Al3+ + H2O Al2Si2O5(OH)4 + 6H+ . [2.6]

The net reaction gives the Eq.2.7

(Al,Fe,Mg)2Si4O10(OH)2 + m Al3+ + 6H2O

Al2Si2O5(OH)4 + 6H

+

+ aFe

3+

+ bMg

2+

.. [2.7]Where m = 2 + a + b.

The addition of Al3+ contributed to the progress of the reaction and to kaolinite

production to a limit where smectite appeared to be stabilized. No amorphous silica

product was formed, and the pH declined during the progress of the reaction as

suggested by the equations. Some unreacted smectite remained, and some discrete

kaolinite was formed (Dixon and Weed, 1989).

2.4. Structure of clays minerals

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The term clay mineral is commonly used to refer to minerals with the sheet silicate

structures of the phyllosilicates. Phyllosilicates are one of many subclasses of the

silicate mineral class (Fig. 2.3). The basic structural feature of all minerals in these

subclasses is the SiO4 tetrahedron linked by the sharing of three of four oxy anions to

form sheets with a pseudohexagonal network in the ab crystallographic plane. The

sheet has the composition (Si2O5)2-. The three shared O2- ions are termed basal O2- ions

while the fourth apical O2- is not shared with another SiO4 tetrahedron and may bond

to other structures. The tetrahedra are interconnected in sheet-like structures. These

sheets are commonly referred to as silica or tetrahedral sheets (Fig. 2.3). Cations in

each tetrahedral sheet are in fourfold coordination, i.e., each Si cation is bonded to

four O2- ions arranged in each tetrahedron. Aluminum may replace as many as half of

the Si, producing a formula composition such as (AlSi3O10)5- or (Al2Si2O10)6-. Minerals

in which tetrahedral sheets of silica dominate are often referred to as layer silicates

(Olson et al. 2000).

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arranged around the positively charged cations in an octahedral pattern; one anion at

each corner of an octahedron surrounding a central cation. Adjacent octahedra share

anions to form a planar network referred to as an octahedral sheet (Schulze, 1989). In

the octahedral sheet, connections between each octahedron, M, to neighbouring

octahedra are made by sharing edges. The edge-shared octahedra form sheets of

hexagonal or pseudo-hexagonal symmetry (Fig. 2.3). Common tetrahedral cations are

Si4+, Al3+, and Fe3+. Octahedral cations are usually Al3+, Fe3+, Mg2+, and Fe2+, but other

cations, such as Li+, Mn2+, Co2+, Ni2+, Cu2+, Zn2+, V3+, Cr3+, and Ti4+ were identified.

Octahedra show two different topologies related to hydroxyl position, i.e., the cis- and

the trans-orientation (Fig. 2.3).

Phyllosilicates may be structurally classified into dioctahedral and trioctahedral types.This classification is based on the cation valence and configuration of the octahedral

site in two- or three-layer structures. There are two ways to fill an octahedral site,

depending on the cation valence. A divalent cation (Ca 2+, Mg2+) when placed into the

octahedral site produces a trioctahedral arrangement (Olson et al., 2000).

Phyllosilicates strongly influence chemical and physical properties of soils because of

their small particle size, high surface area, and cation exchange properties. Among the

phyllosilicate minerals important to soil development, the most common are discussedhere: mica, vermiculite, chlorite, smectite, interstratified minerals, kaolinite, and talc

and pyrophyllite. Depending on the arrangement of octahedral and tetrahedral sheets

the clay minerals can be grouped into 1:1, 2:1 and 2:1:1 type clay minerals (Olson et

al., 2000).

2.4.1. The 1:1 Type Minerals

The 1:1 layer structure consists of a unit made up of one octahedral and one

tetrahedral sheet, with the apical O2- ions of the tetrahedral sheets being shared with

(and part of) the octahedral sheet. There are three planes of anions (fig. 2.4). One

plane consistsof the basal O2 ions of the tetrahedral sheet, the second consists of O2

ions common toboth the tetrahedral and octahedral sheets plus OH- belonging to the

octahedral sheet,and the third consists only of OH- belonging to the octahedral sheet

(Dixon and Weed1989).

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Fig. 2.4. Models of a 1:1 layer structure. Oa, Ob, and Ooct refer to tetrahedral apical, tetrahedral basal,and octahedral anionic position, respectively. M and T indicate the octahedral and tetrahedral cation,respectively (Brigatti et al., 2006).

The common 1:1 type clay minerals will be described.

2.4.1.1. Kaolinite

Kaolinite occurs commonly in soils, often as hexagonal crystals with an effectivediameter of 0.2 to 2 m. Hydrogen bonding between adjacent unit layers prevents

expansion (swelling) of the mineral beyond its basal spacing of 7.2 (Fig 2.5).

Surface area is limited to external surfaces and hence is relatively small, ranging from

10 to 20 m2/g. Kaolinite is coarse clay with low colloidal activity, including low

plasticity and cohesion, and low swelling and shrinkage (Bhon et al., 1979). Kaolinite

is dioctahedral and the layer is electrically neutral. Other kaolin minerals, dickite and

nacrite, have the same basic 0.7-nm unit layer thickness but stacking sequences of

layers along the c axis differ (Olson et al., 2000).

Fig. 2.5. Schematic structure of Kaolinite Al2Si2O5 (OH)4 (Bear, 1964)

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The ideal unit formula for kaolinite [Al2Si2O5(OH)4] has an Si/Al ratio of 1 which

suggests little or no isomorphic substitution. Most of the 1 to 10 cmol kg -1 cation

exchange capacity of kaolinite has been attributed to dissociation of OH groups on

clay edges. However, if only one Si4+ of every 400 in the silica sheet were substituted

by Al3+, the net negative charge would be cmol kg -1. Chemical analysis of clay

systems is not sufficiently sensitive to prove or disprove this extent of isomorphic

substitution. Such limited substitution is not sufficient to alter the unit cell formula

given for kaolinite. The cation exchange capacity of kaolinite is highly pH-dependent

suggesting that isomorphic substitution is not the predominant source of charge (Bhon

et al., 1979).

