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- Herausgeber: John Wiley & Sons
- Kategorie: Wissenschaft und neue Technologien
- Sprache: Englisch
Principles and Practice of Soil Science, Fourth Editionprovides a current and comprehensive introduction to soil science for students in the fields of environmental and agricultural science, ecology, soil and land management, natural resource management and environmental engineering.
- Covers all aspects of soil science including soil habitat, processes in the soil environment and soil management.
- Emphasizes the applications of soil science to the solution of practical problems in soil and land management.
- Highlights real world examples drawn from the author’s international experience in the field.
- Includes an expanded colour section of soil profiles and other features, and greater coverage of international soil classification
- Features new problem sets and questions at the end of each chapter, designed to reinforce important principles. An answer key is provided at the end of the text.
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Contents
Preface to the Fourth Edition
Units of Measurement and Abbreviations used in this Book
Part 1 The Soil Habitat
1 Introduction to the Soil
1.1 Soil in the making
1.2 Concepts of soil
1.3 Components of the soil
1.4 Summary
2 The Mineral Component of the Soil
2.1 The size range
2.2 The importance of soil texture
2.3 Mineralogy of the sand and silt fractions
2.4 Mineralogy of the clay fraction
2.5 Surface area and surface charge
2.6 Summary
3 Soil Organisms and Organic Matter
3.1 Origin of soil organic matter
3.2 Soil organisms
3.3 Changes in plant remains due to the activities of soil organisms
3.4 Properties of soil organic matter
3.5 Factors affecting the rate of organic matter decomposition
3.6 Summary
4 Peds and Pores
4.1 Soil structure
4.2 Levels of structural organization
4.3 Soil micromorphology
4.4 The creation and stabilization of soil structure
4.5 Soil porosity
4.6 Summary
Part 2 Processes in the Soil Environment
5 Soil Formation
5.1 The soil-forming factors
5.2 Parent material
5.3 Climate
5.4 Organisms
5.5 Relief
5.6 Time
5.7 Summary
6 Hydrology, Soil Water and Temperature
6.1 The hydrologic cycle
6.2 Properties of soil water
6.3 Infiltration, runoff and redistribution of soil water
6.4 Soil water retention relationship
6.5 Evaporation
6.6 Soil temperature
6.7 Summary
7 Reactions at Surfaces
7.1 Charges on soil particles
7.2 Cation exchange
7.3 Anion adsorption and exchange
7.4 Particle interaction and swelling
7.5 Clay–organic matter interactions
7.6 Summary
8 Soil Aeration
8.1 Soil respiration
8.2 Mechanisms of gas exchange
8.3 Effects of poor soil aeration on root and microbial activity
8.4 Oxidation–reduction reactions in soil
8.5 Summary
9 Processes in Profile Development
9.1 The soil profile
9.2 Pedogenic processes
9.3 Freely drained soils of humid temperate regions
9.4 Soils of the tropics and subtropics
9.5 Hydromorphic soils
9.6 Salt-affected soils
9.7 Summary
10 Nutrient Cycling
10.1 Nutrients for plant growth
10.2 The pathway of nitrogen
10.3 Phosphorus and sulphur
10.4 Potassium, calcium and magnesium
10.5 Trace elements
10.6 Summary
Part 3 Soil Management
11 Maintenance of Soil Productivity
11.1 Traditional methods
11.2 Productivity and soil fertility
11.3 Soil acidity and liming
11.4 The importance of soil structure
11.5 Soil erosion
11.6 Summary
12 Fertilizers and Pesticides
12.1 Some definitions
12.2 Nitrogen fertilizers
12.3 Phosphate fertilizers
12.4 Other fertilizers including micronutrient fertilizers
12.5 Plant protection chemicals in soil
12.6 Summary
13 Problem Soils
13.1 A broad perspective
13.2 Water management for salinity control
13.3 Management and reclamation of salt-affected soils
13.4 Soil drainage
13.5 Summary
14 Soil Information Systems
14.1 Communication about soil
14.2 Traditional classification
14.3 Soil survey methods
14.4 Soil information systems
14.5 Summary
15 Soil Quality and Sustainable Land Management
15.1 What is soil quality?
15.2 Concepts of sustainability
15.3 Sustainable land management
15.4 Summary
Answers to questions and problems
Index
Dedication
This book is dedicated to my wife Esme Annette White without whose support and encouragement it would not have been completed.
© 1979, 1987, 1997, 2006 by Blackwell Science Ltd, a Blackwell Publishing company
BLACKWELL PUBLISHING350 Main Street, Malden, MA 02148-5020, USA9600 Garsington Road, Oxford OX4 2DQ, UK550 Swanston Street, Carlton, Victoria 3053, Australia
The right of Robert E. White to be identified as the Author of this Work has been asserted in accordance with the UK Copyright, Designs, and Patents Act 1988.
All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted, in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, except as permitted by the UK Copyright, Designs, and Patents Act 1988, without the prior permission of the publisher.
First edition published 1979
Second edition published 1987Third edition published 1997
Fourth edition published 2006
1 2006
Library of Congress Cataloging-in-Publication Data
White, R. E. (Robert Edwin), 1937–
Principles and practice of soil science : the soil as a natural resource / Robert E. White. – 4th ed.
p. cm.
Includes bibliographical references.ISBN-13: 978-0-632-06455-7 (pbk. : alk. paper)ISBN-10: 0-632-06455-2 (pbk. : alk. paper)1. Soil science. I. Title.
S591.W49 2006
631.4–dc22
A catalogue record for this title is available from the British Library.
For further information onBlackwell Publishing, visit our website:
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Preface to the Fourth Edition
Dr Samuel Johnson is reputed to have said ‘what is written without effort is in general read without pleasure’. This edition of Principles and Practice of Soil Science has certainly taken much effort to complete, so I hope it will be enjoyed and provide valuable information to as wide an audience of interested readers as possible. The people I would expect to be interested in learning more about soils are not only soil scientists and others concerned with production systems, but also the various scientists and natural historians who are concerned about Earth’s ecology in its broadest sense.
