The Biogeochemistry of Submerged Soils [1 ed.] 9780470863015, 0470863013

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The Biogeochemistry of Submerged Soils Guy Kirk National Soil Resources Institute Cranfield University, UK and formerly International Rice Research Institute, Philippines

Copyright  2004

John Wiley & Sons Ltd, The Atrium, Southern Gate, Chichester, West Sussex PO19 8SQ, England Telephone (+44) 1243 779777

Email (for orders and customer service enquiries): [email protected] Visit our Home Page on www.wileyeurope.com or www.wiley.com 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, scanning or otherwise, except under the terms of the Copyright, Designs and Patents Act 1988 or under the terms of a licence issued by the Copyright Licensing Agency Ltd, 90 Tottenham Court Road, London W1T 4LP, UK, without the permission in writing of the Publisher. Requests to the Publisher should be addressed to the Permissions Department, John Wiley & Sons Ltd, The Atrium, Southern Gate, Chichester, West Sussex PO19 8SQ, England, or emailed to [email protected], or faxed to (+44) 1243 770620. This publication is designed to provide accurate and authoritative information in regard to the subject matter covered. It is sold on the understanding that the Publisher is not engaged in rendering professional services. If professional advice or other expert assistance is required, the services of a competent professional should be sought. Other Wiley Editorial Offices John Wiley & Sons Inc., 111 River Street, Hoboken, NJ 07030, USA Jossey-Bass, 989 Market Street, San Francisco, CA 94103-1741, USA Wiley-VCH Verlag GmbH, Boschstr. 12, D-69469 Weinheim, Germany John Wiley & Sons Australia Ltd, 33 Park Road, Milton, Queensland 4064, Australia John Wiley & Sons (Asia) Pte Ltd, 2 Clementi Loop #02-01, Jin Xing Distripark, Singapore 129809 John Wiley & Sons Canada Ltd, 22 Worcester Road, Etobicoke, Ontario, Canada M9W 1L1 Wiley also publishes its books in a variety of electronic formats. Some content that appears in print may not be available in electronic books. Library of Congress Cataloging-in-Publication Data Kirk, G. J. D. The biogeochemistry of submerged soils / Guy Kirk. p. cm. Includes bibliographical references (p. ). ISBN 0-470-86301-3 (cloth : alk. paper) 1. Hydromorphic soils. 2. Soil chemistry. 3. Biogeochemistry. I. Title. S592.17.H93K57 2004 2003019773 631.4′ 1—dc22 British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library ISBN 0-470-86301-3 Typeset in 10/12pt Times by Laserwords Private Limited, Chennai, India Printed and bound in Great Britain by Antony Rowe Ltd, Chippenham, Wiltshire This book is printed on acid-free paper responsibly manufactured from sustainable forestry in which at least two trees are planted for each one used for paper production.

Contents

Preface

ix

Acknowledgements

xi

1 Introduction 1.1 Global Extent of Submerged Soils and Wetlands 1.2 Biogeochemical Characteristics 1.3 Types of Submerged Soil 1.3.1 Organic Soils 1.3.2 Mineral Soils 1.3.3 Relation between Soils and Landform

1 1 3 9 9 10 12

2 Transport Processes in Submerged Soils 2.1 Mass Flow 2.2 Diffusion 2.2.1 Diffusion Coefficients in Soil 2.2.2 Propagation of pH Changes Through Soil 2.3 Ebullition 2.4 Mixing by Soil Animals

17 19 22 22 35 38 39

3 Interchange of Solutes between Solid, Liquid and Gas Phases A. WATER 3.1 Composition of the Water 3.1.1 Acid and Bases 3.1.2 Speciation 3.1.3 Equilibrium Calculations 3.2 pH Buffer Capacity 3.3 Equilibrium with the Gas Phase 3.3.1 Floodwater CO2 Dynamics 3.4 Gas Transport Across the Air–Water Interface 3.4.1 CO2 Transfer Across the Air–Water Interface B. SOIL 3.5 The Solid Surfaces in Soils 3.6 The Solid Surfaces in Submerged Soils 3.6.1 Organic Matter in Submerged Soils 3.7 Solid–Solution Interactions 3.7.1 Adsorption

45 45 45 46 47 50 53 54 56 58 61 65 65 69 74 76 76

vi

Contents

3.7.2 3.7.3 3.7.4 3.7.5 4

5

Precipitation Co-Precipitation in Solid Solutions Inhibition of Precipitation Equations for Solid—Solution Interactions

Reduction and Oxidation 4.1 Thermodynamics and Kinetics of Redox Reactions 4.1.1 Electron Activities and Free Energy Changes 4.1.2 Redox Potentials 4.1.3 Relation between pe and Concentration of Redox Couples 4.1.4 pe–pH Diagrams 4.1.5 Energetics of Reactions Mediated by Microbes 4.2 Redox Conditions in Soils 4.2.1 Changes with Depth in the Soil 4.2.2 Changes with Time 4.2.3 Calculated Changes in pe, pH and Fe During Soil Reduction 4.2.4 Measurement of Redox Potential in Soil 4.3 Transformations of Nutrient Elements Accompanying Changes in Redox 4.3.1 Transformations of Carbon 4.3.2 Transformations of Nitrogen 4.3.3 Transformations of Sulfur 4.3.4 Transformations of Phosphorus 4.4 Oxidation of Reduced Soil 4.4.1 Kinetics of Fe2+ Oxidation 4.4.2 Simultaneous Diffusion and Oxidation in Soil Biological Processes in the Soil and Floodwater 5.1 Microbiological Processes 5.1.1 Processes Involved in Sequential Reduction 5.1.2 Nitrate Reduction 5.1.3 Iron and Manganese Reduction 5.1.4 Sulfate Reduction 5.1.5 Methanogenesis 5.1.6 Aerobic Processes 5.2 Macrobiological Processes 5.2.1 Net Primary Production and Decomposition 5.2.2 The Floodwater–Soil System 5.2.3 Floodwater Properties 5.2.4 Floodwater Flora 5.2.5 Fauna 5.3 Is Biodiversity Important?

79 82 85 87 93 93 93 97 97 99 102 106 107 109 113 116 119 120 120 122 124 127 128 131 135 135 136 141 142 143 144 147 150 150 151 152 154 159 163

Contents

vii

6 Processes in Roots and the Rhizosphere 6.1 Effects of Anoxia and Anaerobicity on Plant Roots 6.1.1 Adaptations to Anoxia 6.1.2 Armstrong and Beckett’s Model of Root Aeration 6.2 Architecture of Wetland Plant Root Systems 6.2.1 Model of Root Aeration versus Nutrient Absorption 6.2.2 Root Surface Required for Nutrient Absorption 6.3 Nutrient Absorption Properties of Wetland Plant Roots 6.3.1 Ion Transport in Roots 6.3.2 Ion Transport in Wetland Roots 6.4 Root-Induced Changes in the Soil 6.4.1 Oxygenation of the Rhizosphere 6.4.2 The pH Profile Across the Rhizosphere 6.5 Consequences of Root-induced Changes 6.5.1 Nitrification–Denitrification in the Rhizosphere 6.5.2 Solubilization of Phosphate 6.5.3 Solubilization of Zinc 6.5.4 Immobilization of Cations 6.6 Conclusions

165 165 167 170 171 172 177 180 180 184 190 191 194 196 196 197 200 200 202

7 Nutrients, Toxins and Pollutants 7.1 Nutrient and Acidity Balances 7.1.1 Nutrient Balances in Ricefields 7.1.2 Acidity Balances in Ricefields 7.1.3 Peat Bogs 7.1.4 Riparian Wetlands 7.1.5 Tidal Wetlands 7.2 Toxins 7.2.1 Acidity 7.2.2 Iron Toxicity 7.2.3 Organic Acids 7.2.4 Salinity 7.3 Trace Elements 7.3.1 Global Cycling of Trace Elements 7.3.2 Transport Through Soil and into Plant Roots 7.3.1 Mobilities of Individual Trace Elements

203 203 203 208 210 210 211 212 212 214 215 216 218 218 218 220

8 Trace Gases 8.1 Methane 8.1.1 Global Budget 8.1.2 Processes Governing Methane Emissions from Rice 8.1.3 Modelling Methane Emission

233 233 233 234 237

viii

Contents

8.2

8.3

8.4

8.5

8.1.4 Estimating Emissions at the Regional Scale 8.1.5 Possibilities For Decreasing Emissions Nitrogen Oxides 8.2.1 Global Budget 8.2.2 Processes Governing Nitrous and Nitric Oxide Emissions from Rice 8.2.3 Differences between Rice Production Systems Ammonia 8.3.1 Global Budget 8.3.2 Processes Governing Ammonia Emissions from Rice Sulfur Compounds 8.4.1 Global Budget 8.4.2 Emissions from Ricefields Carbon Sequestration

244 246 247 247 249 250 252 252 254 256 256 256 258

References

259

Index

283

Preface

This book is about the movements and transformations of energy and matter in soils that are continuously or intermittently submerged with water. Submerged soils cover a huge area, from 5 to 7 per cent of the Earth’s land surface, and they are undoubtedly of great practical importance: in local, regional and global element cycles, as habitats for plants and wildlife, and in food and fibre production. The submerged soils in ricefields, for example, produce the basic food of more than 2 billion people, a third of the world population. But submerged soils are also inherently interesting scientifically, and that is the main theme of the book. When a soil is submerged, air is excluded and the soil quickly becomes anoxic. A submerged soil is therefore an open, anoxic chemical system, surrounded by oxic systems with very different characteristics. Energy enters through photosynthesis, and inorganic matter enters with percolating water and by gas exchange. Chemical reactions occur through a complicated interchange between solid, liquid and gas phases, largely mediated by biological processes. Further, because plants are the main conduits for gas exchange between the soil and overlying atmosphere, they have a particularly important influence. Submerged soils therefore provide a unique natural laboratory for studying a great range of physical, chemical and biological processes that are important in environmental systems. They form under a wide range of hydrological, geological and topographical conditions, but because of the overriding influence of anoxia, the soils and plants and microbes adapted to them have various characteristics in common. The book describes the physical, chemical and biological processes operating in submerged soils and links them to the dynamics of nutrients, toxins, pollutants and trace gases. Far less research has been done on these topics for submerged soils than for dryland soils, in spite of their importance. But knowledge and understanding of them have increased substantially in the past few decades. Much of the research has been on rice soils, particularly at the International Rice Research Institute (IRRI) which has been involved in research on submerged soils since it was founded in 1960. But there is also much in the ecological and environmental literatures concerned with natural wetlands. In preparing the book I have aimed to deal with generic principles relevant to both natural and artificial wetlands with the aim of serving audiences for both.

Acknowledgements

I thank the following friends and colleagues for their help in planning the book and reviewing draft chapters: Dave Bouldin, Roland Buresh, Ralph Conrad, Achim Dobermann, Dennis Greenland, Peter Nye, Bill Patrick, John Sheehy, Siobhan Staunton, Dick Webster and Oswald van Cleemput. I am indebted to the Director General of IRRI, Ron Cantrell, for the award of a consultancy to write the book and for his encouragement throughout. Most of the writing was done during a sabbatical in the Department of Plant Sciences, University of Cambridge, and I am grateful to the Head of Department, Roger Leigh, and member of the Department for their hospitality. The book was completed during my first months at the National Soil Resources Institute, Cranfield University, and I am indebted to the Director, Mark Kibblewhite, for his encouragement and forbearance. For help with the artwork I am grateful to Edwin Javier, Ely Tabaquero and Gene Hettel, all of IRRI.