Kaolinite in soils is both authigenic (formed in place) and allogenic (formedelsewhere). It can form from the weathering of primary and secondary minerals,

subsequent to the release of Si4+ and Al3+. In most soils, however, kaolinite is inherited

from the weathering of older sediments (Olson et al., 2000).

Kaolinite particles are often >1 m in size and have a platy or book-like appearance

when viewed with an electron microscope. Although not considered an expanding

clay mineral, kaolinite will swell with the intercalation of some small polar molecules

such as formamide, hydrazine or urea (Olson et al., 2000).

2.4.1.2. Halloysite

Halloysite is similar in structure to kaolinite except for a layer of water molecules that

is intercalated within the 1:1 layer. It differs in morphology from kaolinite in that it is

often tubular (Singh 1996). Fig 2.6 shows the schematic structure of Halloysite.

Halloysite has water molecules between each 1:1 layer and the ability to adsorb large

quantities of monovalent cations such as NH4+. Drying will cause the water molecules

to be removed and the clay layers to collapse together (Newman and Brown, 1966).

Halloysite is also characterized by a tubular morphology, whereas kaolinite, when

examined with microscopic techniques, has a platy structure (Sparks, 1995).

Halloysite is found primarily in soils of volcanic origin and typically weathers to

kaolinite. If not properly handled during sampling and analysis, it is difficult to

distinguish halloysite from kaolinite; once dehydrated, the layer spacing decreases to

close to that of kaolinite (Olson et al., 2000).

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Fig. 2.6. Schematic structure of Halloysite Si4Al4O10(OH)8.4 H2O (Bear, 1964)

2.4.2. The 2:1 type Minerals

The 2:1 layer structure consists of two tetrahedral sheets bound to either side of an

octahedral sheet. There are four planes of anions (fig. 2.4). The outer two planes

consist of the basal oxygens of the two tetrahedral sheets, while the two inner planes

consist of oxygens common to the octahedral sheet and one of the tetrahedral sheets,

plus the hydroxyls of the octahedral sheet (Dixon and Weed, 1989). The common 2:1

type clay minerals will be discussed.

Fig. 2.7. Models of a 2:1 layer structure. Oa, Ob, and Ooct refer to tetrahedral apical, tetrahedral basal,and octahedral anionic position, respectively. M and T indicate the octahedral and tetrahedral cation,respectively(Brigatti et al., 2006)

2.4.2.1. Smectite (Montmorillonite)

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The term smectite is a generic name for 2:1 clay minerals that swell or collapse easily

depending on their pressure potential, osmotic potential, and saturating cations

(Sumner, 2000). Smectites are 2:1 layer silicates with layer charge of 0.25 to 0.6 per

formula unit. Because of the relatively low layer charge, smectites are freely

expanding. The c-spacing varies with the exchangeable cation and the degree of

interlayer solvation. Complete drying yields a spacing of 9.6 , and full hydration can

swell the layer to distance of tens or even hundreds of Angstroms (Bhon et al., 1979).

Clay minerals in the smectite-saponite group are characterized by a layer charge of

0.2-0.6 per half-cell formula unit. The group includes the subgroups dioctahedral

smectites and trioctahedral saponites. The dioctahedral smectites are represented by

montmorillonite, beidellite, and nontronite. The ideal half-cell chemical formula for

montmorillonite is M0.33, H2OAl1.67 (Fe2+, Mg2+)0.33 Si4O10(OH)2, where M refers to a

metal cation in the interlayer space between sheets. The tetrahedral cations are Si 4+

and the octahedral cations are Al3+, Fe2+ and Mg2+ (Fig 2.8). One can calculate the

negative charge in the tetrahedral sheet as 0 and in the octahedral sheet as - 0.33.

Thus, the net negative charge is - 0.33 that is balanced by exchangeable cations

represented by M0.33. The other feature that characterizes montmorillonite is the

presence of water molecules in the interlayer space. This causes montmorillonite to

take on shrink-swell characteristics (Spark, 1995).

Fig. 2.8. Schematic structure of montmorillonite Nax(Al2-xMgx)Si4O10(OH)2 (Bear, 1964)

The major difference between montmorillonite and the other two dioctahedral

smectites, beidellite and nontronite is that isomorphous substitution in these minerals

occurs in die tetrahedral sheet (i.e., Al3+ Substitutes for Si4+) rather than in the

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M is most often Mg2+. In ambient conditions, 6 water molecules, arranged in two

sheets, form an octahedron around each Mg2+. Vermiculites saturated with Ca2+ or

Mg2+ and equilibrated at 54% relative humidity expands to a lesser degree than

smectites, i.e., X-ray diffraction (XRD) d-spacing are about 1.4 nm rather than 1.7 nm

(Olson et al., 2000).

Isomorphic substitution in the tetrahedral sheet yields an excess negative charge of

1.3. Isomorphic substitution in the octahedral sheet yields an excess positive charge of

0.6. So the net charge is -0.7 per unit formula. The layer charge in vermiculite gives

rise to a cation exchange capacity of from 120 to 150 cmol kg-1, which is considerably

higher than the exchange capacity of montmorillonite. As with montmorillonite, the

cation exchange capacity is only slightly pH-dependent. Vermiculite swells less thanmontmorillonite because of its higher layer charge. The mineral is nonswelling (with

a c-spacing of 10 ) when saturated with K+ or NH4+ ions. Such ions are commonly

termed fixed and cannot be exchanged with ordinary salt solutions. Total surface

areas of vermiculite, when not K+ or NH4+ saturated, range from 600 to 800m2/g

(Bhon et al., 1979).