At the time of the third edition (1997) I wrote about the ‘new generic concept’ of ecologically sustainable development (ESD) that was being promoted by international agencies and appearing with increasing frequency in government policy documents. However, through the 1990s and into the early years of the 21st century, more has been written about ESD than has been achieved on the ground in implementation of the policy. I have expanded on the topic of ‘sustainability’ in Chapter 15, drawing particularly on examples in Australia where a relatively fragile landscape continues to be put under pressure from ‘development’. The largest areas affected are rural areas, especially in the better watered coastal zone and the expanding irrigation regions, and areas of urban concentration (mainly along the coasts also). In this context, the quality and quantity of water have become key issues attracting much public and political attention. In recent years in Australia, these twin issues have become enmeshed with the question of climate change – by how much is it changing and where, and what are the possible positive and negative effects – which is directly linked to the emission of greenhouse gases from natural and human-influenced systems. Underlying these issues is soil behaviour because virtually all the precipitation that falls on land interacts with soil in some way. Hence, knowledge of the spatial distribution of different soil types and the pathways of water, with their associated physical, chemical and biological processes, in these various soil types becomes a very important component of land and water management. We need to be aware that everyone lives in a catchment and that the quality of life in that catchment depends on individual and collective human activities in that catchment. I have expounded on this subject in my 2003 G. W. Leeper Memorial Lecture ‘What has soil got to do with water?’, which is available on the Australian Society of Soil Science Inc. website (<wwwc>http://www.asssi.asn.au/asssi/flash/). Important tools for use in unravelling the complexity of water, energy and nutrient fluxes in catchments are models of the biophysical processes, incorporating a digital elevation model (DEM) and digital soil map, dynamically coupled with a Geographic Information System (GIS). I refer to these tools in Chapters 14 and 15.
Apart from updating and revising each chapter and adding colour photographs, I have provided sets of illustrative problems and questions at the end of each chapter, based on my experience in teaching undergraduate classes on soil resources and their management at The University of Melbourne. I have benefited from feedback from students and also from advice given by friends and colleagues, notably Dr Nick Uren and Dr Robert Edis. To all those who contributed I am most grateful, but the ultimate responsibility for any errors and omissions rests with me. I am also grateful to Debbie Seymour, Rosie Hayden and Hannah Berry at Blackwell Publishing who have been very tolerant and supportive while I was preparing this edition.
Robert E. WhiteMelbourne13 December 2004
Units of Measurement and Abbreviations used in this Book
SI units
Basic unit
Abbreviation
metre
m
hectare
ha
gram
g
mole
mol
second
s
temperature
K
ampere
A
becquerel
Bq
Derived units
Unit
Abbreviation
Value
Celsius
°C
K-273
newton
N
kg.m/s
joule
J
N.m
pascal
P
N/m
2
Volt
V
J/A/s
siemen
S
A/V
Coulomb
C
A.s
litre
L
m
3
/1000
tonne
t
kg.1000
bar
bar
Pa.10
5
Faraday’s constant
F
96500 J/mol/V
Universal gas constant
R
8.3143 J/K/mol
Non-SI units used in soil science
Prefixes and suffixes to units
Prefix/suffix
Abbreviation
Value
tera-
T
10
12
giga-
G
10
9
mega-
M
10
6
kilo-
k
10
3
deca-
da
10
1
deci-
d
10
−1
centi-
c
10
−2
milli-
m
10
−3
micro-
μ
10
−6
nano-
n
10
−9
pico-
p
10
−12
Miscellaneous symbols
( )
denotes ‘activity’
[ ]
denotes ‘concentration’
≅
approximately equal to
∼
of the order of
<
less than
<
greater than
≤
less than or equal to
≥
greater than or equal to
log
log
10
ln
log
e
exp
exponential of
Abbreviations
ACIAR
Australian Centre for International Agricultural Research
ACLEP
Australian Collaborative Land Evaluation Program
ADAS
Agricultural Development and Advisory Service
AEC
anion exchange capacity
AM
arbuscular mycorrhizas
AMO
ammonia mono-oxygenase
ANZECC
Australian and New Zealand Environment and Conservation Council
AR
activity ratio
ASC
Australian Soil Classification
ASRIS
Australian Soil Resources Information System
ASSSI
Australian Society of Soil Science Inc.
ATC
4-amino-1,2,4-triazole
ATP
adenosine triphosphate
AWC
available water capacity
BET
Brunauer, Emmet and Teller
BIO
microbial biomass
BMP
best management practice
BP
before present
CEC
cation exchange capacity
CFCs
chlorofluorocarbons
CPMAS
cross-polarization, magic angle spinning
CREAMS
Chemicals, Runoff and Erosion from Agricultural Management Systems
CRF
controlled-release fertilizer
CSIRO
Commonwealth Scientific and Industrial Research Organization
DAP
diammonium phosphate
DCD
dicyandiamide
DCP
dicalcium phosphate
DCPD
dicalcium phosphate dihydrate
DDL
diffuse double layer
DDT
dichlorodiphenyltrichloroethane
DEM
digital elevation model
DL
diffuse layer
DM
dry matter
DNA
desoxyribose nucleic acid
DOC
dissolved organic carbon
DPM
decomposable plant material
DTPA
diethylene triamine pentaacetic acid
E
evaporation
EC
electrical conductivity
ECEC
effective cation exchange capacity
EDDHA
ethylenediamine di (O-hydroxyphenylacetic acid)
EDTA
ethylenediamine tetraacetic acid
EMR
electromagnetic radiation
ENV
effective neutralizing value
EOC
extracted organic C
ESD
ecologically sustainable development
ESP
exchangeable sodium percentage
E
t
evapotranspiration
EU
European Union
FA
fulvic acid
FAO
Food and Agriculture Organization
FC
field capacity
FESLM
Framework for Sustainable Land Management
FTIR
Fourier Transform Infrared
FYM
farmyard manure
GIS
Geographic Information System
GLC
gas-liquid chromatography
GPS
Global Positioning System
GR
gypsum required
HA
humic acid
HAp
hydroxyapatite
HARM
hull acid rain model
HUM
humified organic matter
HYV
high-yielding variety
IBDU
isobutylidene urea
IOM
inert organic matter
IPCC
Intergovernmental Panel on Climate Change
IPM
integrated pest management
IR
infiltration rate
IS
inner sphere
IUSS
International Union of Soil Sciences
KE
kinetic energy
LAI
leaf area index
LF
leaching fraction
LR
leaching requirement
LRA
land resource assessment
LSI
Langelier saturation index
MAFF
Ministry of Agriculture, Fisheries and Food
MAH
monocyclic aromatic hydrocarbons
MAP
monoammonium phosphate
MCP
monocalcium phosphate
MDB
Murray-Darling Basin meq milli-equivalent
MPN
most probable number
MWD
maximum potential soil water deficit
NASIS
National Soil Information System
NCPISA
National Collaborative Project on Indicators for Sustainable Agriculture
NDS
non-linear dynamic systems
NHMRC
National Health and Medical Research Council
NMR
nuclear magnetic resonance
NRCS
Natural Resources Conservation Service
NSESD
National Strategy for Ecologically Sustainable Development
NV
neutralizing value
OCP
octacalcium phosphate
o.d.