1 Introduction

Submerged soils behave and affect the environment in substantially different ways to dryland soils. This chapter discusses the main characteristics and environmental effects of submerged soils and the wetlands they support, and their extent across the globe. 1.1 GLOBAL EXTENT OF SUBMERGED SOILS AND WETLANDS For the purposes of the book I define wetlands as lands that are intermittently or permanently inundated with water to a depth of no more than a few metres. Depending on the precise definition applied, estimates of the total global wetland area range from 700 to 1000 Mha (Aselmann and Crutzen, 1989; Scharpenseel, 1997; Mitsch and Gosselink, 2000). Figure 1.1 shows their approximate distribution and Table 1.1 the extents of different types distinguished by hydrology, vegetation and soil characteristics. The largest areas are the bogs and fens in polar and boreal regions in North America and Russia (34 % of total area); tropical swamps, especially in East Africa and South America (14 % of total area); tropical floodplains, especially of the Amazon and the rivers of South East Asia (10 %); and temperate and tropical ricefields (4 and 12 %, respectively). Almost half the global wetland area is in the tropics. There has been considerable loss of wetlands in many parts of the world over the past 200 years as a result of conversion to agricultural and aquacultural uses. In the US for example, it is estimated that the area has declined from 89 Mha in the 1780s to 49 Mha in the 1980s (Mitsch and Gosselink, 2000). A special class of wetland is the lowland ricefield, which accounts for almost a fifth of the wetland area worldwide. Much of our knowledge and understanding of submerged soils has been gained from research on rice soils. The success of rice as a food crop stems from its origins as a wetland plant and its ability to withstand soil submergence with the attendant improvements in water and nutrient supplies. A corollary is that rice is more sensitive to water deficiency than most other crops, and the critical factors in its productivity are the supply of water to the soil, from rain, river, reservoir or groundwater, and the ability of the soil to retain water. Hence most rice is produced and the highest yields attained on the alluvial deposits associated with major rivers and their deltas. More than 90 % of the production is in Asia, distributed unevenly over four rice The Biogeochemistry of Submerged Soils Guy Kirk  2004 John Wiley & Sons, Ltd ISBN: 0-470-86301-3

2

Introduction

40° 20° 0° Equator 20°

40°

Major Wetland Area Area with Abundant Wetlands

160° 140° 120° 100° 80°

60°

40°

20°

0°

20° 40°

60°

80° 100° 120° 140° 160°

Figure 1.1 Global distribution of wetlands (Mitsch and Gosselink, 2000). Reproduced by permission of Wiley, New York Table 1.1

Global extent of wetlands of different types Area (Mha)

Bogs Fens Swamps Marshes Floodplains Shallow lakes Ricefields Total

Polar

Boreal

Temperate

Tropical

Total

21 54 — — — — — 75

104 62 1 — — — — 167

42 32 10 17 8 1 29 139

20 — 102 10 74 11 80 297

187 148 113 27 82 12 109 678

Definitions of wetland types: Bogs are raised peat-producing wetlands formed in wet climates where organic material has accumulated over long periods. Because they are raised, water and nutrients are entirely derived from the atmosphere, and they are therefore nutrient deficient and acid. Sphagnum moss is the main vegetation, though other types of vegetation are also present in tropical regions. Fens are peat-producing wetlands that receive water and nutrients through inflow from neighbouring land. They are generally less acid than bogs and may be alkaline, and tend to be dominated by grasses and sedges. Because of their better nutrient status they are generally more prolific than bogs. Swamps are forested, freshwater wetlands on submerged soils in which little peat accumulates. This is the US definition; elsewhere the term also includes non-forested wetlands with reeds. Swamps tend to form in warmer climates. Marshes are herbaceous freshwater, non-peat-producing wetlands dominated by grasses, sedges or reeds. The distinction between swamps and marshes may be blurred. Floodplains are periodically inundated areas along rivers or lakes. Their vegetative cover is variable. Shallow lakes are open water bodies a few metres deep. Only considered foe temperate and tropical regions; in polar and boreal regions it is difficult to separate shallow lakes from bogs and fens. Ricefields exclude upland rice. The physical area is calculated from the sum of irrigated rice (51 Mha of which 47 % is double- or triple-cropped with rice and 33 % under rice–wheat rotation), rainfed lowland rice (54 Mha) and deepwater rice (4 Mha). The riceland of China is taken to be all temperate. Sources: Aselmann and Crutzen (1989); Huke and Huke (1997); IRRI (2002).

3

Biogeochemical Characteristics Table 1.2

Rice ecosystems in the main rice-producing countries in Asia Harvested area (kha) Irrigated WSa

DSa

India 15 537 4123 China 20 490 9146 Indonesia 2963 2963 Bangladesh 351 2267 Thailand 274 665 Vietnam 1630 1630 Myanmar 1812 1386 Philippines 1175 1029 Pakistan 2125 0 Cambodia 140 165 Nepal 706 24 Korea, Rep. of 776 0 Sri Lanka 377 251 Total 49 211 24 003

Rainfed lowland 0–30b

30–100b

11 985 1990 2872 3271 6382 1963 2033 911 0 1069 406 326 213 34 056

4447 0 1006 2873 1778 651 478 341 0 349 166 0 26 12 131

Flood-prone

Upland

Total

1364 0 2 1220 342 177 362 0 0 152 118 0 0 3737

5060 499 1209 697 203 322 214 165 0 24 68 1 0 8853

42 516 32 125 11 015 10 679 9644 6373 6285 3621 2125 1899 1488 1103 867 13 1991

a

Wet/dry season. Depth of floodwater (cm). Definitions of ecosystem types: Irrigated. Grown in levelled, bunded fields with good water control. Crop is transplanted or direct seeded in puddled soil, and a shallow floodwater is maintained on the soil surface so that the soil is predominantly anoxic during crop growth. Rainfed lowland. Grown in level to gently sloping, bunded fields that are flooded for at least part of the cropping season. Water depths exceed 100 cm for no more than 10 consecutive days. Crop is transplanted in puddled soil or direct seeded on puddled or ploughed dry soil. During season soil alternates between oxic and anoxic conditions of variable duration and frequency. Flood-prone. Distinguished from rainfed lowland rice by extent and duration of flooding. Fields are flooded to at least 100 cm and often much more for at least 10 consecutive days in the growing season. Crop is transplanted in puddled soil or direct seeded on ploughed dry soil; soil may alternate between oxic and anoxic conditions during season. Upland. Grown in level to steeply sloping fields that are rarely flooded. No effort is made to impound water as for other rice ecosystems. Crop is direct seeded on ploughed dry soil or dibbled in wet, non-puddled soil. Source: IRRI (2002). Reproduced by permission of IRRI. b

‘ecosystems’ distinguished by land and water characteristics and adaptations of the rice plant to them. These are defined in Table 1.2 together with their extents in the main rice-producing countries in Asia.

1.2 BIOGEOCHEMICAL CHARACTERISTICS Wetlands are intermediate between upland systems and true aquatic systems, both in terms of their hydrologies, being intermittently to permanently flooded, and in terms of their biogeochemistries, being sources, sinks and transformers of

4

Introduction

nutrients and carbon, whereas uplands tend to be sources and aquatic systems sinks. Three types of wetland are distinguished based on hydrology (Figure 1.2): a. fluxial, which receive water wholly or in part from surface flow, such as in runoff or streams; b. phreatic, which receive water from groundwater that rises to the soil surface for at least part of the year; and c. pluvial, which receive water entirely from rainfall. In fluxial wetlands water flowing in from neighbouring upland brings with it sediment and nutrients which are only slowly lost to deepwater areas downslope, and may be supplemented by seasonal inflow from deepwater areas. Because of the net inflow of nutrients, the abundance of water, and beneficial changes in the soil resulting from chemical reduction under anoxia, fluxial wetlands are among the most productive ecosystems on Earth. By contrast pluvial wetlands rely on nutrients brought in by rainfall or fixed biologically from the atmosphere, and they therefore tend to be much less productive. Phreatic wetlands are intermediate. As sources, sinks and transformers of matter and energy, wetlands have important roles in element cycles at local, regional and global scales. They contribute to the global stability of carbon dioxide, methane and sulfur in the atmosphere and of available nitrogen and phosphorus in surface waters, and they are important regionally as sinks for organic and inorganic pollutants released into them accidentally or otherwise. These topics are introduced in the following sections; all are returned to in greater detail later in the book. Carbon Balances in Wetlands Table 1.3 shows the net primary production of different wetlands compared with upland and aquatic ecosystems. The generally greater productivity of wetlands is evident. Net primary production (NPP) is the gross rate of carbon fixation in photosynthesis less the rate of loss in plant respiration. The chief factors governing NPP are radiation, temperature, water, nutrients and toxins. Hence for a given type of wetland, NPP tends to increase from polar to tropical regions as incident radiation and day length increase; correspondingly nutrients and temperature become increasingly limiting. There are of course interactions between these changes. For example, the greater productivity of temperate compared with tropical ricefields on a per crop basis shown in Table 1.3 arises because of interactions between radiation and temperature: in temperate rice areas with high radiation during the growing season, low night-time temperatures result in lower respiratory losses compared with tropical areas and hence greater net productivity. Because of their often high biological productivity and low rates of decomposition under anoxia, wetlands are one of the largest terrestrial sinks for carbon. They account for about a third of the soil carbon globally (Table 1.4). However there are large differences between wetland types. Organic wetland soils tend

Figure 1.2

Generally high

Generally low

Net import

Net export or import, and transformation

Net export

Low to medium

Permanently deeply flooded

…… …… … …… … …

…… … …… ……

Intermittently -permanently flooded

Aquatic

Wetland

Dry

Upland

Upland Wetland

Generally medium to high

Net import and transformation

Intermittently -permanently flooded

…… …… ……… … … …

(b) Phreatic

Wetlands in (a) fluxial, (b) phreatic and (c) pluvial landscapes

Energy conversion

Nutrient regime

Water regime

(a) Fluxial Upland

Upland

(c) Pluvial Wetland

Generally low

Transformation

Intermittently -permanently flooded

… …… …… …

Upland

5

6 Table 1.3

Introduction Net primary production of wetlands compared with other ecosystems Net primary production (g C m−2 year−1 )

Wetlands Bogsa Fensa Swampsa Marshesa Floodplainsa Shallow lakesa Wetland riceb Others Forestc Grasslandc Arabled Desertc

Polar

Boreal

Temperate

Tropical

100–300 100–300

300–700 400–700 500–1000

400–800 400–1200 700–1500 800–2000 800–1800 400–600 850

600–1200

430

650 320 750 50 µm

0.5–50 µm

Pz (2.34) where Pz is the hydrostatic pressure at depth z (= Patm + ρgz). This inequality may be met either because of accumulation of dissolved gases or because of changes in pressure, as for example when a core of mud is brought up from an anaerobic sediment. In flooded soil or sediment, bubbles form through heterogeneous nucleation at the surface of solid particles, rather than by homogeneous nucleation in free solution. Because of this, bubbles form easily and the sum of the partial pressures of volatile solutes tends to be maintained at or near the hydrostatic pressure. Therefore, for a methanogenic sediment, PN2 + PCO2 + PCH4 + PH2 O = Patm + ρgz

(2.35)