2.4.2.3. Mica

Micas occur in almost any geologic environment and are abundant in many rocks

including shales, slates, phyllites, schists, gneiss, granites, and in sediments derived

from these rocks. As clay minerals, they can be derived from preexisting micas by

mechanical fragmentation but may also form in situ. Micas are 2:1 phyllosilicates

having a charge imbalance that is satisfied by a tightly held, nonhydrated, interlayer

cation (Olson et al. 2000). The idealized end member micas are dioctahedral

muscovite- K2Al2,Si6Al4O20(OH)4, and its isomorphous analogue paragonite

Na2Al2,Si6Al4O20(OH)4, and trioctahedral biotite- K2Al2Si6 (Fe2+, Mg)6O20(OH)4 and

phlogopite- K2Al2,Si6Mg6O20(OH)4. A variety of additional ion substitutions, beyond

those indicated for the end members, such as Li for Mg, more Al for Si, F or O for

OH, and possibly OH for O, can occur in micas. Layer charge, represented equivalent

to K, in the end-members, decreases in some biotites because of substitution of Al for

Mg, or Fe3+ for Fe2+ in the trioctahedral layer. The layer charge is extremely high in

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margarite-Ca2Na2Al2,Si6Al4O20(OH)4 which is representative of the brittle micas

(Bear, 1964).

Micas have an octahedral sheet sandwiched between two tetrahedral sheets (Fig.

2.10). One of the 4 tetrahedral sites typically contains Al3+

rather than Si4+

. Thisresults in excess negative charge per formula unit. The negative charge is balanced by

monovalent cations, usually K+, occupying interlayer sites between two

tetrahedraloctahedral- tetrahedral layers. The general formula is Mx(R(2+)3-y, R(3+)y) (Si4-

X,Alx)O10(OH)2 where M is usually K+ or Na+. Micas can be either dioctahedral or

trioctahedral. The mica group consists of many species because Fe2+ and Fe3+

substitute for Mg2+ and Al3+ in octahedral sheets and Na+ and Ca2+ may substitute for

K+

. Occasionally, Ba2+

, Rb2+

, Cs, Sr+

, or NH4+

also substitute for K+

, particularly inbiotite mica (Olson et al., 2000).

Fig. 2.10. Schematic structure of Mica K[Al2 (Si3Al)O10(OH)2] (Bear, 1964)

Although micas may be either dioctahedral or trioctahedral, isomorphic substitution in

the tetrahedral layer creates negative charge close to the unit layer surface. Thischarge strongly retains the K+ ion. Such interlayer K+ is so strongly adsorbed that it is

not exchanged in standard cation exchange capacity determinations. Thus despite the

relatively large layer charge (about -1 for many micas), the CEC is only 20 to 40 cmol

kg-1. Total surface area is about 70 to 120 m2/g and is restricted to external surfaces.

Micas are nonswelling and are only moderately plastic. So-called fixed K+ is released

slowly during weathering and is an indigenous source of K+ for many soils (Bhon et

al., 1979).

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Chlorites are often referred to as 2:1:1 clays since they are 2:1 clays with a hydroxide

interlayer, either gibbsite-like [Al(OH)x] or brucite-like [Mg(OH)x] where x is


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or (2) diagenesis. Similarly, Sawhney (1989) states that naturally occurring mixed

layer minerals form by

1. Hydrothermal alteration,

2. Weathering involving partial removal of interlayer K from mica or removal of a

hydroxide-interlayer from chlorite, or

3. Uptake of K and

4. Formation of a brucite-like or gibbsite-like interlayer in expanding layer silicates.

2.4.3.1. MontmorilloniteVermiculiteChlorite Intergrade

Precipitation of coatings of hydrous, positively charged sesquioxides on the negative

surfaces of layer silicates has long been recognized as a phenomenon characteristic of

soils, as has their fundamental influence on soil chemical properties such as phosphate

fixation. Precipitation of hydroxy-aluminum, hydroxy sesquioxides, and possibly

magnesium hydroxide as gibbsite-like (or brucite-like) structures in the interlayer

spaces of montmorillonite and vermiculite produces structures the properties of which

are intergradient between those of the expansible mineral and those of chlorite. Most

natural clay intergrades so far described have shown 14 diffraction spacing and

have been termed dioctahedral vermiculite, vermiculite, chlorite-like orchlorite, according to the end-member to which they are compared. The interlayers

are heterogeneous with respect to islands of the brucite-like structures or

gibbsitelike structures which are distributed in interlayer spaces otherwise filled

with water and exchangeable cations, as is characteristic of vermiculite or

montmorillonite. The cation exchange capacity is consequently decreased to the

extent of the positive charge of the nonreplaceable hydroxy cation interlayers. The

interlayer specific surface is also decreased. The interlayer space becomesincreasingly resistant to collapse by heating as hydroxy interlayers become more

extensive. The products have appropriately been designated as intergradient 2:1 to

2:2 layer silicates or intergradient montmorillonite-vermiculite-chlorite (Bear, 1964).

2.4.3.2. Swelling 2:1 to 2:2 intergrades

Swelling chlorite has been described as having one surface of the brucite-like layer

unattached to a silicate layer, the interlayer positions thus would be heterogeneous

with respect to brucite and water-cation layers, and the complex is a swelling 2:1 to

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2:2 intergrades. A sharp 18 peak with Mg saturation and glycerol solvation, and a

greatly broadened 10 peak with K saturation and 500 C heating, characterize the

expanding montmorillonite-chlorite intergrade type of clay (Bear, 1964).

In alkaline soils of arid regions, montmorillonite is relatively stable to weathering andso has a high frequency distribution. Characteristically, it is partially interlayered. The

weathering mechanism by which swelling intergradient 2:1 to 2:2

montmorillonitechlorites is formed appears to arise from the swollen character of

montmorillonitewhen moist. The presence of abundant montmorillonite creates moist

conditions and,as a result, the clay in the soil is kept swollen to 20 to 40 spacing.