oven-dry
OS
outer sphere
p, pp.
page, singular and plural
P
precipitation
PAH
polycyclic aromatic hydrocarbons
PAM
polyacrylamide
PAPR
partially acidulated phosphate rock
PAW
plant available water
PBC
phosphate buffering capacity
PEG
polyethyleneglycol
POM
particulate organic matter
PR
phosphate rock
PSCU
polymer-coated sulphur-coated urea
PVA
polyvinyl alcohol
PVAc
polyvinylacetate
PVC
polyvinyl chloride
PWP
permanent wilting point
PZC
point of zero charge
Q/I
quantity/intensity
RAW
readily available water
RH
relative humidity
RNA
ribose nucleic acid
RPM
resistant plant material
RPR
reactive phosphate rock
RUSLE
Revised Universal Soil Loss Equation
RWEQ
Revised Wind Erosion Equation
SAR
sodium adsorption ratio
SCU
sulphur-coated urea
SGS
Sustainable Grazing Systems
SI
Système International
SIR
substrate-induced respiration
SLM
sustainable land management
SOM
soil organic matter
SOTER
World Soils and Terrain Database
sp, spp.
species, singular and plural
SRF
slow-release fertilizer
SSP
single superphosphate
ST
Soil Taxonomy
SUNDIAL
Simulation of Nitrogen Dynamics in Arable Land
SWD
soil water deficit
TCP
tricalcium phosphate
TDR
time domain reflectometer/reflectometry
TDS
total dissolved salts
TEC
threshold electrolyte concentration
TSP
triple superphosphate
UF
urea formaldehyde
UN
United Nations
UNEP
United Nations Environment Program
USDA
United States Department of Agriculture
USLE
Universal Soil Loss Equation
VD
vapour density
VP
vapour pressure
WCED
World Commission on Environment and Development
WEPP
Water Erosion Prediction Project
WEPS
Wind Erosion Prediction System
WEQ
Wind Erosion Equation
WHO
World Health Organization
WRB
World Reference Base for Soil Resources
VFA
volatile fatty acid
XRD
X-ray diffraction
Part 1
The Soil Habitat
‘Soils are the surface mineral and organic formations, always more or less coloured by humus, which constantly manifest themselves as a result of the combined activity of the following agencies; living and dead organisms (plants and animals) parent material, climate and relief.’
V. V. Dokuchaev (1879), quoted by J. S. Joffe in Pedology
‘The soil is teeming with life. It is a world of darkness, of caverns, tunnels and crevices, inhabited by a bizarre assortment of living creatures …’
J. A. Wallwork (1975) in The Distribution and Diversity of Soil Fauna
Redrawn from Reganold J. P., Papendick R. I. & Parr J. F. (1990) Sustainable agriculture. Scientific American 262(6), 112–20.
Chapter 1
Introduction to the Soil
1.1 Soil in the making
With the exposure of rock to a new environment – following an outflow of lava, an uplift of sediments, recession of a water body, or the retreat of a glacier – a soil begins to form. Decomposition proceeds inexorably towards decreased free energy and increased entropy. The free energy of a closed system, such as a rock fragment, is that portion of its total energy that is available for work, other than work done in expanding its volume. Part of the energy released in a spontaneous reaction, such as rock weathering, appears as entropic energy, and the degree of disorder created in the system is measured by its entropy. For example, as the rock weathers, minerals of all kinds are converted into simpler molecules and ions, some of which are leached out by water or escape as gases.
Weathering is hastened by the appearance of primitive plants on rock surfaces. These plants – lichens, mosses and liverworts – can store radiant energy from the sun as chemical energy in the products of photosynthesis. Lichens, which are symbiotic associations of an alga and fungus, are able to ‘fix’ atmospheric nitrogen (N2) and incorporate it into plant protein, and to extract elements from the weathering rock surface. On the death of each generation of these primitive plants, some of the rock elements and a variety of complex organic molecules are returned to the weathering surface where they nourish the succession of organisms gradually colonizing the embryonic soil.
A simple example is that of soil formation under the extensive deciduous forests of the cool humid areas of Europe, Asia and North America, on calcareous deposits exposed by the retreat of the Pleistocene ice cap (Table 1.1). The profile development is summarized in Fig. 1.1. The initial state is little more than a thin layer of weathered material stabilized by primitive plants. Within a century or so, as the organo-mineral material accumulates, more advanced species of sedge and grass appear, which are adapted to the harsh habitat. The developing soil is described as a Lithosol (Entisol or Rudosol*). Pioneering micro-organisms and animals feed on the dead plant remains and gradually increase in abundance and variety. The litter deposited on the surface is mixed into the soil by burrowing animals and insects, where its decomposition is hastened. The eventual appearance of larger plants – shrubs and trees – with their deeper roots, pushes the zone of rock weathering farther below the soil surface. After a few hundred more years, a Brown Forest Soil (Inceptisol or Tenosol) emerges. We shall return to the topic of soil formation, and the wide range of soils that occur in the landscape, in Chapters 5 and 9.