This equation can be used to calculate the composition of bubbles and rates of ebullition from rates of gas formation and the volatility of the different species. Thus for a methanogenic sediment in which rates of CH4 and CO2 generation are balanced by their rates of loss to the atmosphere above by diffusion and ebullition, we have for each volatile solute (cf. Morel and Herring, 1993, Equations 144–146) 

CZ /KH CZ − C0 +ε −R =0 (2.36) D Z Patm + ρgZ

39

Mixing by Soil Animals

FE = 2.9

FD = 2.1

FE = 0.3

FD = 4.7

FE = 0.8

FD = −0.8

Depth in overlying water (cm)

0 10

CH4

CO2

N2

20 30 40 50 0.0 0.3 0.6 0.9 0.0 0.3 0.6 0.9 1.2 1.5 1.8 2.1 2.4 0.0 0.3 0.6 Concentration of dissolved gas in water (mM)

Figure 2.10 Concentrations and fluxes of CH4 , CO2 and N2 in anoxic acidic marsh (after Morel and Herring, 1993). FE and FD are the fluxes by ebullition and diffusion, respectively Reproduced by permission of Wiley, New York

where Z is the depth of overlying water, D is the diffusion coefficient of the solute in water, CZ and C0 are the concentration of dissolved solute at sediment surface and water surface, respectively, ε is the rate of ebullition of all gases together, KH is Henry’s law constant and R is the rate of generation of the solute in the sediment. An equation of this type can be written for N2 , CH4 and CO2 and combined with Equation (2.35) and the resulting equation solved to obtain the rates of ebullition and the concentrations of each gas at the sediment surface given the ambient atmospheric concentrations, the rate of methanogensis and the depth of the water. Figure 2.10 compares the relative contributions of ebullition and diffusion to fluxes of CH4 , CO2 and N2 in an anoxic marsh so calculated. The figure shows that CO2 escapes mainly by diffusion whereas more than half the CH4 escapes by ebullition. The bubbles contain 69 % CH4 , 19 % N2 , 5 % H2 O and only 7 % CO2 . In practice gas bubbles may become entrapped under irregularly shaped soil particles, and so the simple steady state described by Equation (2.36) does not hold. The rate of ebullition is then sensitive to mechanical disturbances, induced for example by wading animals or by the action of wind on plants in the sediment. This is discussed further in Chapter 8. 2.4 MIXING BY SOIL ANIMALS The upper few centimetres of the soil are subject to mixing by invertebrates burrowing through the soil and ingesting soil particles. If populations are sufficiently

40

Transport Processes in Submerged Soils

dense, this may have a large effect on solute transfer between the soil and overlying water. Oligochaete worms are often present in submerged soils in populations exceeding several thousand per m2 with burrows extending to several centimetres (Chapter 5). Once the burrows are constructed, the worms remain in them feeding with their heads downward and their posterior ends upward in the overlying water. By waving their posteriors and moving their bodies in a peristaltic motion they cause the water in the burrows to be mixed with the overlying water. Solid particles also fall into the burrows and are mixed. Hence solutes diffusing into a burrow will be rapidly transferred to the surface, and vice versa. The ecology of tubificids and other organisms in the soil and floodwater are discussed in Chapter 5. I here discuss approaches to modelling their effects on solute transfer between the soil and floodwater. Three approaches have been taken to the analogous problem of mixing by invertebrates in marine sediments (Aller, 1980a; Berner, 1980). The simplest approach has been to lump together all the processes involved and to assume that mixing is random and complete to a specified depth. This has been applied successfully to the long-term mixing of sediments under the combined action of invertebrates and waves or currents, but is inappropriate for less perturbed systems and short times. A second approach has been to express the effect of burrowing as increased effective diffusion coefficients of solutes in the pore water, derived by fitting diffusion equations to empirical data (Aller, 1980a; Berner, 1980; van Rees et al., 1996). But the physical basis of this approach is doubtful. A third approach was developed by Aller (1980a, b) who studied solute fluxes in near-shore marine sediments showing seasonal variation. In this approach, the geometry of the burrow–sediment system is allowed for explicitly and transport in the sediment between the burrows is described with appropriate diffusion equations. It is assumed that the burrows are oriented normal to the sediment surface and distributed uniformly or randomly in the horizontal plane (Figure 2.11). Thereby a cylindrical zone of influence is ascribed to each burrow with a radius

burrow (radius = r1)

zone of influence (radius = r2 = 1/√πN)

Figure 2.11 Distribution of worm burrows and cylinders of influence represented by boundary conditions for Equation (2.37)

41

Mixing by Soil Animals

such that the whole sediment volume is accounted for. The water in a burrow is assumed to mix instantaneously with the overlying seawater, and solutes diffuse radially between the burrow and the sediment surrounding it as well as vertically between the sediment and overlying water. The corresponding continuity equation for transport in the sediment influenced by a particular burrow is

 ∂ ∂C ∂C 1 ∂ ∂C = D + rD +R (2.37) ∂t ∂z ∂z r ∂r ∂r where z is the distance from the sediment surface, r is the radial distance from the centre of the cylinder and R is the rate of production or consumption of solute in the sediment. In Equation (2.37), the first term on the right-hand side accounts for diffusion in the vertical direction; the second term accounts for radial diffusion across the cylinder. The following boundary conditions apply. At the sediment–water and sediment–burrow interfaces, the concentrations are the same as in the overlying water: z = 0 C = C0 r = r1

C = C0

At the boundary between adjacent cylinders, there is effectively no transfer of solute: r = r2 dC/dr = 0 √ where the radius of the cylinder, r2 , = 1/ πN , where N is the density of worms per unit sediment surface area. At the bottom of the cylinder, the flux of solute is constant: z = L dC/dz = B The value of B is specified from empirical observations. Aller (1980b) shows that if the mean distance between burrows is small compared with their length, then a steady state (∂C/∂t = 0) will be attained rapidly, and he provides an analytical solution of Equation (2.37) for the steady state subject to the above boundary conditions. (The solution is complicated, involving Bessel functions, and is not reproduced here.) The mean concentration at a particular depth is found by integrating the concentration across the cylinder of sediment at that depth:  r2

2πrC.dr

r Cz = 1

(2.38)

r2

2πr.dr

r1

Aller uses the model to explain seasonally fluctuating profiles of NH4 + concentration in sediments in Long Island Sound. In this system NH4 + is produced

42

Transport Processes in Submerged Soils

in anoxic decomposition of organic matter in the sediment at a rate decreasing exponentially with depth, R = R0 exp(−αz) + R1

(2.39)

and it is removed by nitrification in the overlying water and in worm burrows. The rate of NH4 + formation and the density of the worms vary with seasonal temperature changes. Figure 2.12 shows concentration profiles of NH4 + in the sediments measured over 2 years and the corresponding profiles predicted by the model using independently measured parameter values. It shows that the main features of the profiles and their seasonal dynamics are satisfactorily predicted. By comparison, a model using the same parameter values but only allowing for diffusion in the vertical direction over-predicted the concentrations several fold. Aller found similar good agreement between observed and predicted profiles and fluxes of SO4 2− and Si in the sediments. He concluded that the model accounted satisfactorily for the important processes operating. The application of this approach is illustrated in Figures 2.13 and 2.14, which show the effects of tubificid worms on the movement of P between a submerged Concentration of NH4+ in solution (µM)

Depth (cm)

0

0

100 200 300 400

0

0

100 200 300 400

0

5

5

5

10

10

10

15

15

0

100 200 300 400

November 1974 0

0

100 200 300 400

March 1975 0

5

5

5

10

10

10

15

15

15

July 1975

100 200 300 400

15

July 1974 0

0

October 1975

0

100 200 300 400

March 1976

Figure 2.12 Concentration profiles of NH4 + at different times in sediments in Long Island Sound. Points are measured data; lines are predicted with Equations (2.37) and (2.39) using independently measured parameter values (after Aller, 1980a). Reprinted with permission from Elsevier

43

Mixing by Soil Animals

0.0

0

Concentration of P in solution (µM) 100 200 300 400 500 600

Depth (cm)

0.5 1.0

N= 30 000

1.5

15 000 5000

2.0 2.5

1000 3.0

0.8

L = 5 cm 0.6 3

0.4

1.5 0.2 0.0

0

10 000 20 000 N (m−2)

30 000

Ratio of soil surface flux to total

−

P flux (mmol m−2 day 1)

Figure 2.13 Effect of mixing of pore water by tubificid worms on profiles of P concentration in submerged soil calculated with Equations (2.37) and (2.40). Numbers on curves are densities of tubificids

1.0 0.8 0.6 1.5 0.4 3 5

0.2 0.0

0

10 000

20 000

30 000

N (m−2)

Figure 2.14 Effect of mixing by tubificids on flux of P between soil and floodwater calculated with Equations (2.37) and (2.40). Numbers on curves are depths of mixing

soil and overlying floodwater. The primary productivity of the floodwater, including the fixation of N by photosynthetic aquatic organisms, is often limited by the supply of P from the soil. So enhanced P transport resulting from tubificid activities can be important. Figure 2.13 shows calculated concentration profiles of P in the soil near the floodwater for realistic densities of tubificids (see Chapter 5) and other parameters, and Figure 2.14 shows the corresponding fluxes from the soil into the floodwater. Following Aller (1980b) for Si desorption in marine sediments, the rate of P desorption from the soil is calculated with the formula R = k(Ceq − C)

(2.40)

44

Transport Processes in Submerged Soils

where k is a rate constant and Ceq an apparent equilibrium P concentration. Values of k = 10−6 s−1 and Ceq = 0.5 mm were used for the calculations in Figures 2.13 and 2.14. The values of the other parameters used were θL = 0.6, √ fL = 0.4, b = 100, B = 0, r1 = 0.5 mm, r2 = 1/ πN , where the values of N are given in the figures and L = 3 cm. It will be seen that the tubificids have a large effect and the flux of P to the floodwater increased several fold for realistic numbers and dimensions. For comparison, the fluxes of P from the soil required to sustain typical rates of primary production in the floodwater in ricefields are in the range 0.05–0.25 mmol m−2 day−1 , calculated from measured primary production and the P contents of likely photosynthetic organisms given by Roger (1996). Because the tubificids depend upon the photosynthetic organisms for their carbon, there will be a positive feedback between mixing by tubificids and net primary production in the floodwater. Note that the sensitivity of the net flux between the soil and water to the worms’ activities depends on the relation between the rate R and the solute concentration. For the calculations in Figures 2.13 and 2.14, R varies linearly with concentration as specified in Equation (2.40), and the flux is sensitive to worm activity. But where the rate is independent of concentration, as for NH4 + formation in Equation (2.39), the net flux, which in this case is roughly R0 /α + LR1 , is necessarily independent of worm activity, though the distribution of the flux between burrows and the sediment surface and the concentration profile are not. In practice the rate will always depend to some extent on concentration. But the predictions here for the idealized steady state indicate the expected sensitivities.