Intercalatedhydroxyaluminum sheets can attach to only one layer of montmorillonite

so long as the layers are widely separated. In a feldspar rich mantle, alkalinity andhydrolysisyield a plentiful supply of freshly precipitated aluminum hydroxide, which

has apositive charge when below pH 8.3. The (+) valence charges on the hydroxyl

sesquioxide interlayer cations attach to (-) exchange sites. Laboratory syntheses also

have involved attachment of gibbsite interlayers to one surface while the

montmorillonite is in a swollen state. The cation exchange capacity was considerably

lowered by the introduction of hydroxyaluminum interlayers in synthetic preparations

(Bear, 1964).

2.5. Soil mineral weathering sequences

Soil mineral weathering has been studied for many decades; most of the available

information, however, is empirical in nature and cannot be used for quantitative

predictions. In this section some of the available empirical weathering sequences

along with the predictions based on thermodynamic data have discussed (Dixon and

Weed, 1989).

Although there are exceptions, minerals that are far removed from their formation

environment tend to be the least stable. Based on generalizations from accumulated

data, (Goldich, 1938) proposed a mineral stability sequence (Fig. 2.12) that supports

the above hypothesis, in which the least stable minerals (in an aqueous environment)

are the first to crystallize during cooling of magma, and the most stable minerals

crystallize last. Because weathering (in this case, essentially the rate of dissolution) is

highly dependent on the nature of the weathering environment, however, the main

drawback of this sequence is that it does not consider the weathering environment.

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Furthermore, secondary minerals that are weathering products of the preexisting

minerals are not considered (Dixon and Weed, 1989).

Fig. 2.12. Mineral stability sequence. Stability of minerals increases as one proceeds towards quartz,and as a group, salic minerals (plagioclases, feldspars, etc.) are more stable than mafic minerals(olivine, augite, etc.) (Goldich, 1938)

Jenny (1941) considered soil development to be a function of five factors: parent

material, topography, biosphere, climate, and time. The resultant weathering therefore

was considered to be a function of these factors.

Jackson and Sherman (1953) related the relative degree of soil development to the

types of minerals present in the clay fraction (Table 2.4). The explicit assumption in

this sequential development is that there is a continuous leaching of the soil profile

with time. Although, in general, the sequence is applicable to all soils and can be used

to categorize the soils according to their degree of development, this sequence cannot

be used to determine the specific chemical environmental conditions needed to bring

about these changes, or whether a given soil under a specific weathering environment

will proceed through all stages.

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Table 2.3

Types of minerals as indicators for relative degree of soil development (the higher the number, thehigher the degree of development

Relative degree of

soil development

Prominent minerals in soil clay fraction

1 Gypsum, and other soluble salts

2 Calcite, dolomite, and apatite

3 Olivine-hornblende minerals

4 Biotite, glauconite, ferromagnesian chlorite

5 Feldspars

6 Quartz

7 Muscovite-illite

8 Interstratified 2:1 layer silicates and vermiculite

9 Montmorillonites

10 Kaolinite and halloysite

11 Gibbsite and allophane

12 Hematite-goethite

13 Anatase-leucoxene

Source: Jackson and Sherman, 1953

Numerous other researchers have identified the weathering products of various rocksand minerals and have proposed empirical weathering sequences. In general, it is

recognized that the weathering of a mineral is affected by its composition, coefficients

of expansion, cleavage, and original defects in the crystals, hardness, and specific

surface. Additional external factors are the physical, biological, and chemical

(oxidation/reduction, hydration, hydrolysis, pH, chelation, cation exchange,

carbonation) conditions of the weathering environment (Dixon and Weed, 1989).

The rate of weathering of primary minerals (e.g., 80,000 and 8800 years) of mean

lifetime at 25C and pH 5 for 1-mm crystals of albite and enstatite, respectively, as

well as the rate of formation of crystalline clay minerals, is so slow that the bulk

mineralogy of any given soil is not expected to change considerably in the lifetime of

any given individual. Therefore, it is easy to analyze the sample mineralogically and

find the exact distribution of the minerals. In the case of trace elements or hazardous

elements, however, the type or sequence of mineral formation is very important

because it may control the concentrations of these elements in groundwaters (Dixonand Weed, 1989).

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Thermodynamic data can be used to predict the formation and weathering of minerals,

and perhaps the sequence of mineral weathering, because the of a mineral is

dependent on the composition, crystallinity, nature of chemical bonds, etc., and

because the precise weathering environment can be quantified in terms of activities of

constituent ions (Dixon and Weed, 1989).

Most major soil minerals contain one or more elements in common. Rarely, if ever,

can these minerals all be in equilibrium with one another. Within the group competing

for common elements, individual minerals are usually stable (least soluble) over a

rather narrow range of solution compositions. Thus, it is possible to understand the

soil system over a small portion of its solution composition range by considering the

stabilities of only a few of its minerals (e.g., gibbsite, kaolinite, montmorillonites, andamorphous silica) (Dixon and Weed, 1989).

The use of thermodynamic data in making graphs to portray mineral weathering

sequences or stability sequences has been discussed by many authors (Garrels and

Christ, 1965). As an example, the following reaction describes the equilibrium

between kaolinite and its dissolved species:

Al2Si2O5(OH)4 + 6H+ 2Al3+ + 2H4SiO40 + H2O

Kaolinite and H2O have unit activity and thus can be ignored in the equilibrium

expression. The equilibrium constant (K) for the above reaction is therefore

log k = 2 log (Al3+) + 2 log (H4SiO40)- 6 log (H+) .. [2.8]

Simplifying, rearranging, and substituting the above equation (Eq. 2.8), eq.2.9 derived

log (Al3+) + 3pH = 3.4 log (H4SiO40) . [2.9]

This algebraic manipulation of the equilibrium constant (Eq. 2.9) permits a graphic

representation of this solubility reaction, as shown in fig. 2.13. Similarly, gibbsite,

amorphous silica, and montmorillonite solubilities are also plotted in fig. 2.13.