1.2 Concepts of soil
The soil is at the interface between the atmosphere and lithosphere (the mantle of rocks making up the Earth’s crust). It also has an interface with bodies of fresh and salt water (collectively called the hydrosphere). The soil sustains the growth of many plants and animals, and so forms part of the biosphere.
* See Box 1.1 for a discussion of soil names.
Fig. 1.1 Stages in soil formation on a calcareous parent material in a humid temperate climate.
There is little merit in attempting to give a rigorous definition of soil because of the complexity of its make-up, and of the physical, chemical and biological forces that act on it. Nor is it necessary to do so, for soil means different things to different users. For example, to the geologist and engineer, the soil is little more than finely divided rock material. The hydrologist may see the soil as a storage reservoir affecting the water balance of a catchment, while the ecologist may be interested only in those soil properties that influence the growth and distribution of plants and animals. The farmer is naturally concerned about the many ways in which soil influences crop growth and the health of his livestock, although frequently his interest does not extend below the depth of soil disturbed by a plough (15–20 cm).
Fig. 1.2 Sites of primitive settlements in the Middle East (after Gates, 1976).
In view of this wide spectrum of potential user-interest, it is appropriate when introducing the topic of soil to readers, perhaps for the first time, to review briefly the evolution of our relationship with the soil and identify some of the past and present concepts of soil.
Soil as a medium for plant growth
Human’s use of soil for food production began two or three thousand years after the close of the last Pleistocene ice age, which occurred about 11,000 years BP (before present). Neolithic people and their primitive agriculture spread outwards from settlements in the fertile crescent embracing the ancient lands of Mesopotamia, Canaan and southern Turkey (Fig. 1.2) and reached as far as China and the Americas within a few thousand years. In China, for example, the earliest records of soil survey (4000 years BP) show how soil fertility was used as a basis for levying taxes on landholders. To study the soil was a practical exercise of everyday life, and the knowledge of soil husbandry that had been acquired by Roman times was passed on by peasants and landlords, with little innovation, until the early 18th century.
From that time onwards, however, the rise in demand for agricultural products in Europe was dramatic. Conditions of comparative peace, and rising living standards as a result of the Indus Revolution, further stimulated this demand throughout the 19th century. The period was also one of great discoveries in physics and chemistry, the implications of which sometimes burst with shattering effect on the conservative world of agriculture. In 1840, von Liebig established that plants absorbed nutrients as inorganic compounds from the soil, although he insisted that plants obtained their nitrogen (N) from the atmosphere: Lawes and Gilbert at Rothamsted subsequently demonstrated that plants (except legumes) absorbed inorganic N from the soil. In the 1850s, Way discovered the process of cation exchange in soil. During the years from 1860 to 1890, emin bacteriologists including Pasteur, Warington and Winogradsky elucidated the role of microorganisms in the decomposition of plant residues and the conversion of ammonia to nitrate.
Fig. 1.3 Soil development in relation to the landscape and underlying regolith.
Over the same period, botanists such as von Sachs and Knop, by careful experiments in water culture and analysis of plant ash, identified the major elements that were essential for healthy plant growth. Agricultural chemists drew up balance sheets of the quantities of these elements taken up by crops and, by inference, the quantities that should be returned to the soil in fertilizers or animal manure to sustain growth. This approach, whereby the soil was regarded as a relatively inert medium providing water, mineral* ions and physical support for plants, has been called the ‘nutrient bin’ concept.
* The term ‘mineral’ is used in two contexts: first, as an adjective referring to the inorganic constituents of the soil (ions, salts and particulate matter); second, as a noun referring to specific inorganic compounds found in rocks and soil, such as quartz and feldspars (Chapter 2).
Soil and the influence of geology
The pioneering chemists who investigated a soil’s ability to supply nutrients to plants tended to see the soil as a chemical and biochemical reaction medium. They little appreciated soil as part of the landscape, moulded by natural forces acting on the land surface. In the late 19th century, great contributions were made to our knowledge of soil by geologists who defined the mantle of loose, weathered material on the Earth’s surface as the regolith, of which only the upper 50–150 cm, superficially enriched with organic matter, could be called soil (Fig. 1.3). Below the soil was the subsoil that was largely devoid of organic matter. However, the mineral matter of both soil and subsoil was recognized as being derived from the weathering of underlying rocks, which led to an interest in the influence of rock type on the soils formed. As the science of geology developed, the history of the Earth’s rocks was subdivided into a time scale consisting of eras, periods and epochs, going back some 550 million years BP. Periods within the eras are usually associated with prominent sequences of sedimentary rocks that were deposited in the region now known as Europe. But examples of these rocks are found elsewhere, so the European time divisions have gradually been accepted worldwide (although the European divisions are not necessarily as clear-cut in all cases outside Europe). Studies of the relationship between soil and the underlying geology led to the practice of classifying soils loosely in geological terms, such as granitic (from granite), marly (derived from a mixture of limestone and clay), loessial (derived from wind-blown silt-size particles), glacial (from glacial deposits) and alluvial (from river deposits).
Table 1.1 The geological time scale.
A simplified version of the geological time scale from the pre-Cambrian period to the present is shown in Table 1.1.
The influence of Russian soil science
Russia is a vast country covering many climatic zones in which, at the end of the 19th century, crop production was limited not so much by soil fertility, but by primitive methods of agriculture. Early Russian soil science was therefore concerned not with soil fertility, but with observing soils in the field and studying relationships between soil properties and the environment in which the soil had formed. From 1870 onwards, Dokuchaev and his school emphasized the distinctive features of a soil that developed gradually and distinguished it from the undifferentiated weathering rock or parent material below. This was the beginning of the science of pedology*.