3 Interchange of Solutes between Solid, Liquid and Gas Phases

This chapter is concerned with how ions and uncharged solutes in the water and soil solution in submerged soils interchange between the solid, liquid and gas phases present. This is a large topic. I give here the bare essentials needed to understand the transport and transformation processes discussed elsewhere in the book, and I give references to more detailed treatments where appropriate. The water and atmosphere overlying the soil are dealt with first and then the additional complexities in the soil. A. WATER 3.1 COMPOSITION OF THE WATER The water contains: • dissolved matter – free ions; – inorganic and organic complexes; – uncharged molecules. • particulate matter – large organic and inorganic polymers; – oxides; – clay minerals; – organic matter. Because of their large surface areas, charged particles are very efficient scavengers of ions from solution, and where the sediment load is large the concentration of adsorbed ions may greatly exceed the concentration in solution. Similarly for ions that form complexes with organic or inorganic ligands, their total concentration in solution may be far greater than the concentration of the free ion. Complexation and sorption are especially important in regulating the concentrations of trace metals in natural water systems. The interactions between ions and charged particles are discussed in the sections on soil. The Biogeochemistry of Submerged Soils Guy Kirk  2004 John Wiley & Sons, Ltd ISBN: 0-470-86301-3

46

Interchange of Solutes between Solid, Liquid and Gas Phases

3.1.1 ACIDS AND BASES The concentrations of dissolved species in natural waters depend ultimately on the dissolution of basic rocks–carbonates, silicates and aluminosilicates–induced by the action of weak acids in the water derived from dissolved gases–e.g. H2 CO3 derived from CO2 . Anions produced in acid–base reactions balance cations produced in dissolution reactions. The charge balance is:
m[cationm+ ] =
n[anionn− ]

(3.1)

Table 3.1 shows the main weak acids present in natural waters and typical concentration ranges. Table 3.2 shows the corresponding equilibrium constants. Table 3.1 shows that carbonic acid is by far the dominant acid with concentrations typically of the order of several mM. It arises from the dissolution of carbonate rocks and atmospheric CO2 , and from the respiration of aquatic and soil organisms. The concentrations of dissolved silica are 5–10 times smaller, and those of ammonium and orthophosphate smaller again, although NH4 + concentrations in the mM range may arise in the water in ricefields following fertilizer applications. The hydrolysis products of certain metals, such as Fe(III) and Al(III), also behave as weak acids and may be important under particular circumstances. Dissolved amino, organic and humic acids are rarely a large part of the charge balance in solution but may be important as metal ligands. The distributions of different acid–base pairs with pH are shown in Figure 3.1. Bicarbonate (HCO3 − ) is the dominant carbonate species at near neutral pH; silicic acid (H4 SiO4 ) is essentially undissociated at all pHs of interest; and the ammonium ion (NH4 + ) is the dominant form of ammoniacal-N at pHs below about 8. Orthophosphate and sulfide have acidity constants near neutral pH. For a given concentration of a particular dissolved acid, the proportions of the component species in the equilibrium solution will depend on the alkalinity of the solution; that is, the balance of cations and non-dissociating anions present. This can be calculated as shown in Table 3.3 for the aqueous carbonate equilibria Table 3.1

Concentrations of weak acids and bases in natural waters Global mean for freshwatera

Range in water

Range in submerged soil solutions

0.97 mM 0.22 mM 0–10 µM 0.7 µM — — —

0.01–10 mM 0.1–0.5 mM 0.001–1 mM 0.5–25 µM Trace Trace 0.001–1 mM

5–100 mM 0.1–1.5 mM 0.001–1 mM 0.5–100 µM 0.01–10 µM 0.1–10 µM 0.1–10 mM

Carbonate Silicate Ammonium Phosphate Sulfide Amino acids Organic acids Source: a Morel and Herring (1993).

47

Composition of the Water Table 3.2 Equilibrium constants for acid–base equilibria at 25 ◦ C, I = 0 Equilibrium H2 O = H+ + OH− CO2 (g) + H2 O = H2 CO3 ∗ H2 CO3 ∗ = H+ + HCO3 − HCO3 − = H+ + CO3 2− H4 SiO4 = H+ + H3 SiO4 − H3 SiO4 − = H+ + H2 SiO4 2− NH3 (g) = NH3 (aq) NH4 + = H+ + NH3 (aq) H3 PO4 = H+ + H2 PO4 − H2 PO4 − = H+ + HPO4 2− HPO4 2− = H+ + PO4 3− H2 S(g) = H2 S(aq) H2 S(aq) = H+ + HS− HS− = H+ + S2− CH2 NH2 COOH = H+ + CH2 NH2 COO− CH3 COOH = H+ + CH3 COO−

− log K 14.0 1.46 6.35 10.33 9.86 13.1 −1.76 9.24 2.15 7.20 12.35 0.99 7.02 13.9 9.78 4.76

with the equilibrium constants in Table 3.4. Similar calculations can be made for the other dissolved acids. Table 3.3 gives the equilibria in a closed system in which the total carbonate concentration, CT , is fixed. In an open system, such as the water on the surface of a submerged soil, CT is variable and the resulting changes in pH depend on the balance of charge between the non-carbonate anions and cations present. Likewise if a quantity of strong acid, HX, or base, MOH, is added to the solution, the equilibria will adjust so as to neutralize part of the H+ or OH− added and so buffer the change in pH. The changes in [H+ ] with alkalinity or dissolved CO2 can be found from (see Equation 9, Table 3.3): CB − CA = [HCO3 − ] + 2[CO3 2− ] + [OH− ] − [H+ ]

(3.2)

where CA is the concentration of non-carbonate anions after the addition of acid HX and CB is the concentration of cations after addition of base MOH. If CB > CA , the difference CB − CA is the alkalinity of the solution; if CA > CB , the difference CA − CB is the mineral acidity. 3.1.2 SPECIATION Many ions and uncharged molecules are present in solution as more than one species, depending on the concentrations of ligand ions and molecules and the

48

Interchange of Solutes between Solid, Liquid and Gas Phases (a)

HCO3−

H2CO3*

CO32−

(b)

H4SiO4 H3SiO4− H2SiO42−

log activity (arbitrary units)

H2CO3

(c)

NH4+

NH3

(d)

H2PO4−

HPO42− PO42−

H3PO4

(e)

(f)

HS−

H2S

CH3COO−

CH3COOH

S2−

4

5

6

7

8

9

10

11

4

5

6

7

8

9

10

11

pH

Figure 3.1 Distribution of dissolved acid–base species at constant total concentration in solution (after Morel and Herring, 1993). Reproduced by permission of Wiley, New York

solution pH and ionic strength. Complexation of metals with ligands can result in the total concentration of the metal being far greater than the concentration of the free ion. This topic is covered in detail by Morel and Herring (1993), Stumm and Morgan (1996) and, for chelation by humic substances, by Tipping (2002). A complex is a species in which a metal atom or ion is attached by coordinate bonds to one or more ligand ions or uncharged molecules. The complex may itself be positive, negative or uncharged. In forming a coordinate bond the ligand donates a pair of electrons to the metal. In so doing the ligand is acting as a

49

Composition of the Water Table 3.3

Equilibria in aqueous carbonate solutions

Species

Equilibriab

CO2 (g), CO2 (aq), H2 CO3 , HCO3 − , CO3 2− , H+ , OH− , M+ , X− [H2 CO3 ∗ ]a = [CO2 (aq)] + [H2 CO3 ] [CO2 (aq)]/[CO2 (g)] = H c [H2 CO3 ∗ ]/pCO2 = KH [CO2 (aq)]/[H2 CO3 ] = K [H+ ][HCO3 − ]/[H2 CO3 ] = KH2 CO3 [H+ ][HCO3 − ]/[H2 CO3 ∗ ] = K1

(0) (0a) (1) (2) (2a)

CT = [H2 CO3 ∗ ] + [HCO3 − ] + [CO3 2− ]

(5)

[H+ ][CO3 2− ]/[HCO3 − ] = K2 [H+ ][OH− ] = KW Ionization fractions for constant total carbonate concentration, CT 2−

−

∗

(3) (4)

[H2 CO3 ] = α0 CT [HCO3 ] = α1 CT [CO3 ] = α2 CT 
K1 K2 −1 K1 α0 = 1 + + + + 2 [H ] [H ] −1
+ K2 [H ] α1 = +1+ + K1 [H ]
+2 −1 [H+ ] [H ] α2 = + +1 K1 K2 K2 Electrical neutrality condition [H+ ] + [M+ ] = [HCO3 − ] + 2[CO3 2− ] + [OH− ] + [X− ] a

The ‘apparent’ concentration of H2 CO3 since [CO2 (aq)] ≫ [H2 CO3 ]. Equilibrium constants are defined at constant ionic strength. Dimensionless Henry’s law constant, in which [CO2 (g)] = PCO2 /RT . Source: Stumm and Morgan (1996). Reproduced by permission of Wiley, New York. b c

Table 3.4 Equilibrium constants for carbonate equilibria at 25 ◦ C, I = 0 Equilibrium CO2 (g) + H2 O = H2 CO3 H2 CO3 ∗ = H+ + HCO3 − HCO3 − = H+ + CO3 2− H2 O = H+ + OH− a

PCO2 in kPa.

Constant ∗

a

− log K

KH K1

3.47 6.35

K2 KW

10.33 14.0

(6) (7) (8)

(9)

50

Interchange of Solutes between Solid, Liquid and Gas Phases

Lewis base and the metal as a Lewis acid. A characteristic of ligands is that they have a lone pair of electrons which they can donate to empty electron orbitals on the metal. Some ligands also have empty p- or d-orbitals and can produce complexes in which a double bond is formed with the metal: a sigma bond by donation of the lone pair from the ligand to the metal and a pi bond by back donation of electrons on the metal to empty d-orbitals on the ligand. The term chelate is reserved for species involving polydentate ligands that form a ring of atoms including the metal. Inorganic and organic ligands contain the following electron donor atoms from Groups IVB to VIIB of the Periodic Table: C

N O F P S Cl As Se Br Te I

Formation of coordination complexes is typical of transition metals, but other metals also form complexes. The tendency to form complexes is a function of the metal’s electron configuration and the nature of its outer electron orbitals. Metal cations can be classified into types A and B based on their coordination characteristics, as shown in Table 3.5. A-type cations, which tend to be from the left side of the Periodic Table, have the inert-gas type electron configuration with largely empty d-orbitals. They can be imagined as having electron sheaths not easily deformed under the influence of the electric fields around neighbouring ions. B-type cations have a more readily deformable electron sheath. In consequence, A-type cations form complexes preferentially with the fluoride ion and ligands having oxygen as their electron donor atom. They are attracted to H2 O more strongly than to NH3 or CN− , and they do not form sulfides because OH− ions readily displace HS− or S2− ions. They tend to form sparingly soluble precipitates with OH− , CO3 2− and PO4 3− . By contrast, B-type cations coordinate preferentially with ligands containing I, S or N as electron donors. They may bind NH3 more strongly than H2 O and CN− more strongly than OH− , and they tend not to form complexes with the main functional groups in organic matter, which have O as electron donor. They form insoluble sulfides and soluble complexes with S2− and HS− . Table 3.6 shows the major inorganic species expected in a solution with a composition typical of natural fresh water. Some calculations for organic ligands in submerged soil solutions are given in Section 3.7.