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Fig. 2.13. Relative solubility of a few clay minerals. Solubility lines for Belle Fourche montmorilloniteassume equilibrium with hematite and Na as the exchangeable ion when log Na = -3. The data for SiO2(am) are from Kittrick (1969); for gibbsite, from Kittrick (1966a); for kaolinite, from Kittrick (1966b);and for Belle Fourche montmorillonite, from Kittrick (1970)

Fig. 2.13 is a two-dimensional graph with a two-ion parameter on the ordinate and a

single illustrative contour line. Three ion variables are shown in this graph, with pH

and pMg (-log Mg2+) fixed at selected levels of interest (pMg 3.7 is the average

content of natural waters) and pFe3+ assumed to be fixed by hematite equilibrium and

pH. Only four minerals are considered explicitly, yet the graph appears to explain the

formation of gibbsite, kaolinite, and montmorillonite, and their control of Al 3+ as a

function of H4SiO40. As the H4SiO40 concentration increases, first gibbsite, then

kaolinite, and finally montmorillonite become, in turn, the most stable mineral of the

group (Dixon and Weed, 1989).

The most stable mineral also is the least soluble. At a log (H 4SiO2) activity of -3.0, the

montmorillonite line (pH 6, pMg 3.7) is below the metastable extension of the logAl3+

+ 3 pH kaolinite stability line (dashed), which is below the metastable (dashed)

extension of the gibbsite stability line. Because the pH of the system is fixed, the

lowest value of log Al3+ + 3pH is the lowest value of Al3+. Gibbsite supports almost

10,000 times as much Al3+ as montmorillonite at log (H4SiO40) of -3.0 and pH 7.0. At

log (H4SiO40) of -6.0, their relative solubilities are reversed (Dixon and Weed, 1989).

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The two montmorillonite stability lines (contours) in fig. 2.13 indicate that

montmorillonite becomes more stable as the pH increases and as pMg decreases. This

increased montmorillonite stability comes at the expense of kaolinite, whose stability

range is correspondingly restricted. The point where the montmorillonite and kaolinite

stability lines intersect defines the chemical conditions under which these two

minerals can be in equilibrium with each other. Montmorillonite-amorphous silica and

kaolinite-gibbsite are also compatible pairs. Incompatible pairs are amorphous

silicagibbsite and montmorillonitc-gibbsite (unless solution conditions are such that

montmorillonitc becomes sufficiently stable to eliminate the kaolinite stability field

entirely). Perhaps the most important insight to be gained from fig. 2.13 is an

appreciation that a variety of minerals can be formed in soils in response to a variety

of soil-solution environments and that the chemical weathering environment can be

quantified (Dixon and Weed, 1989).

The relative solubility of selected primary and secondary minerals is Plotted in fig.

2.14. It should be emphasized that the accuracy of the stability of a given mineral

depends on the conditions assumed for the diagram as well as on the accuracy of the

thermodynamic data used in the calculations. These diagrams can be extended to

include other minerals for which thermodynamic data are available and other

environmental conditions not considered in fig. 2.14. In these diagrams, the mineral

that maintains the lowest log Al3+ + 3pH activity at a given H4SiO40 activity is the

most stable. For example, at log (H4SiO40) of -7.0, the minerals in order of increasing

stability are Na-glass, K-glass, vermiculite, low albite, Ca-glass, analcime (same as

Ca-glass), microcline, pyroxene, montmorillonite, anorthite, chlorite (same as

anorthite), kaolinite and gibbsite (Dixon and Weed, 1989).

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Fig. 2.14. Relative solubility of selected primary and secondary minerals at pH 7.0 when log Mg2+= logNa+ log K = -3, log Ca2+ = -2.5, and log Fe3+ is in equilibrium with hematite (Kittrick, 1970)

The stability sequences based on thermodynamic data depend on the chemical

conditions in the weathering environment, such as H4SiO40 activity, pH, and the

precise activities of other ions; by contrast, the empirical sequences are fixed, and

thus changes in weathering environment cannot be taken into consideration. In

general, however, the thermodynamic data support the empirical observations, namely

(Dixon and Weed, 1989):

1. The primary minerals dissolve to form clay minerals

2. Montmorillonites are stable in solutions with high H4SiO40 activity (poorly drained

conditions)

3. Kaolinites are stable in solutions of moderate H4SiO40 activity (well drained) and

4. Gibbsite is stable only under conditions where H4SiO40 activity is very low

In addition to the relative stabilities of minerals, information regarding environmental

conditions needed for the possible formation of these minerals also can be obtained

from the stability diagrams.

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2.6. Pridiction of clay mineralogy

The ratio of CEC by ammonium acetate at pH 7 to clay can be used to predict clay

mineralogy. If the ratio is multiplied by 100, the product is cmol(+)/kg clay (Soil

Survey Staff, 2006). The typical ratio for some classes of clay mineralogy arementioned in Table 2.4

Table 2.4

The typical ratio of CEC to clay for some classes of clay mineralogy

Source: Soil Survey Staff, 2006

These ratios are more valid when some detailed mineralogy data are available. If the

ratio of 1500 kPa water to clay is 0.25 or less or 0.6 or more, the ratio of CEC by

ammonium acetate to clay is not valid. Ratios of 1500 kPa water to clay of 0.6 or

more are typical of poorly dispersed clays, andic and spodic materials with an isotic

mineralogy class, and ratios of less than 0.3 are common in some soils that contain

large amounts of gypsum ((Soil Survey Staff, 2006)).

A ratio of CEC at pH 8.2 to 1500 kPa water of more than 1.5 and more exchange

acidity than the sum of bases plus KCl extrctable Al imply a soil with a high pH-

dependent charge. Along with bulk density data, they help to distinguish soils that

have andic and spodic materials or soils that have materials with an isotic mineralogy

class from soils with materials that are more crystalline (Soil Survey Staff, 2006).