* From the Greek word for ground or earth.
Following the Russian lead, scientists in other countries began to appreciate that factors such as climate, parent material, vegetation, topography and time interacted in many ways to produce an almost infinite variety of soil types. For any particular combination of these soil-forming factors (Chapter 5), a unique physicochemical and biological environment was established that led to the development of a distinctive soil body – the process of pedogenesis. A set of new terms was developed to describe soil features, such as:
Soil profile – constituting a vertical face exposed by excavating the soil from the surface to the parent material;
soil horizons – layers in the profile distinguished by their colour, hardness, texture, the occurrence of included structures, and other visible or tangible properties. The upper layer, from which materials are generally washed downwards, is described as eluvial; lower layers in which these materials accumulate are called illuvial.
In 1932, an international meeting of soil scientists adopted the notation of A and B for the eluvial and illuvial horizons, respectively, and C horizon for the parent material. The A and B horizons comprise the solum. Unweathered rock below the parent material is called bedrock R. Organic litter on the surface, not incorporated in the soil, is designated as an L layer. A typical Alfisol (ST) or Chromosol (ASC) soil profile showing a well-developed A, B and C horizon is shown in Fig. 1.4.
Soil genesis is now known to be much more complex than this early work suggested. For example, many soils are polygenetic in origin; that is, they have undergone successive phases of development due to changes in climate and other environmental factors over time. In other cases, two or more layers of different parent material are found in one soil profile. Nevertheless, the Russian approach was a considerable advance on traditional thinking, and recognition of the relationship between a soil and its environment encouraged soil scientists to survey and map the distribution of soils. The wide range of soil morphology that was revealed in turn stimulated studies of pedogenesis, an understanding of which, it was believed, would enable the copious field data on soils to be collated more systematically. Thus, Russian soil science provided the inspiration for many of the early soil classifications.
Fig. 1.4 Profile of an Alfisol (ST) or Chromosol (ASC) showing well-developed A, B and C horizons (see Plate 1.4).
A contemporary view of soil
Between the two World Wars of the 20th century, the philosophy of the soil as a ‘nutrient bin’ was prevalent, particularly in the western world. More and more land was brought into cultivation, much of which was marginal for crop production because of limitations of climate, soil and topography. With the balance between crop success and failure made even more precarious than in favourable areas, the age-old problems of wind and water erosion, encroachment by weeds, and the accumulation of salts in irrigated lands became more serious. Since 1945, demand for food, fibre and forest products from an escalating world population (now > 6 billion) has led to increased use of fertilizers to improve yields, and pesticides to control pests and diseases (Chapter 12). Such practices have resulted in some accumulation of undesirable pesticide residues in soil, and in increased losses of soluble constituents such as nitrate and phosphates to surface waters and groundwater. There has also been widespread dispersal of the very stable pesticides (e.g. organo-chlorines) in the biosphere, and their accumulation to concentrations potentially toxic to some species of birds and fish.
More recently, however, scientists, producers and planners have acknowledged the need to compromise between maximizing crop production and conserving a valuable natural resource. Emphasis is now placed on maintaining the soil’s natural condition by minimizing the disturbance when crops are grown, matching fertilizer additions more closely to crop demand in order to reduce losses, using legumes to fix N2 from the air, and returning plant residues and waste materials to the soil to supply some of a crop’s nutrient requirements. In short, more emphasis is being placed on the soil as a natural body (Box 1.2) and on the concept of sustainable land management (Chapter 15).
1.3 Components of the soil
We have seen that a combination of physical, chemical and biotic forces acts on organic materials and weathered rock to produce a soil with a porous fabric that retains water and gases. The mineral matter derived from weathered rock consists of particles of different size, ranging from clay (the smallest), to silt, sand, gravel, stones, and in some cases boulders (Section 2.1). The particle density ρp (rho p) varies according to the mineralogy (Section 2.3), but the average ρp is 2.65 Mg/m3. Organic matter has a lower density of 1–1.3 Mg/m3, depending on the extent of its decomposition. Water has a density of 1.0 Mg/m3 at normal temperatures (c. 20°C)*.
Fig. 1.5 Proportions of the main soil components by volume.
Soil water contains dissolved organic and inorganic solutes and is called the soil solution. While the soil air consists primarily of N2 and oxygen (O2), it usually contains higher concentrations of carbon dioxide (CO2) than the atmosphere, and traces of other gases that are by-products of microbial metabolism. The relative proportions of the four major components – mineral matter, organic matter, water and air – may vary widely, but generally lie within the ranges indicated in Fig. 1.5. These components are discussed in more detail in the subsequent chapters of Part 1.
1.4 Summary
Soil forms at the interface between the atmosphere and the weathering products of the regolith. Physical and chemical weathering, erosion and redeposition, combined with the activities of a succession of colonizing plants and animals, moulds a distinctive soil body from the milieu of rock minerals in the parent material. The process of soil formation, called pedogenesis, culminates in a remarkably variable differentiation of soil material into a series of horizons that constitute a soil profile. Soil horizons are distinguished by their visible and tangible properties such as colour, hardness, texture and structural organization. The intimate mixing of mineral and organic matter to form a porous fabric, permeated by water and air, creates a favourable habitat for a variety of plant and animal life. Soil is a fragile component of the environment. Its use for food and fibre production, and waste disposal, must be managed in a way that minimizes the off-site effects of these activities and preserves the soil for future generations. This is the basis of sustainable soil management.
* The density of water is 1.000 Mg/m3 at 4°C and 0.998 Mg/m3 at 20°C, which is rounded to 1.0.
References
Avery B. W. (1980) Soil Classification for England and Wales. Soil Survey Technical Monograph No. 14. Rothamsted Experimental Station, Harpenden.