3.1.3 EQUILIBRIUM CALCULATIONS Complete calculations of chemical equilibria in natural waters and soil solutions are complicated because such a large number of solutes, solids and gases are

51

Composition of the Water Table 3.5

Classification of metal ions

A-type metal cations

Transition-metal cations

B-type metal cations

Electron configuration of inert gas, low polarizability, ‘hard spheres’

One to nine outer shell electrons, not spherically symmetric

Electron number corresponds to Ni0 , Pd0 and Pt0 (10 or 12 outer shell electrons), low electronegativity, high polarizability, ‘soft spheres’

(H+ ), Li+ , Na+ , K+ , Be2+ , Mg2+ , Ca2+ , Sr2+ , Al3+ , Sc3+ , La3+ , Si4+ , Ti4+ , Zr4+ , Th4+

V2+ , Cr2+ , Mn2+ , Fe2+ , Co2+ , Ni2+ , Cu2+ , Ti3+ , V3+ , Cr3+ , Mn3+ , Fe3+ , Co3+

Cu+ , Ag+ , Au+ , Tl+ , Ga+ , Zn2+ , Cd2+ , Hg2+ , Pb2+ , Sn2+ , Tl3+ , Au3+ , In3+ , Bi3+

Ligands F > O > N = Cl > Br > I>S OH− > RO− > RCOO−

Ligands S > I > Br > Cl = N > O > F

CO3 2− ≫ NO3 −

PO4 3− ≫ SO4 2− ≫ ClO4 − Source: Stumm and Morgan (1996). Reproduced by permission of Wiley, New York.

involved. However general computer programs are available to perform such calculations using successive approximation (Melchior and Bassett, 1990; Mangold and Tsang, 1991; Sposito, 1994). WHAM (Tipping, 1994, 2002) gives particular attention to reactions involving humic substances. An important component of equilibrium calculations is the conversion between ion activities, which equilibrium constants refer to, and ion concentrations, which mass balance and electrical neutrality equations refer to. The conversion is made with activity coefficients defined by the relation: ai = γi Ci

(3.3)

Various empirical relations are available for calculating individual ion activity coefficients [discussed by Stumm and Morgan (1996) for natural waters and Sposito (1984a, b), for soil solutions]. In the calculations in this book I used the Davies equation: log γ = −AZ 2

 √

I √ − 0.3I 1+ I



(3.4)

where I is ionic strength (= 12
Ci Zi 2 ), Z is ionic charge and A = 1.82 × 106 (εT )−1.5 , where ε is the dielectric constant (A ≈ 0.5 for water at 25 ◦ C). This relation is valid for I < 0.5 m.

52

Interchange of Solutes between Solid, Liquid and Gas Phases

Table 3.6

Major inorganic species in representative natural water

Condition

Element

Major species

Fresh water [Mn+ /MT ]

B(III)

H3 BO3 , B(OH)4 −

V(V)

HVO4 2− , H2 VO4 −

Hydrolysed,

Cr(VI)

CrO4 2−

anionic

As(V)

HASO4 2−

Se(VI)

SeO4 2−

Mo(VI)

MoO4 2−

Si(IV)

Si(OH)4

Li Na

Li+ Na+

1.00 1.00

Mg K

Mg2+ K+

0.94 1.00

Ca

Ca2+

0.94

Predominantly free aquo ions

2+

Sr Cs

Sr Cs+

0.94 1.00

Ba

Ba2+

0.95

Be(II) Al(III)

BeOH+ , Be(OH)2 0 Al(OH)3 (s), Al(OH)2 + , Al(OH)4 −

Ti(IV) Mn(IV) Fe(III)

TiO2 (s), Ti(OH)4 0 MnO2 (s) Fe(OH)3 (s), Fe(OH)2 + , Fe(OH)4 −

Co(II)

2+

Co , CoCO3 2+

0

0

1.5 × 10−3 1 × 10−9

2 × 10−11 0.5

2+

Ni(II)

Ni , NiCO3 (Ni , NiCl)

0.4

Complexation with

Cu(II)

CuCO3 0 , Cu(OH)2 0

0.01

OH− , CO3 2− ,

Zn(II)

Zn2+ , ZnCO3 0 (Zn2+ , ZnCl)

0.4

−

Ag(I)

Ag+ , AgCl0 (AgCl2 − , AgCl)

0.6

Cd(II) La(III)

Cd2+ , CdCO3 0 (CdCl2 ) LaCO3 + , La(CO3 )2 −

Tl(I), Tl(III)

Tl+ , Tl(OH)3 0 , Tl(OH)4 −

2 × 10−21

Hg(II)

Hg(OH)2 0 (HgCl4 2− )

1 × 10−10

HCO3 , Cl

−

Pb(II) Bi(III)

0

+

PbCO3 (PbCl , PbCO3 ) Bi(OH)3

0.5 8 × 10−3

5 × 10−2 7 × 10−16

Fresh water conditions: pH = 8, Alk = 2 mM, [SO4 2− ]T = 0.3 mM, [Cl− ] = 0.25 mM, [Ca2+ ]T = 1 mM, [Mg2+ ]T = 0.3 mM, [Na+ ]T = 0.25 mM, O2 at saturation with air, I = 5 mM. Source: Stumm and Morgan (1996). Reproduced by permission of Wiley, New York.

53

pH Buffer Capacity

3.2 pH BUFFER CAPACITY The extent to which the pH of a solution is buffered against additions or removals of protons is measured by the solution’s pH buffer capacity. This is defined as the amount of strong acid or base required to produce unit change in pH. The buffering depends on the transfer of protons between donors and acceptors, i.e. Bronsted acids and bases, which form conjugate acid–base pairs. The pH buffer capacity of a solution is calculated from the buffer capacities of the individual acid–base pairs present. Consider a generic acid–base pair HX–X− representing the various acid–base pairs in a solution. The pH buffer capacity of the HX–X− pair is defined as bHX =

d[X− ] d pH

(3.5)

The total concentration of the pair is [HX] + [X− ] = [Xtotal ] and from the acidity constant, K, [HX] = [H+ ][A]/K, hence [X− ] =

[Xtotal ] [H ]/K + 1

(3.6)

+

Substituting in Equation (3.5) for [X− ] from Equation (3.6)   d 1/ [H+ ]/K + 1 bHX = [Xtotal ] d pH

(3.7)

Hence bHX or

  d[H+ ] d 1/ [H+ ]/K + 1 K[H+ ] = [Xtotal ] = 2.303[Xtotal ]  2 + d pH d[H ] K + [H+ ] bHX = 2.303

[HX][X− ] [HX] + [X− ]

(3.8)

The total buffer capacity of the solution is equal to the sum of the buffer capacities of the individual acid–base pairs present, each given by an equation like Equation (3.8). In an aqueous solution of acid HA, three acid–base pairs are present: HA–A− , H3 O+ –H2 O and H2 O–OH− . Because [H3 O+ ] and [OH− ] are both negligible compared with [H2 O], in Equation (3.8) [X− ] = ([HX] + [X− ]) for H3 O+ –H2 O and [HX] = ([HX] + [X− ]) for H2 O–OH− . Hence bH3 O+ = 2.303[H3 O+ ]

(3.9)

bOH− = 2.303[OH ]

(3.10)

−

Therefore
 [HA][A− ] bsolution = 2.303 [H3 O+ ] + [OH− ] + [HA] + [A− ]

(3.11)

54

Interchange of Solutes between Solid, Liquid and Gas Phases

If other acid–base pairs are present the buffer capacity is
 [HB][B− ] [HA][A− ] + − + bsolution = 2.303 [H3 O ] + [OH ] + + .... [HA] + [A− ] [HB] + [B− ] (3.12) Polyprotic acids can be treated as a mixture of monoprotic acids. For example, consider the diprotic acid H2 C which forms the acid–base pairs H2 C = HC− + H+ and HC− = C2− + H+ . The acidity constants are K1 = [H+ ][HC− ]/[H2 C] and K2 = [H+ ][C2− ]/[HC− ], respectively. The total buffer capacity of the solution is therefore   − 2− − [HC [H ][C ] C][HC ] 2 bsolution = 2.303 [H3 O+ ] + [OH− ] + + [H2 C] + [HC− ] [HC− ] + [C2− ] (3.13) For example, for a solution buffered by the CO2 –H2 O–H2 CO3 –HCO3 − system, application of Equation (3.13) gives bsolution = 2.303{[H3 O+ ] + [OH− ] + [α1 (α0 + α2 ) + 4α2 α0 ] CT }

(3.14)

where α0 , α1 and α2 are ionization fractions defined in Table 3.3. This equation predicts that the buffer capacity will pass through minima at pH 4–4.5 where [H3 O+ ] and [HCO3 − ] are both low, at pH 8.3 where [H2 CO3 ] and [CO3 2− ] are low, and at pH 10.5–11 where [HCO3 − ] and [OH− ] are low. At these pH ranges, changes in the concentrations of acids or bases in the solution will cause large pH changes.

3.3 EQUILIBRIUM WITH THE GAS PHASE The equilibrium distribution of a volatile solute between gas and liquid phases is described by Henry’s law. For the equilibrium A(l) = A(g) in a dilute solution at low gas pressure, [A(l)] = KH pA (3.15) where [A(l)] is the concentration of the dissolved gas in solution, pA is the partial pressure in the gas phase and KH is the Henry’s law constant. (At high concentrations or gas pressures, [A(l)] and pA are replaced by the corresponding activities and fugacities.) The constant is also sometimes expressed in dimensionless form, H, such that [A(l)] = H [A(g)] (3.16) where [A(g)] is the concentration in the gas phase. Hence H = RT KH

(3.17)

55

Equilibrium with the Gas Phase Table 3.7 Henry’s law constants for important gases in submerged soils at 25 ◦ C and typical partial pressures in the atmosphere Typical partial pressure

KH (M kPa−1 )

(kPa)

−6

N2 O2 CO2 CH4 NH3 H2 S NO2 NO N2 O

6.52 × 10 1.24 × 10−5 3.35 × 10−4 1.27 × 10−5 5.63 × 10−1 1.04 × 10−3 9.87 × 10−5 1.88 × 10−5 2.54 × 10−5

78 21 3.5 × 10−2 1.7 × 10−4 0.1–5 × 10−7 < 2 × 10−8 1–5 × 10−7 1–5 × 10−8 3 × 10−5

where R is the gas constant (= 8.314 kPa L mol−1 K−1 ) and T is the temperature (K). Values of KH for important gases in submerged soils are given in Table 3.7. For some volatile solutes, slow reactions influence the rate of equilibration between the gas and liquid phases. Generally the rate of gas transfer across the liquid–gas interface is the rate-limiting step, as discussed in Section 3.4. But there may also be slow hydration or other reactions in solution that must be allowed for. An important example is the hydration of CO2 , whose half-life may be comparable to rates of transfer of CO2 across the air–water interface. Kinetics of CO2 Hydration The kinetics of the hydration and dehydration reactions are slow in comparison with some processes in the water. The reactions are kf 1

−− −− → CO2 + H2 O − ← − H2 CO3

(3.18)

kb1

and

kf 2

− + −− −− → CO2 + H2 O − ← − HCO3 + H

(3.19)

kb2

and the corresponding rate law is −

   d[CO2 ]  = kf 1 [CO2 ] − kb1 [H2 CO3 ] + kf 2 [CO2 ][OH− ] − kb2 [HCO3 − ] dt (3.20)

56

Interchange of Solutes between Solid, Liquid and Gas Phases

Substituting from Table 3.3 for the equilibrium constant for dissociation of H2 CO3 , which is fast, − or

   d[CO2 ]  = kf 1 + kf 2 [CO2 ] − kb1 + kb2 KH2 CO3 [H2 CO3 ] dt −

d[CO2 ] kH CO = kCO2 [CO2 ] − 2 3 [H+ ][HCO3 − ] dt KH2 CO3

(3.21)