2.7. Ganges Floodplain soils of Bangladesh

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Typical ratio of CEC to clay classes of clay mineralogy

0.7 Smectite


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Bangladesh is apparently flat alluvial plain with apparent homogenous land and soil

condition and it, in fact, has a complex land pattern and soil condition. It has three

distinct land system: Floodplain, Terrace and Hill (Rahman, 2005).

Huizing (1971) made a reconnaissance study on the sand mineralogy of Bangladeshsoils and reported that hill and terrace soils Contained low feldspathic and micaceous

minerals while floodplain soils were rich in feldspar, mica and amphiboles.

Hasibullah et al. (1971) studied the clay mineralogy of Nunni series of_Northern and

Eastern Piedmont Plains, Borda series of Middle Meghna River Floodplain and

Noadda and Chhiata series of terrace areas. They reported that all these soils were

dominated by mica and kaolinite except for Borda series. Vermiculite was dominant

in the two floodplain soils while randomly interstratified minerals dominated the twoterrace soils.

Hassan and Razzq (1981) studied four soil profiles from the Sunderban forest area on

Ganges Tidal Floodplain and found that all soils were dominated by mica and

smectite with some katolinite, chlorite; vermiculite and interstratified minerals. white

(l985) reported qualitative estimates of clay mineralogical composition of a large

number of soils and opined that mica (muscovite) and kaolinite were the predominant

minerals in the clay fraction of most floodplain soils. He speculated that mica wastransformed to vermiculite under acidic condition, in the BrahmaPutra, Meghna and

Tista floodplain soils; while under neutral to alkaline reaction and poor drainage

conditions mica was transformed tosmectite (montmorillonite). He also reported the

occurrence of halloysite in the surface and smectite in the lower horizons of some

terrace soils.

Saheed (1985) reported three groups of mineralogical association in Bangladesh soils

mica, vermiculite and kaolinite in most floodplain soils; smectite along with mica inGangeas floodplain soils; and mica and halloysite in terrace soils. Mineralogy of silt,

and fine and coarse clays of four soil profiles was studied by Islam and Lotse (1986).

Using X-ray dilfraction (XRD), ion exchange and selective dissolution techniques.

Mica and smectite were found to be dominant in Batra and Ghior series of Ganges

River Floodplain, whereas mica and kaolinite were dominant in Naraibag and Ghatail

series of Old Meghna Estuarine Floodplain and Old BrahmaPutra Floodplain,

respectively.

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than in calcareous soils; calcareous soils contained smectite in large quantities while

noncalcareous soils were rich in vermiculite. Mazumder (1996) reported the clay

mineralogical composition of seven soil profiles of Brahmaputra floodplain and found

that all soils were rich in mica, kaolinite and vermiculite while chlorite was identified

in some soils.

Aramaki (1996) studied 11 soil profiles for the clay mineralogical composition and

reported that mica was dominant in all soils, smectite in Ganges floodplain soils,

interstratified mica-vermiculite-smectite and interstratified kaolinite-smectite or

kaolinite in terrace soils. kaolinite was rich in hill soils and even greater than mica,

while Old Brahmaputra and Meghna floodplain soils contained good amounts of

chlorite and vermiculite. Khan et al. (1997) studied five benchmark sail profiles fromdifferent floodplain of Bangladesh and found that mica was the dominant mineral in

all soils. Smectite was dominant in Ganges floodplain soils only. Varing amounts of

kaolinite and vermiculite were present in all soils.

Idenification of kaolinite as a dominant mineral in most floodplain soils was a matter

of great controversy among the reports. Mia (1990), White (1985), Saheed (1985), Ali

(1994), Islam and Lotse (1986), Mazumder ( 1996), and Khan et al. (1997) reported

high quantity of kaolinite in different floodplains soils while others found smalleramounts of kaolinite in those soils.

The clay mineralogical map of Bangladesh (Fig. 2.15) was prepared by Hussain et al.

(1999).

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Fig. 2.15 A clay mineralogical map of Bangladesh soils. Dots indicate location of soil sampling sites.

Legends: 1. mica-chlorite suite; 2. mica-smectite suite; 3. mica- vermiculite- kaolinite suite; 4. mica-

chlorite- vermiculite suite; 5. mica -mixed-layer-minerals- kaolinite suite; 6. kaolinite -mica suite; 7.

mica kaolinite- vermiculite suite; and 8. mica-kaolinite- vermiculite suite ( Hussain et al., 1999).

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3. Material And Methods

A study was conducted to asses the mineralogy of some soils of Ganges river

floodplain area in the south-western region of Bangladesh.

3.1. Study area

The samples were collected from different locations of Jessore and Satkhira district.

General information of sampling sites are presented in Table 3.1

Table 3.1 General information of sampling sites

Sample No. Address Physiography Agricultural Land Use

1

Village: Mohakal

Union: Noapara

Thana: Avoynagar

District: Jessore

Ganges TidalFloodplain Sesbania

2

Village: Ghoragasa

Union: Norendropur

Mouza: Ghoragasa

District: Jessore

Ganges MeanderFloodplain

Vegetables-Ginger-Tumeric-Fallow

3

Village: Ramnagar

Union: Ramnagar

Mouza: Sutigata

District: Jessore


Teak plants

4

Village: TeghoriUnion: Deora

Mouza: Teghori

Thana: Kotowali

District: Jessore


Amon-Jute -Pulses

5

Village:Meghla

Union:Chachra

Thana:Kotowali

Dist:Jessore


Fallow

6

Village: Navaron

Union: Navaron

Mouza: Jhikorgacha

Thana: Jhikorgacha

District: Jessore


Marigold

7

Village: Jhapaghat

Union: Helatola

Mouza: Brozobaksha

Thana: kolaroa

District: Satkhira


Jute-Potato-Rice

8 Village: Jhapaghat

Union: Helatola

Mouza: BrozobakshaThana: Kolaroa


Vegetables

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District: Satkhira

3.2. Collection of soil samples

A total of 8 soil samples were collected for laboratory analysis. Representative soil

samples were collected from surface only. The samples were placed in plastic bags.