Baldwin H., Kellogg C. W. & Thorpe J. (1938) Soil classification, in Soils and Man. United States Government Printing Office. Washington DC.
FAO (1998) World Reference Base for Soil Resources. World Resources Report No. 84. FAO, Rome.
Gates C. T. (1976) China in a world setting: agricultural response to climatic change. Journal of the Australian Institute of Agricultural Science 42, 75–93.
Isbell R. F. (2002) The Australian Soil Classification, revised edn. Australian Soil and Land Survey Handbooks Series Volume 4. CSIRO Publishing, Melbourne.
Soil Survey Staff (1999) Soil Taxonomy. A Basic Classification for Making and Interpreting Soil Surveys, 2nd edn. United States Department of Agriculture Handbook No. 436. Natural Resources Conservation Service, Washington DC.
Further reading
Hillel D. (1991) Out of the Earth: Civilization and the Life of the Soil. The Free Press, New York.
Jenny H. (1980) The Soil Resource – Origin and Behaviour. Springer-Verlag, New York.
Marschner H. (1995) Mineral Nutrition of Higher Plants, 2nd edn. Academic Press, London.
McKenzie N. & Brown K. (2004) Australian Soils and Landscapes: an Illustrated Compendium. CSIRO Publishing, Melbourne.
Chapter 2
The Mineral Component of the Soil
2.1 The size range
Rock fragments and mineral particles in soil vary enormously in size from boulders and stones down to sand grains and very small particles that are beyond the resolving power of an optical microscope (< 0.2 μm in diameter). Particles smaller than c. 1 μm are classed as colloidal. Particles that do not settle quickly when mixed with water are said to form a colloidal solution or sol; if they settle within a few hours they form a suspension. Colloidal solutions are distinguished from true solutions (dispersions of ions and molecules) by the Tyndall effect. This occurs when the path of a beam of light passing through the solution can be seen from either side at right angles to the beam, indicating a scattering of the light rays.
An arbitrary division is made by size-grading soil into material:
That passes through a sieve with 2 mm diameter holes – the
fine earth
, and
that retained on the sieve (> 2 mm ) –
the stones or gravel
, but smaller than
fragments > 600 mm, which are called
boulders
. The separation by sieving is carried out on air-dry soil that has been gently ground by mortar and pestle, or crushed between wooden rollers, to break up the aggregates. Air-dry soil is soil allowed to dry in air at ambient temperatures (between 20 and 40°C).
Particle-size distribution of the fine earth
The distribution of particle sizes determines the soil texture, which may be assessed subjectively in the field or more rigorously by particle-size analysis in the laboratory.
Size classes
All soils show a continuous range of particle sizes, called a frequency distribution, which is obtained by plotting the number (or mass) of particles of a given size against their actual size. When the number or mass in each size class is summed sequentially we obtain a cumulative distribution of soil particle sizes, some examples of which are given in Fig. 2.1. In practice, it is convenient to subdivide the continuous distribution into several class intervals that define the size limits of the sand, silt and clay fractions. The extent of this subdivision, and the class limits chosen, vary from country to country and even between institutions within countries. The major systems in use are those adopted by the Soil Survey Staff of the USDA, the British Standards Institution and the International Union of Soil Sciences (IUSS). These are illustrated in Fig. 2.2. All three systems set the upper limit for clay at 2 μm diameter, but differ in the upper limit chosen for silt and the way in which the sand fraction is subdivided.
Fig. 2.1 Cumulative frequency distributions of soil particle sizes in a typical clay, sandy silt loam and sandy soil.
Field texture
A soil surveyor assesses soil texture by moistening a sample with water until it glistens. It is then kneaded between fingers and thumb until the aggregates are broken down and the soil grains thoroughly wetted. The proportions of sand, silt and clay are estimated according to the following qualitative criteria:
Coarse sand
grains are large enough to grate against each other and can be detected individually by sight and feel;
fine sand
grains are much less obvious, but when they comprise more than about 10% of the sample they can be detected by biting the sample between the teeth;
silt
grains cannot be detected by feel, but their presence makes the soil feel smooth and silky and only slightly sticky;
clay
is characteristically sticky, although some dry clays, especially of the expanding type (Section 2.3), require much moistening and kneading before they develop their maximum stickiness.
High organic matter contents tend to reduce the stickiness of clay soils and to make sandy soils feel more silty. Finely divided calcium carbonate also gives a silt-like feeling to the soil.
Depending on the estimated proportions of sand, silt and clay, the soil is assigned to a textural class according to a triangular diagram (Fig. 2.3). The triangle in Fig. 2.3a is used by the Soil Survey of England and Wales and is based on the British Standards system of particle-size grading (Fig. 2.2); the one in Fig. 2.3b is used in Australia and is based on the International system (Fig. 2.2). The USDA system is very similar to the British Standards system. Note that in these systems there are 11 textural classes, but the Australian system has broader classes for the silty clays, silty clay loams and silty loams than the British or USDA systems.
Fig. 2.2 Particle-size classes most widely adopted internationally.
Fig. 2.3 (a) Triangular diagram of soil textural classes adopted in England and Wales (after Hodgson, 1974). (b) Triangular diagram of soil textural classes adopted in Australia (after McDonald et al., 1998).
Soil surveyors become expert at texturing after years of experience, which is gained by their checking field assessments of texture against a laboratory analysis of a soil’s particle-size distribution.
Particle-size analysis in the laboratory
The success of the method relies on the complete disruption of soil aggregates and the addition of chemicals that ensure dispersion of the soil colloids in water. Full details of the methods employed are given in standard texts, for example Klute (1986) and Rayment and Higginson (1992). The sand particles are separated by sieving; silt and clay are separated using the differences in their settling velocities in suspension. The principle of the latter technique is outlined in Box 2.1.