(3.22)

where kCO2 = kf1 + kf2 and kH2 CO3 = kb1 + kb2 KH2 CO3 . Equation (3.22) corresponds to the simplified scheme kCO2

fast

+ − −− −− → −−− −− → CO2 + H2 O ← ← − H + HCO3 − H2 CO3 −

(3.23)

kH2 CO3

That is, the hydration reaction is first order with respect to dissolved CO2 . The rate constant kCO2 = 0.025–0.04 s−1 (25 ◦ C) and activation energy ≈63 kJ mol−1 . For the dehydration reaction, kH2 CO3 = 10–20 s−1 (20–25 ◦ C) and activation energy ≈67 kJ mol−1 . 3.3.1 FLOODWATER CO2 DYNAMICS The pH of the water on the surface of a submerged soil often depends on the activity of photosynthetic organisms. Photosynthesis by aquatic plants and algae removes dissolved CO2 during the day, but at night the net respiratory activity of the organisms returns CO2 to the water and the concentration of dissolved CO2 and acidity increase: photosynthesis

−− −− −− −− −− −− → CO2 + H2 O − ← − CH2 O + O2

(3.24)

respiration

where CH2 O is organic matter produced in photosynthesis or consumed in respiration. As a result the pH may rise as high as 10 during the day but fall by two or three pH units at night. Figure 3.2 shows measured diurnal changes in pH and carbonate species in the floodwater of a ricefield. The relations between pH, alkalinity and carbonate equilibria are described by Equation (3.2). Equation (3.24) shows that photosynthesis and respiration do not affect the alkalinity of the water per se. The pH increases or decreases with the change in CT at constant alkalinity. The change in pH depends on the alkalinity as it affects the initial pH and the consequent acid–base system operating. At pHs below pK1 (= 6.3), CO2 (aq) is the dominant species and there is little change in pH with CT . Between pK1 and + pK2 (= 10.3)HCO− 3 is the dominant species and roughly 1 mol of H is consumed − + per C fixed in photosynthesis (HCO3 + H → CH2 O + O2 ), with a correspondingly greater pH change. At pHs above pK2 , CO3 2− is the dominant species and roughly 2 mol of H+ are consumed per C fixed (CO3 2− + 2H+ → CH2 O + O2 ),

57

Equilibrium with the Gas Phase 10 60

90

50 pH 40

80 15

10

5

30

20

Free CO2 H2CO3

CO32−

9

8

pH

HCO3−

Free CO2 content (mg L−1)

Percent mole fraction of H2CO3, HCO3− and CO32−

100

7

10

0 0 600 800 1000 1200 1400 1600 1800 2000 2200 Time

6

Figure 3.2 Diurnal changes in pH and concentrations of carbonate species in the floodwater in a ricefield (Mikkelsen et al., 1978). Reproduced by permission of Soil Sci. Soc. Am.

and the pH change is correspondingly larger again. Figure 3.3 shows calculated changes in pH for a sinusoidally varying floodwater [H2 CO3 ∗ ] over the day for two different alkalinities. The dissolved CO2 concentrations are the same in Figure 3.3(a) and (b); only the alkalinities differ. In principle, the alkalinity of the water will also be affected by the balance of nutrient ions consumed and released by organisms in the water. But in practice these have a minor affect compared with CO2 . The average composition of the algal biomass in natural waters is given by the Redfield formula (Redfield, 1934) as C106 H263 O110 N16 P. Therefore for the complete stoichiometry of algal photosynthesis and respiration, we have with NO3 − as the source of N 106CO2 + 16NO3 − + H2 PO4 − + 122H2 O + 17H+ = C106 H263 O110 N16 P + 133O2 and with NH4 +

(3.25)

106CO2 + 16NH4 + + H2 PO4 − + 106H2 O = C106 H263 O110 N16 P + 106O2 + 15H+

(3.26)

The corresponding changes in alkalinity are +17/106 = +0.16 molc per mol C fixed for NO3 − nutrition and −15/106 = −0.14 molc per mol C fixed for NH4 + nutrition. More significant changes in the alkalinity of ricefield floodwater are

58

Interchange of Solutes between Solid, Liquid and Gas Phases (a) [Alk] = 10 mM 1.0

9.0 8.5

30

pH

8.0

0.6

7.5

0.2

CO32−/C T

H2CO3*/C T

6.5

0.0

0

(b) [Alk] = 0.5 mM 1.0

40

Free CO2

0.8

HCO3−/C T

7.0

10

30

Free CO2 (mg L−1)

Ratio of H2CO3*, HCO3− or CO32− to C T

20 0.4

6.0 pH

Free CO2

0.8

40 HCO3−/C T

9.0 8.5 8.0

0.6 pH

20

7.5

0.4 0.2 0.0

7.0

H2CO3*/C T

10 6.5

6

8

10 12 14 16 18 Time (h past midnight)

20

0 22

6.0

Figure 3.3 Calculated diurnal changes in the pH and concentrations of carbonate species in ricefield floodwater for sinusoidally varying [H2 CO3 ∗ ] with (a) [Alk] = 10 mM, (b) [Alk] = 0.5 mM. The free CO2 concentrations are in mg L−1 to be consistent with Figure 3.2

caused by additions of nitrogenous fertilizers. Effects on pH again depend on initial pH and corresponding buffer systems operating. 3.4 GAS TRANSPORT ACROSS THE AIR–WATER INTERFACE The floodwater is for the most part not in equilibrium with the atmosphere because rates of production of volatile solutes in the water exceed rates of gas exchange across the air–water interface. In particular, during the day, rates of CO2 consumption and O2 production by photosynthesizing organisms are generally sufficient to cause undersaturation of CO2 and supersaturation of O2 . Conversely, at night, respiration causes depletion of O2 and supersaturation of CO2 . The underlying soil is also a large sink for O2 and source of CO2 . The resulting diurnal changes in dissolved CO2 can cause large changes in floodwater pH, often from near neutral at night to pH 10 during the day.

59

Gas Transport Across the Air–Water Interface turbulent bulk air

still air layer

dzG

still water layer

dz L

turbulent bulk water

Figure 3.4

The air–water interface

Two main approaches have been taken to modelling the air–water interface in natural systems so as to calculate rates of volatilization and dissolution (Liss and Slater, 1974; Frost and Upstill-Goddard, 1999; McGillis et al., 2001). In the simpler the interface is viewed as two thin still layers, one in the air and one in the water, separating well-mixed bulk phases (Figure 3.4). Transport across the still layers is by diffusion. The still layers arise because of the increased viscosity of the air and water near the interface. Their thicknesses depend on such factors as wind speed and surface roughness. Under turbulent conditions, the thickness of the still layers is reduced and rates of gas transport correspondingly increased. At steady state the fluxes across the layers are equal. Therefore, if the gas undergoes no reactions, we have from Fick’s first law F =−

DG DL (CG0 − CG ) = − (CL − CL0 ) δzG δzL

(3.27)

where subscripts G and L indicate the gas and liquid phases, respectively, and subscript 0 indicates the interface. The alternative approach considers that turbulent eddies periodically mix the surface layers with the bulk fluids. The flux across the interface is related to the concentration difference by a transfer coefficient equal to the square root of the diffusion coefficient divided by a characteristic time, τ , representing the frequency of mixing. Thus   DG DL F =− (CG0 − CG ) = − (CL − CL0 ) (3.28) τG τL Neither model accounts completely for the processes operating in the interface, and they provide similar fits to empirical data (Frost and Upstill-Goddard, 1999). However the first model has the advantage of conceptual simplicity and I use it in the following sections.

60

Interchange of Solutes between Solid, Liquid and Gas Phases

If the gas obeys Henry’s law, then CL0 = H CG0

(3.29)

where H is the dimensionless Henry’s law constant. Eliminating CG0 between Equations (3.27) and (3.29) gives F =

1 1 (CG − CL0 /H ) = (CL0 − CL ) kG kL

(3.30)

where kG (= DG /δzG ) and kL (= DL /δzL ) are transfer coefficients for the gas and liquid phases, respectively. Hence CL0 can be eliminated to obtain the following equation for the flux through the water film F =

1 (CG − H CL ) 1/kL + H /kG

(3.31)

In Equation (3.31), 1/kL is the resistance to transfer through the liquid film and H /kG is the resistance to transfer through the gas film. The relative importance of these resistances is given by the ratio resistance in gas phase kL =H resistance in liquid phase kG

(3.32)

Table 3.8 gives values of this ratio for important gases in submerged soils. Most of the gases are sparingly soluble, and the resistance in the liquid phase is much greater than that in the gas. This is because diffusion coefficients in water are two orders of magnitude smaller than those in air and because, for these gases, H is small. But for very soluble gases, such as NH3 , resistance in the gas phase may be limiting. Solubility varies much more between different gases than the diffusion coefficients and is therefore the main determinant of whether gas or liquid phase resistance is limiting. If H is less than about 5, transport in the Table 3.8 Relative importance of resistances in air (rG ) and in water (rL ) to gas transfer across an air–water interface at 25 ◦ C and 1 atm (Equation 3.32)

O2 CO2 CH4 H2 S N2 NH3 N2 O NO a

DG (dm2 s−1 )

DL (dm2 s−1 )

H

rG /rL a

2.05 × 10−3 1.55 × 10−3 2.20 × 10−3 1.66 × 10−3 2.04 × 10−3 2.19 × 10−3 1.55 × 10−3 2.04 × 10−3

2.26 × 10−7 1.93 × 10−7 1.73 × 10−7 2.00 × 10−7 2.02 × 10−7 2.49 × 10−7 1.98 × 10−7 2.55 × 10−7

3.08 × 10−2 8.29 × 10−1 3.16 × 10−2 2.57 1.62 × 10−2 1.39 × 103 6.29 × 10−2 4.65 × 10−2

3.40 × 10−6 1.03 × 10−4 2.48 × 10−6 3.09 × 10−4 1.59 × 10−6 0.159 8.03 × 10−6 5.81 × 10−6

For δzL = δzG and assuming gases do not react with water.

Gas Transport Across the Air–Water Interface

61

liquid phase is limiting; if it is greater than about 500, transport in the gas phase is limiting. However, this simple picture only applies to gases that do not undergo reactions in the boundary layers. For gases that do react, for example through hydration and acid–base reactions, the net flux depends on the simultaneous movement of all the solutes involved, and the flux will not be the simple function of concentration expressed in Equation (3.25). An example is CO2 , which reacts with water to form carbonic acid and carbonate species–H2 CO3 , HCO3 − and CO3 2− . The situation is complicated because the exchange of H+ ions in the carbonate equilibria results in a pH gradient across the still layer, and it is therefore necessary to account for the movement of H+ ions across the still layer as well as the movement of carbonate species. The situation is further complicated in the case of CO2 by the kinetics of hydration and dehydration, which may be slow in comparison with transport. 3.4.1 CO2 TRANSFER ACROSS THE AIR–WATER INTERFACE Under equilibrium conditions, the bulk of the dissolved CO2 is present as HCO3 − or CO3 2− or both if the pH is greater than about 6. Therefore, where a gradient of CO2 concentration exists across a solution, the net flux of CO2 will be greatly increased if there is rapid equilibration between the dissolved CO2 and carbonate species. Consequently, most plants and animals have evolved enzyme systems to catalyse the hydration–dehydration equilibria and the enzyme responsible—carbonic anhydrase—is present in most plant and animal cells. It is likely that this enzyme will often be present extracellularly in natural waters. This is because many aquatic plants use HCO3 − for photosynthesis under low CO2 conditions by catalysing the conversion of HCO3 − to CO2 outside the plasma membranes of leaf cells. The mechanism involves catalysis by extracellular carbonic anhydrase in conjunction with H+ extrusion across the plasma membrane (Graham et al., 1984; Tsuzuki and Miyachi, 1989). Since at least some forms of the enzyme are soluble, appreciable concentrations should arise in the water under intense algal growth, though the stability of the enzyme under high light and O2 conditions is unknown. The presence of carbonic anhydrase or similar enzymes catalysing CO2 hydration has been demonstrated in seawater with corresponding differences in rates of CO2 exchange (Berger and Libby, 1969). The following calculations show the range of effects from infinitely slow hydration–dehydration to infinitely fast. Emerson (1975) and Kirk and Rachhpal-Singh (1992) and have made calculations allowing for the kinetics of the uncatalysed hydration–dehydration reactions, giving intermediate results. We have for the flux of CO2 across the still air layer an equation of the type FG = −