3.3. Processing of soil samples

The collected soil samples were dried in air by spreading on separate sheet of papers

after it was transported to laboratory. After drying in air, the larger aggregates were

broken gently by crushing in a wooden hammer. A portion of the crushed soil was

passed through a 2.0 mm sieve. The sieved soil were then preserved in plastic bag and

labeled properly. These were later used for various chemical analyses. The chemical

analyses were carried out on those collected soil samples in the laboratory of the Soil

Science discipline in Khulna University.

3.4. Soil analyses

Different parameters of soil were analyzed by following the available procedures.

3.4.1. Particle size analysis

The particle size was determined by hydrometer method presented by Gee and Bauder

(1986).

3.4.2. pH

Soil pH (1:2.5) was determined electrochemically by using of glass electrode with the

procedure of Jackson (1973).

3.4.3. ECe

The ECe of the soil was measured at a soil water ratio of 1:1 by EC meter (USDA,

2004)

3.4.4. Organic carbon

Soil organic carbon was determined by of Wakley and Blacks (1934) wet oxidation

method as described by Jackson (1973).

3.4.5. Cation exchange capacity

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CEC was determined by method as described by Jackson (1973); extracting the soil

with 1N NH4OAc solution of pH 7 followed by replacing ammonium from the

exchange site by 2N KCl (Jackson, 1973).

3.4.6. Exchangeable K+

For the determination of exchangeable K+ soil was extracted with 1N NH4OAC

solution of pH 7.0 as described by Piper (1950) and Jackson (1973).

3.4.7. Water soluble K+

Water soluble K+ was determined after extraction with distilled water as described by

Jackson (1973).

3.4.8.

Water soluble Mg2+

Water soluble Mg2+ was determined in a combination determination process of

Ca2+and Mg2+, after extraction with distilled water as described by Piper (1950).

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4. Result and Discussion

The results obtained from laboratory analysis will be discussed in this section.

4.1. Physical Characteristics

The texture of the samples varied from loam to clay loam (Table 4.1). This result

resembles to several findings by other researches. SRDI staff (1965-86) found that on

the young Ganges meander floodplain, soils of ridges and inter-ridges depressions are

silt loam to silty clay; and on the old Ganges meander floodplain, soils of ridges and

depressions are loamy to clay in texture.

Highest value of sand percentage was observed in sample no. 1 (51%) and lowest

value of sand percentage was in sample no. 2 (19%) (Table 4.1). Highest value of siltpercentage was in sample no. 5 (60%) and lowest in sample no. 6 (38%) (Table 4.1).

Highest value of clay percentage was observed in sample no. 6 and 7 (Table 4.1) and

lowest in sample no. 1 (10%) (Table 4.1).

4.2. Chemical Characteristics

4.2.1. pH

Among the samples pH values were varied from 5.9 to 7.22 (Table 4.1). pH 5.9 was

found in sample no. 4 whereas 7.22 was found in sample no. 7.

4.2.2. EC

EC values were varied from 0.59 dS/m (sample no. 5) to 3.84 dS/m (sample no. 1)

among the samples (Table 4.1).

4.2.3. CEC

CEC values were varied from 42.56 cmolc/Kg (sample no. 7) to 10.92 cmolc/Kg soil

(sample no. 3) among the samples (Table 4.1).

4.2.4. K+ content

Water soluble K+ content were varied from 0.04 cmolc/Kg (sample no. 6) to 0.10

cmolc/Kg (sample no. 8) among the samples (Table 4.1).

Exchangeable K+ content were varied from 0.04 cmolc/Kg (sample no. 5) to 0.32


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Table 4.1 Different properties of studied soils

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Soil Property 1 2 3 4 5 6 7 8

Physical PSA

Sand (%) 51 19 27 28 27 27 21 24

Silt (%) 39 61 58 48 60 38 44 42

Clay (%)10

20 15 24 13 35 35 34

Texture LoamSilt

loamSilt

loamLoam

Siltloam

Clayloam

Clayloam

Clayloam

Chemical

pH 6.9 6.1 7.1 5.9 6.76 6.67 7.22 7.02

EC (dS/m) 3.84 0.69 0.94 0.77 0.59 0.73 0.98 1.37

CEC (cmolc/Kg) 19.10 18.75 10.92 15.04 17.04 31.91 42.56 27.17

K content(cmolc/Kg)

Exchang-eable

0.11 0.10 0.20 0.12 0.04 0.05 0.32 0.23

Watersoluble

0.07 0.05 0.07 0.05 0.03 0.04 0.07 0.10

Mg content(cmolc/Kg)

Watersoluble

1.43 0.41 0.34 1.19 0.58 1.50 1.77 0.72

Organic C (%) 0.78 0.62 0.64 0.60 0.53 0.41 0.90 0.47

Humus (%) 1.34 1.08 1.11 1.04 0.91 0.71 1.55 0.81


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4.2.5. Mg2+ content

Water soluble Mg2+ content were varied from 1.77cmolc/Kg (sample no. 7) to 0.34


4.2.6. Organic C and humus

Highest value of Organic C (0.90%) and humus (1.55%) was found in sample no. 7

and lowest value of these was 0.41% and 0.71%, respectively in sample no. 6 (Table

4.1).

4.3. Prediction of clay minerals

The clay mineralogy of the soil samples was predicted from the ratio of CEC by

ammonium acetate at pH 7 to clay (Soil Survey Staff, 2006) and the result ispresented in Table 4.2.

Table 4.2

The clay mineralogy of the soil samples

Samples No. Ratio of CEC to clay Probable Clay Mineralogy

1 1.91 Smectitic

2 0.94 Smectitic

3 0.73 Smectite

4 0.63 Mixed or Smectitic5 1.31 Smectitic

6 0.91 Smectitic

7 1.22 Smectitic

8 0.799 Smectitic

4.4. Weathering stability of clay minerals

According to Faure (1992), pedological indices for studied soil samples are shown in

Table 4.3.