The result of particle-size analysis is expressed as the mass of the individual fractions per 100 g of oven-dry (o.d.) soil (fine earth only). Oven-dry soil is soil dried to a constant weight at 105°C. When the coarse and fine sand fractions are combined, the soil may be represented by one point on the triangular diagrams of Fig. 2.3 (a, b). Alternatively, by stepwise addition of particle-size percentages, graphs of cumulative percentage against particle diameter of the kind shown in Fig. 2.1 are obtained.
2.2 The importance of soil texture
Soil scientists are primarily interested in the texture of the fine earth fraction. Nevertheless, in some soils the size and abundance of stones cannot be ignored because they can have a marked influence on the soil’s suitability for agriculture. As the stone content increases, a soil holds less water than a stoneless soil of the same fine-earth texture, so that crops become more susceptible to drought. Conversely, such soils may be better drained and therefore warm up more quickly in spring in cool temperate regions. Large stones on the soil surface act as sinks during daytime for heat energy that is slowly released at night – this is of benefit in cool climate vineyards, such as in the Rhône Valley, France, where frost in spring and early summer can damage flowering and fruit set (Fig. 2.4). Stoniness also determines the ease, and to some extent the cost of cultivation, as well as the abrasive effect of the soil on tillage implements.
(B2.1.1)
(B2.1.2)
Fig. B2.1.1 A settling soil suspension and plummet balance (courtesy of J. Loveday).
Texture is one of the most stable soil properties and is a useful index of several other properties that determine a soil’s agricultural potential. Fine and medium-textured soils, such as clays, clay loams, silty clays and silty clay loams, are generally more desirable than coarse-textured soils because of their superior retention of nutrients and water. Conversely, where rapid infiltration and good drainage are required, as for irrigation or liquid waste disposal, sandy or coarse-textured soils are preferred. In farming terms, clay soils are described as ‘heavy’ and sandy soils as ‘light’, which does not refer to their mass per unit volume, but to the power required to draw a plough or other implements through the soil. Because it is easy to estimate, and is routinely measured in soil surveys, texture (and more specifically clay content) has been used as a ‘surrogate’ variable for other soil properties that are less easily measured, such as the cation exchange capacity (Section 2.5).
Texture has a pronounced effect on soil temperature. Clays hold more water than sandy soils, and the presence of water considerably modifies the heat required to change a soil’s temperature because:
Its specific heat capacity is 3–4 times that of the soil solids;
considerable latent heat is either absorbed or evolved during a change in the physical state of water, for example, from ice to liquid or
vice versa
. Thus, the temperature of wet clay soils responds more slowly than that of sandy soils to changes in air temperature in spring and autumn (Section 6.6).
Fig. 2.4 Boulders and stones covering the soil surface in a vineyard in the central Rhône Valley, France (see also Plate 2.4).
Texture should not be confused with tilth, of which it is said that a good farmer can recognize it with his boot, but no soil scientist can describe it. Tilth refers to the condition of the surface of ploughed soil prepared for seed sowing: how sticky it is when wet and how hard it sets when dry. The action of frost in cold climates breaks down the massive clods left on the surface of a heavy clay soil after autumn ploughing, producing a mellow ‘frost tilth’ of numerous small granules (Section 4.2).
2.3 Mineralogy of the sand and silt fractions
Simple crystalline structures
Sand and silt consist almost entirely of the resistant residues of primary rock minerals, although small amounts of secondary minerals (salts, oxides and hydroxides) formed by weathering also occur. The primary rock minerals are predominantly silicates, which have a crystalline structure based upon a simple unit – the silicon tetrahedron, (Fig. 2.5). An electrically neutral crystal is formed when cations, such as Al3+, Fe3+, Fe2+, Ca2+, Mg2+, K+ and Na+, become covalently bonded to the O atoms in the tetrahedron and the surplus valencies of the O2‒ ions in the group are satisfied. An example of this kind of structure is the primary mineral olivine, which has the composition (Mg, Fe)2SiO4.
Fig. 2.5 Diagram of a Si tetrahedron (interatomic distances not to scale).
Coordination number
The packing of the O atoms, which are the largest of the more abundant elements in the silicates, determines the crystalline dimensions. In quartz, for example, O occupies 98.7% and Si only 1.3% of the mineral volume. The size ratio of Si to O is such that four O atoms can be packed around one Si, and larger cations such as iron (Fe) can accommodate more O atoms. The cation to oxygen radius ratio determines the coordination number of the cation (Table 2.1). Many of the common metal cations have radius ratios between 0.41 and 0.73, which means that an octahedral arrangement of six O atoms around the cation (coordination number 6) is possible. Larger alkali and alkaline earth cations, such as K+ and Ba2+, that have radius ratios > 0.73 form complexes of coordination number 8 or greater. Aluminium, which has a cation to oxygen radius ratio close to the maximum for coordination number 4 and the minimum for coordination number 6 (0.41), can exist in either fourfold (IV) or sixfold (VI) coordination.
Table 2.1 Cation to oxygen radius ratios and coordination numbers for common elements in the silicate minerals. After Schulze, 1989.
Isomorphous substitution
Elements of the same valency and coordination number frequently substitute for one another in a silicate structure – a process called isomorphous substitution. The structure remains electrically neutral. However, when elements of the same coordination number but different valency are exchanged, there is an imbalance of charge. The most common substitutions are Mg2+, Fe2+ or Fe3+ for Al3+ in octahedral coordination, and Al3+ for Si4+ in tetrahedral coordination. The excess negative charge is neutralized by the incorporation of additional cations, such as unhydrated K+, Na+, Mg2+ or Ca2+ into the crystal structure, or by structural arrangements that allow an internal compensation of charge (see the chlorites).
More complex crystalline structures
Chain structures
These are represented by the pyroxene and amphibole groups of minerals, which collectively make up the ferromagnesian minerals. In the pyroxenes, each Si tetrahedron is linked to adjacent tetrahedra by the sharing of two out of three basal O atoms to form a single extended chain (Fig. 2.6). In the amphiboles, two parallel pyroxene chains are linked by the sharing of an O atom in every alternate tetrahedron (Fig. 2.7). Cations such as Mg2+, Ca2+, Al3+ and Fe2+ are ionic-covalently bonded to the O2‒ ions to neutralize the surplus negative charge.