DG (CG − CL0 /H ) δzG

(3.33)

62

Interchange of Solutes between Solid, Liquid and Gas Phases

At steady state the flux of CO2 gas must equal the net flux of dissolved CO2 and carbonate species, therefore FGC = FLC = FLH2 CO3 ∗ + FLHCO3 − + FLCO3 2−

(3.34)

where H2 CO3 ∗ represents CO2 (aq) + H2 CO3 . The fluxes of the uncharged solutes, CO2 and H2 CO3 , are given by equations of the type DLA (CLA − CLA0 ) (3.35) FLA = − δzL The fluxes of charged solutes depend on the diffusion potential arising from differences in the mobilities of individual ions, as well as on an ion’s own concentration gradient (Equation 2.21). The effect of diffusion potentials will be important if the carbonate species are a large part of the total ion concentration, as they often will be. Therefore we have for the net flux of ion B FLB = −DLB


DLi dCLi /dz dCLB + ZB CLB DLB dz
Zi 2 DLi CLi

(3.36)

where subscript i refers to all the co- and counter-ions in solution and ZB and Zi are the ionic charges. This gives FLB ≈ −

ZB DLB (CLB0 − CLB ) + (CLB + CLB0 )DLB Φ δzL 2

(3.37)

where Φ=


DLi dCLi /dz
Zi 2 DLi CLi

There is an equation of this type for each of the ions present. The principal cations and anions in floodwaters are generally Ca2+ , Cl− , 2− − HCO− 3 , CO3 and OH . Therefore we have five equations of type (3.37) for the fluxes of the five charged species, Equation (3.33) for CO2 gas and Equation (3.35) for H2 CO3 ∗ . These seven equations contain six unknowns—the concentrations of H2 CO3 ∗ and the five ions in solution at the interface—and these are found with the following six equations: Equation (3.34), Equations (1)–(3) in Table 3.3, and FLCa2+ = 0 and FLCl− = 0—i.e. no net flux of Ca2+ and Cl− across the interface, their concentration gradients being balanced by their diffusion potential gradients. Note that charge balance between the diffusing ions is inherent in Equation (3.37). Note also that the movement of H+ ions formed in the carbonate equilibria is allowed for in the movements of the various conjugate acid–base pairs present: H2 CO3 –HCO3 − , HCO3 − –CO3 2− and H2 O–OH− . For each mol of CO2 entering or leaving the water, 1 mol of H+ is added or removed at the

63

Gas Transport Across the Air–Water Interface

interface and transferred to or from the bulk solution by the diffusion of conjugate acid–base pairs. Thus (3.38)

FGC = FLH2 CO3 ∗ − FLCO3 2− − FLOH−

which is inherent in the mass and charge balances. Figure 3.5 shows calculated concentration profiles in the still water layer for realistic conditions in ricefields. In figure 3.5(a) the CO2 pressure is large, the pH in the bulk solution correspondingly low (pH 6.7), and the movement of dissolved CO2 to the interface primarily as H2 CO3 ∗ . The loss of CO2 raises the pH at the interface (to pH 8.2), tending to offset the depletion of HCO3 − and the gradient of HCO3 − concentration is small. In figure 3.5(b) the CO2 pressure is small, the pH in the bulk solution correspondingly high (pH 10.6), and the movement of dissolved CO2 away from the interface is primarily as HCO3 − . Dissolution of CO2 lowers the pH at the interface (to pH 8.3) and there is therefore a gradient of decreasing OH− towards the interface. The gradient of CO3 2− is also negative. Since the mobility of OH− is about five times that of HCO3 − and CO3 2− , there is therefore an excess negative potential at the interface and as a result Ca2+ diffuses to the interface and Cl− away. (a) 0.00 0.0

0.25

Concentration (mM) 0.50 0.75

Depth (mm)

Cl−

Ca2+

1.00

1.25

HCO3−

OH− H2CO3*

0.1 (b) 0.0 Cl−

Ca2+

Depth (mm)

H2CO3*

HCO3−

CO32− OH−

0.1

Figure 3.5 Profiles of CO2 , HCO3 − , etc. across still water layer. Still layer thickness both = 1000 µm, [Ca2+ ]L∞ = 0.5 mM, [Cl− ]L∞ = 0.15 mM, PCO2 L∞ = 1 kPa (a), 2.5 × 10−5 kPa (b)

64

Interchange of Solutes between Solid, Liquid and Gas Phases

CO2 flux (kg C ha−1 h−1)

15

10

with CO2 hydration

5

0 without −5 −10 0.01

0.1

1 10 100 1000 CO2 pressure (Pa)

Figure 3.6 Flux of CO2 as a function of CO2 pressure with and without carbonate equilibria

Figure 3.6 shows how the flux of CO2 across the interface varies with CO2 pressure in the bulk solution, with and without equilibration between CO2 and carbonate species in the boundary layer. A positive flux indicates dissolution and a negative flux volatilization. The figure shows that the effect of the carbonate equilibria is very marked at small CO2 pressures, but insignificant at large pressures where transport across the boundary layer is primarily as H2 CO3 ∗ . At small CO2 pressures the rate of dissolution is enhanced many fold by the carbonate equilibria, the effect increasing as the CO2 pressure decreases and the pH of the bulk solution correspondingly increases. An important practical problem in ricefields is the loss of N fertilizer through volatilization of NH3 from the floodwater. Loss of NH3 is sensitive to the pH of the floodwater, and hence is intimately linked to the dynamics of dissolved CO2 (Bouldin and Alimago, 1976). To quantify this it is necessary to consider the simultaneous transfers of CO2 and NH3 across the air–water interface and their coupling through acid–base reactions. There is an equation of type (3.33) for the flux of NH3 across the still air layer and, as for the dissolved CO2 and carbonate species, the flux across the still water layer is FGN = FLN = FLNH4 + + FLNH3 + FLNH4 OH

(3.39)

The acid–base pairs involved are NH4 + –NH3 and NH4 + –NH4 OH, in addition to those listed above, and we have FGC − FGN = FLH2 CO3 − FLCO3 2− − FLNH3 − FLNH4 OH − FLOH

(3.40)

Equation (3.40) is inherent in the mass and charge balances. These equations can be solved as before to calculate the simultaneous fluxes of CO2 and NH3 across the air–water interface.

The Solid Surfaces in Soils

65

B. SOIL In addition to the factors considered for water, we need to consider for soil: (a) the far greater importance of interactions with solid surfaces and the buffering of ions in solution by ions adsorbed on the surfaces; and (b) the more-strongly reducing conditions that develop in soil because of the greater sink for O2 , resulting in transformations of soil surfaces as well as of species in solution. Figure 3.7 shows the concentrations of cations in solution following submergence of four representative rice soils; the corresponding changes in EH , pH, HCO3 − , CEC and soil Fe are shown in Figures 2.6 and 2.7. The main anion in solution is HCO3 − derived from CO2 together with Cl− : any NO3 − ions present before submergence are rapidly consumed in reduction and consequently their concentration is generally negligible after the initial stages, and SO4 2− ions are also reduced though more slowly. These unadsorbed anions determine the overall strength of the soil solution and balance cations derived from exchange, dissolution and redox reactions involving the soil solid. Redox processes are discussed in detail in Chapter 4. The rest of this chapter deals with solid–solution interactions, firstly for soils in general and then for submerged soils. Recent reviews of solid–solution interactions in soils include Sposito (1994), Sparks (2003) and the relevant chapters of Sumner (2000).

3.5 THE SOLID SURFACES IN SOILS The main surfaces with which ions interact are clay-sized particles of layer silicates and Al, Fe and Mn oxides, and organic matter bound to clay particles. Their interactions with ions depend on their functional groups. These are analogous to the functional groups on molecules in solution but differ in that they are held a fixed distance apart and their charge characteristics may be strongly influenced by the neighbouring functional groups.

Layer Silicates The layer silicates comprise tetrahedral sheets of silica and octahedral sheets of aluminium and magnesium hydroxide, with varying amounts of the Si, Al and Mg replaced by cations of lower valence giving the lattice a net negative charge. Two basic combinations occur: 1 tetrahedral sheet with 1 octahedral (e.g. kaolinite, halloysite), and 2 tetrahedral with 1 octahedral (e.g. smectite, vermiculite, illite). 1:1 Type. The tetrahedral and octahedral sheets are bound together because they share an oxygen atom, and the resulting layers are bound together by hydrogen bonds between the oxygens of the tetrahedral sheet and the hydroxyls of the adjacent octahedral sheet. The structure is therefore rigid. Pure kaolinite, formula

66

Interchange of Solutes between Solid, Liquid and Gas Phases 7 Iloilo

6

Fe2+

5 4 Ca2+

3

Mg2+

NH4+

2

Na+

1

K+

Mn2+

0 16 Maahas

14 Concentration in soil solution (mmolc L−1)

12 10

Ca2+

8

Mg2+

6 4

Na+

Fe2+

2 0 20

Nueva Ecija 15 Ca2+ 10 Mg2+ 5

Na+ Fe2+

0 6 5

Ca2+

Tarlac

4 3

Mg2+

2

0

Na+

Fe2+

1 0

20

40

60

80

Time after flooding (days)

Figure 3.7 Changes in the concentrations of cations in solution following flooding of four rice soils (Kirk et al., 2003). The corresponding changes in EH , pH, HCO3 − , CEC and soil Fe are shown in Figures 2.6 and 2.7. Reproduced by permission of Blackwell Publishing

67

The Solid Surfaces in Soils Table 3.9

Structural charge and surface area of layer silicates

Group

Layer structure

Typical formulaa

Structural Specific negative surface charge area −1 (molc kg−1 ) (m2 kg )

Kaolinite 1:1 [Si4 ]Al4 O10 (OH)8 0–0.01 Halloysite 1:1 [Si4 ]Al4 O10 (OH)8· 4H2 O 0–0.01 Illite 2:1 M1.4 – 2 [Si6.8 Al1.2 ]Al3 Fe0.25 Mg0.75 O20 (OH)4 1.9–2.8 Vermiculite 2:1 M1.2 – 1.8 [Si7 Al]Al3 Fe0.5 Mg0.5 O20 (OH)4 1.6–2.5 Smectite 2:1 M0.5 – 1.2 [Si8 ]Al3.2 Fe0.2 Mg0.6 O 20 (OH)4 0.7–1.7 Chlorite 2:1:1 (Al(OH)2.55 )4 [Si6.8 Al1.2 ]Al3.4 Mg0.6 O20 (OH)4 Variable a

0.5–3 1–4.5 8–15 30–50 60–80 2.5–15

M = monovalent cation.