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principle the following equilibrium reactions of [Eq. 4.1] Illite Muscovite will shift

to left hand side. As a result Illite can form in the environment.

1.3 [(Al1.80 Mg0.25) (Si3.50 Al0.50) O10 (OH)2]K0.60 + 0.22K+ + 0.43H+

KAl2Si3AlO10(OH)2 + 1.55 SiO2 (amorphous) + 0.325 Mg2+ + 0.5H2O . [4.1]

In all soil samples, the highest value of log [Mg2+]/[H+]2envwas 12.39 and lowest was

9.57(Table 4.3) which were higher than the value of log [Mg2+]/[H+]2eqi (6.42).

According to Le Chatelirs principle the following equilibrium reactions [Eq. 4.2] of

Kaolinite Mgmontmorillonite will shift to right hand side and as a result formation

of Mg-montmorillonite may be possible in environment.

7 Al2Si2O5(OH)4 + 8 SiO2 (amorphous) + Mg2+

6[(Al2.00) (Si3.67 Al0.33)O10(OH)2] Mg0.167 + 7H2O + 2H+ .. [4.2]

The values of log [Mg2+]/[H+]2envin all soil samples were varied from 12.39 to 9.57

(Table 4.3) whereas the value of log [Mg2+]/[H+]2eqiwas in Faure (1992) equilibrium

equations was 3.29. So the following equilibrium reaction [Eq. 4.3] of Illite Mg-

montmorillonitewill shift to right hand side.

1.013[(Al1.08Mg0.25)(Si3.50Al0.50)O10(OH)2]K0.60 + 0.1245SiO2 (amorphous) + 0.781H+

[(Al2.00) (Si3.67 Al0.33)O10(OH)2] Mg0.167 + 0.086 Mg2+ +0.405 H2O [4.3]

Among all the soil samples, the values of log [Mg 2+]/[H+]2env varied form 12.39 to

9.57 (Table 4.3). according to Faure (1992), the equilibrium value of log [Mg 2+]/

[H+]2eqi was 11.36 (Table 4.4). The higher log[Mg2+]/[H+]2env value indicates the

possibility of formation of Chlorite from Mg-montmorillonite. The values of log

[Mg2+]/[H+]2env was higherthan 11.36 in sample no. 1 (11.65), 7 (12.39) and 8 (11.60).

The equilibrium reaction [Eq. 4.4] of Mg-montmorillonite Chlorite is[(Al2.00)(Si3.67 Al0.33)O10(OH)2] Mg0.167 + 5.658 Mg2+ + 9.32 H2O

1.165 Mg5Al2Si3O10(OH)8 + 0.175 SiO2 + 11.316 H+ .. [4.4]

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5. Summary and Conclusion

The physiography of studied soil samples were both Ganges Tidal and Meander

floodplain. The landscape was nearly leveled basin and poorly drained. The mineralogical

property and weathering stability of minerals were studied in laboratory.

From the ratio of CEC to clay percentage the clay minerals presented in the

soils were predicted. The dominant clay mineral in almost all the soils were

Smectitic. Only Teghori soil (sample no. 4) contained mixed or Smectitic

dominant clay mineral.

Almost all the studied soils had the tendency to transform into Mg-

montmorrillonite. In illite-muscovite system, illite is more stable than

muscovite. Muscovite was more stable than illite in Jhapaghat soil (sample no.

7 and 8). In Mg-montmorrillonite-chlorite system, chlorite was more stable

than Mg-montmorrillonite for Jhapaghat soil (sample no. 7 and 8). Different

cropping pattern had relatively same weathering stability of the minerals, thus

it can be said that stability of minerals was not affected significantly by

cropping pattern of the soils.

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5. References

Adams, W.A.; Evans, L.J. and Abdulla H.H. 1971. Quantitative pedological studies

on soil derived from Silurian mudstones. . Laboratory and in situ

weathering of chlorite. Soil Science, 22: 158-165.

Alam, M.L. 1994. Studies on Clay Mineralogical Characteristics and Their Effects on

K Release and Fixation Behaviors of Bangladesh Soils. Dissertation to The

United Graduate School of Agriculture Sciences, Kagoshima University, pp.

151.

Alam, M.L., Miyauchi, N. and Shinagawa, A. 1993a. Study on clay mineralogical

characteristics of calcareous and noncalcareous soils of Bangladesh. Clay

Science. 9: 81-97.

Alam, M.L., Miyauchi, N. and Shinagawa, A. 1993b. Study on clay mineralogical

characteristics of hill and terrace soils of Bangladesh. Clay Science. 9: 109-

121.

Ali, M.L. 1994. Effect of Alternate Wetting and Drying Cycles on Pedogenic

Processes of Some Representative Bangladesh Soils. Ph.D. Thesis, University

of Bangladesh, Dhaka, Bangladesh.Armaki, K. 1996. Mineralogical Composition of Main Bangladesh Soils. Graduation

Thesis, Faculty of Agriculture, Kyushu University, pp. 38.

Bassett, W.A. 1959. The origin of vermiculite deposits at Libby, Montana. American

Mineralogist, 44: 282-299.

Bear, F.E. 1964. Chemistry of The Soil. 2nd Edition. Oxford and IBH Publishing Co.

Pvt. Ltd., New Delhi.

Bhon, H.L.; McNeal, B.L. and OConnor, G.A. 1979. Soil Chemistry. John Wiley

and Sons. Inc., Canada.

Borchardt, G.A.; Jackson, M.L. and Hole, F.D. 1966. Expansible layer silicate genesis

in soils deplicted in mica pseudomorphs. In: Heller, L. and Weiss, A. (eds.),

Israel Program of Scientific Translations. Proceedings of the International

Clay Conference, Volume 1, 20-24 June 1966, Jerusalem,

Prediction of Mineralogy of Some Soils of Ganges Floodplain Under Different Cropping Patterns

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