Fig. 2.6 Single chain structure of Si tetrahedra as in a pyroxene. The arrows indicate O atoms shared between adjacent tetrahedra (after Bennett and Hulbert, 1986).
Fig. 2.7 Double chain structure of Si tetrahedra as in an amphibole. The arrows indicate O atoms shared between chains (after Bennett and Hulbert, 1986).
Sheet structures
It is easy to visualize an essentially one-dimensional chain structure being extended in two dimensions to form a sheet of Si tetrahedra linked by the sharing of all the basal O atoms. When viewed from above the sheet, the bases of the linked tetrahedra form a network of hexagonal holes (Fig. 2.8). The apical O atoms (superimposed on the Si atoms in Fig. 2.8) form ionic-covalent bonds with other metal cations by, for example, displacing OH groups from their coordination positions around a trivalent Al3+ ion. (Because the unhydrated proton is so small, OH occupies virtually the same space as O.) When Al octahedral units (one is shown diagrammatically in Fig. 2.9) are linked by the sharing of edge OH groups, they form an alumina sheet. Aluminium atoms normally occupy only two-thirds of the available octahedral positions – this is a dioctahedral structure, characteristic of the mineral gibbsite, [Al2OH6]n. If Mg is present instead of Al, however, all the available octahedral positions are filled – this is a trioctahedral structure, characteristic of the mineral brucite [Mg3(OH)6]n.
The bonding together of silica and alumina sheets through the apical O atoms of the Si tetra-hedra causes the bases of the tetrahedra to twist slightly so that the cavities in the sheet become trigonal rather than hexagonal in shape. When two silica sheets sandwich one alumina sheet, the result is a 2 : 1 layer structure characteristic of the micas, chlorites and many soil clay minerals (Section 2.4). Two-dimensional layer crystal structures such as these are typical of the phyllosilicates, some characteristics of which are given in Box 2.2.
Additional structural complexity is introduced by isomorphous substitution. In the dioctahedral mica muscovite, for example, one-quarter of the tetrahedral Si4+ is replaced by Al3+ resulting in a net 2 moles of negative charge per unit cell (Box 2.3). Muscovite has the structural composition:
Fig. 2.8 A silica sheet in plan view showing the pattern of hexagonal holes (after Fitzpatrick, 1971).
Fig. 2.9 Diagram of an Al octahedron (interatomic distances not to scale).
Fig. B2.2.1 General structure of a phyllosilicate crystal.
.
Note that the negative charge is neutralized by K+ ions held in the spaces formed by the juxtaposition of the trigonal cavities of adjacent silica sheets (Fig. 2.10). Because the isomorphous substitution in muscovite occurs in the tetrahedral sheet, the negative charge is distributed over only three surface O atoms. The magnitude of the layer charge and its localization are sufficient to cause cations of relatively small ionic potential*, such as K+, to lose their water of hydration. The unhydrated K+ ions have a diameter comparable to that of the ditrigonal cavity formed between opposing siloxane surfaces and their presence provides very strong bonding between the layers. Such complexes, where an unhydrated ion forms an ionic-covalent bond with atoms of the crystal surface, are called inner-sphere (IS) complexes.
Fig. 2.10 Structure of muscovite mica – a 2 : 1 non-expanding phyllosilicate mineral.
Biotite is a trioctahedral mica which has Al3+ substituted for Si4+ in the tetrahedral sheet and contains Fe2+ and Mg2+ in the octahedral sheet. Again, because of the localization of charge in the tetrahedral sheet, unhydrated K+ ions are retained in the interlayer spaces to give a unit cell formula of:
.
* Determined by the charge to radius ratio. As this ratio increases, the ionic potential of the cation increases.
The most complex of the two-dimensional structures belongs to the chlorites, which have a brucite layer sandwiched between two mica layers. In the type mineral chlorite, the negative charge of the two biotite layers is neutralized by a positive charge in the brucite, developed due to the replacement of two-thirds of the Mg2+ cations by Al3+ (Fig. 2.11). Chlorite is an example of a regular mixed-layer mineral. Predictably, the bonding between layers is strong, but the high content of Mg renders this mineral susceptible to weathering in acidic solutions. For the same reason, and also because of its ferrous (Fe2+) iron content, biotite is much less stable than muscovite. Micas and chlorites weather to form vermiculites and smectites in soil (Section 2.4).
Three-dimensional structures
The most important silicates in this group are silica and the feldspars. Silica minerals consist entirely of polymerized Si tetrahedra of general composition (SiO2)n. Silica occurs as the residual mineral quartz, which is very inert, and as a secondary mineral precipitated after the hydrolysis of more complex silicates. Secondary silica initially exists as amorphous opal that dehydrates over time to form microcrystalline quartz, known as flint or chert. Silica also occurs as amorphous or microcrystalline silica of biological origin. For example, the diatom, a minute aquatic organism, has a skeleton of almost pure silica. Silica absorbed from the soil by terrestrial plants (grasses and hardwood trees in particular) forms opaline structures called phytoliths, which are returned to the soil when the plant dies.
Fig. 2.11 Structure of chlorite – a 2 : 2 or mixed layer mineral.
Quartz or flint fragments of greater than colloidal size are very insoluble, and hence are abundant in the sand and silt fractions of many soils. The feldspars, on the other hand, are chemically more reactive and rarely comprise more than c. 10% of the sand fraction of mature soils. Of these, the potassium feldspars are more resistant to weathering than the Ca and Na feldspars. Their structure consists of a three-dimensional framework of polymerized Si tetrahedra in which some Si4+ is replaced by Al3+. The cations balancing the excess negative charge are all of high coordination number, such as K+, Na+ and Ca2+, and less commonly, Ba2+ and Sr2+. The range of composition encountered is shown in Fig. 2.12