Si4 Al4 O10 (OH)8 , has no structural charge, but moderate structural charge (< − 10 mmolc kg−1 ) arises in most natural kaolinites as a result of substitution of Mg2+ for Al3+ or Al3+ for Si4+ (Table 3.9). More important is the pH-dependent charge that arises as a result of adsorption and desorption of protons by the –OH and –O groups at the lattice edges. This is in the range +10 to −50 mmolc kg−1 over the usual range of soil pHs, with zero charge at about pH 4.6. 2:1 Type. The tetrahedral sheets of adjacent layers cannot form hydrogen bonds with each other, lacking –O groups, but are held together by electrostatic forces arising from their charge. In illites the interaction is strong because the lattice charge is neutralized by K+ ions, which effectively glue the sheets together. In smectites and vermiculites water molecules and solutes can enter between the layers, increasing the layer spacing. Swelling clay minerals have a moderate proportion of substituted atoms and hence a relatively low charge density (Table 3.9). The cations that compensate the permanent charge are therefore bound to the surface by relatively weak purely electrostatic forces. They form a broad diffuse layer in the solution near the surface. Non-swelling 2:1 clay minerals have a much larger proportion of substituted atoms and therefore greater charge density (Table 3.9). As a result the chargecompensating cations may form covalent or ionic bonds with the surface, as well as being held electrostatically. Certain ions, e.g. K+ and NH4 + , fit neatly in holes in the clay structure and produce a rigid stack of clay layers. Only the cations held on outer surfaces and, to a variable extent, on imperfectly fitting layers within the clay structure are then exchangeable with cations in the soil solution. 2:1:1 Type. The interlayer spaces of 2:1 silicates may be blocked by poorly ordered sheets of Al hydroxy polymers, such as [Al(OH)2.5 0.5+ ]n (n ≥ 6). Such Al interlayers neutralize a considerable part of the surface charge and restrict swelling, and effectively convert 2:1 clays into materials similar to kaolinite.

68

Interchange of Solutes between Solid, Liquid and Gas Phases

Even moderate Al interlayering greatly affects clay properties. Such materials are important in submerged soils that have become ferrolysed (Section 7.1). Amorphous Aluminosilicates. These occur in soils influenced by volcanic activity and are associated with very high moisture retention and anion fixation, and low to very high pH-dependent CEC. They may also bind organic matter tightly, protecting it against decomposition. Examples are allophane and imogolite. Oxides Being widespread in the lithosphere and insoluble in the usual range of soil pH, oxides and hydroxides of Al, Fe and Mn are common in soil clays. Red or yellow coloration of soils is apparent at Fe oxide contents of only 0.1 % or less, especially if the Fe is amorphous and coats other minerals. The most visible change occurring when soils are submerged is the conversion of the red and yellow compounds of Fe(III) to the bluish-grey compounds of Fe(II). Metal oxides and hydroxides have little or no structural charge but develop pH-dependent charge as the hydroxyl groups at the lattice edges gain or lose protons. The surface charge is a function of both pH and the concentration of salts in the solution, as these affect the dissociation of the –OH groups. However the pH at which the surface negative charge is equal to the surface positive charge—the point of zero charge (pzc)—is independent of the salt concentration if the salt does not react with the surface. The point of zero charge is an important characteristic of the surface. Table 3.10 gives pzc values for common soil materials. Values are large for metal oxides and hydroxides but small for silica and soil organic matter. In real soils where oxides, layer silicates, organic matter and other materials are present in intimate mixtures, with the oxides and organic matter often coating the surfaces of the other materials, the different functional groups interact with Table 3.10 Points of zero charge (pzc) of oxides and aluminosilicates Material α-Al(OH)3 γ -AlOOH α-FeOOH γ -Fe2 O3 Amorphous Fe(OH)3 MgO δ-MnO2 SiO2 Feldspars Kaolinite Montmorillonite 7:05 pm, Jan 18, 2005

pzc 5.0 8.2 7.8 6.7 8.5 12.4 7.2 2.0 2–2.4 4.6 2.5

The Solid Surfaces in Submerged Soils

69

each other and so the distinction between permanent and pH-dependent charge is blurred. Organic Matter Soil organic matter is a weak acid and becomes negatively charged by losing protons. The main functional groups are carboxylates and, to a lesser extent, phenols, which are weaker acids. Their acid–base behaviour is complicated because of their heterogeneity and because of the effects of neighbouring functional groups on soil surfaces. With increasing dissociation, the build up of negative charge on the surface tends to inhibit further dissociation. Thus a plot of the extent of dissociation of organic functional groups versus pH tends to be steeper than the equivalent plot for simple monoprotic acids, and it approaches a straight line over the usual pH range in soils. This leads to the following rough empirical relation for the negative charge on soil organic matter as a function of pH (McBride, 1994): SOM charge(mmolc g−1 organic C) = −0.6 + 0.5 pH

(3.41)

As a rule of thumb, at near neutral pH, each g of organic C per kg of soil increases the surface negative charge by about 3 cmolc kg−1 soil. A further complication is that soil organic matter becomes more soluble at higher pH as dissociation increases the surface negative charge. Also, organic matter may form coordination complexes with some metals involving covalent bonds. 3.6 THE SOLID SURFACES IN SUBMERGED SOILS Many submerged soils are developed in recent in alluvium and are often young or only weakly weathered (Section 1.3). The overall composition of the clay fraction is therefore often close to that of the parent sediment. Hence the following generalizations can be made for rice soils in the humid tropical lowlands (Kyuma, 1978; Binkman, 1985) • Soils derived from marine alluvial sediments tend to be dominated by montmorillonitic 2:1 clays whereas those from riverine sediments have vermiculitic 2:1 clays with mixtures of 1:1 clays and metal oxides, the sediment being developed under more strongly weathered conditions. • Soils developed in positions higher in the landscape tend to be dominated by more-weathered material. • Soils derived from basic volcanic ejecta, metamorphic rocks and granitic rocks have corresponding mineralogies. Various changes in mineralogy are induced by seasonal flooding. The first factor in this is the change in base status of the soil due to the flow of water

70

Interchange of Solutes between Solid, Liquid and Gas Phases

through and across it. Depending on the alkalinity of the water entering and leaving, the soil may be enriched with bases or depleted. Large quantities of bases are liberated in soil reduction following flooding and may be leached. The balance will depend on the topological and hydrological situation of the soil, and in general soils low in the landscape will accumulate bases and those higher will be depleted. There may be biological fixation of bases, for example by aquatic snails in fields receiving base-rich interflow water or irrigation. For example, after 15 years of intensive irrigation of ricefields at the International Rice Research Institute (Laguna, Philippines) with base-rich water (4 mmolc L−1 Na+ and 1 mmolc L−1 Mg2+ ), there was a marked accumulation of CaCO3 in snail shells and the aerobic soil pH increased from 5.6–6.0 to 6.5–7.0 (Moormann and van Breemen, 1978). The reverse process—decalcification—may also occur. For example, in the calcareous soils of the Ganges and Megha sediments, Bangladesh, where the ricefields are rainfed and the rainwater tends to be acid, Brammer (1971) reported losses of 1 % of the CaCO3 in 25 years from sediments that originally contained 5–10 % CaCO3 . The calcite is dissolved by acids in the rain and CO2 formed during soil reduction, and Ca(HCO3 )2 is leached out of the soil. The second factor is the changing redox state of the soil. Generally iron is the most abundant redox species present. Table 3.11 shows total iron contents of a range of rice soils across Asia. From 20 to 80 % of the iron is present as free Fe(III) oxides and often from 1 to 20 %—and sometimes as much as 90 %—of the free Fe(III) is reduced to soluble and exchangeable Fe(II) following submergence (see for example Figure 2.8). Some of the exchangeable Fe(II) is subsequently reprecipitated as mixed Fe(II)Fe(III) compounds of uncertain composition. There may also be reduction of structural Fe(III) to Fe(II) within clay minerals. These changes take place over a matter of weeks. Upon subsequent drying and re-oxidation, the exchangeable and amorphous Fe(II) are rapidly converted to ferric hydroxides, initially in amorphous forms that recrystallize only very slowly (Figure 3.8). As a result, amorphous ferric hydroxides and similar materials tend to accumulate in the soil at the expense of more stable Table 3.11 Total iron contents (mg Fe g−1 ) of rice soils Country

Mean

Min.

Max.

n

Bangladesh Burma Cambodia India Indonesia Malaysia Philippines Sri Lanka Thailand Japan

40.1 39.7 32.2 72.3 20.8 20.8 54.1 37.4 25.2 42.0

8.0 5.9 0.0 9.0 1.7 1.9 28.8 5.0 0.0 —

66.1 65.5 80.1 117.6 50.3 50.3 86.7 149.6 90.0 —

53 16 16 73 44 41 54 33 80 155

Source: Kyuma (1978).

71

The Solid Surfaces in Submerged Soils Slow reduction

Crystalline Fe(III) oxides

Fast oxidation

Amorphous Fe(III) oxides

Fe(II)

Fast reduction Slow crystallization

Figure 3.8 Accumulation of amorphous Fe(III) compounds under alternating reduction and oxidation (after Moormann and van Breemen, 1978). Reproduced by permission of IRRI

ferric oxides. In turn, the amorphous materials are more easily reduced during soil flooding and over time the iron compounds reach a steady state in which easily reducible amorphous materials are combined with more recalcitrant minerals. The proportion of amorphous and crystalline materials will be influenced by the hydrological regime and climate. Concomitantly there are short- and long-term changes in soil organic matter. Changes in Surface Properties Following Submergence For the most part, the overall composition of the clay fraction in soils is determined by long-term processes and does not change rapidly with changes in conditions. However the properties of the clay surface can change rapidly and the surface is often not in equilibrium with the rest of the solid phase. Thus alternating reduction and oxidation under variable water regimes cause rapid but transient changes in surface properties, as well as more persistent changes in the overall composition of the solid phase. The changes in surface properties following soil submergence are as follows. Dissolution of Oxide Coatings. The net negative charge on the soil solid may increase following reduction as a result of dissolution of oxide and oxyhydroxide coatings on clay surfaces. The pzc of oxides and oxyhydroxide are at or above neutral pH (Table 3.10), and so the coatings on clays are positively charged in neutral and acid soils and neutralize some of the clay’s negative charge. Their dissolution therefore results in an increase in the net negative charge on the surface. Hence, for example, an oxide with composition Fe(OH)2.5 0.5+ , is reduced according to the half reaction [soil—2Fe(OH)2.5 ] + 5H+ + 2e− −−−→ [soil—]− + 2Fe2+ + 5H2 O

(3.42)

where e− represents an electron transferred in the reduction. If the corresponding oxidation of soil organic matter is (Chapter 4) CH2 O + H2 O −−−→ CO2 + 4H+ + 4e−

(3.43)

72

Interchange of Solutes between Solid, Liquid and Gas Phases

then the overall reaction is 2[soil—2Fe(OH)2.5 ] + CH2 O + 6H+ −−−→ 2[soil—]− + 4Fe2+ + CO2 + 9H2 O (3.44) In Reaction (3.44), for each mol of Fe reduced the surface negative charge increases by 0.5 molc and 1.5 mol of H+ are consumed. Roth et al. (1969) found increases in surface negative charge equivalent to 10–60 % of the initial charge for a range of soils and soil clays. The change could be attributed quantitatively to the removal of the positively charged oxide coatings and was reversed by re-oxidizing the samples. Changes in charge with reductive dissolution of oxides have been demonstrated using chemical reducing agents (Roth et al., 1969) and microbial reducing agents (Bloomfield, 1951; Ottow, 1973; Lovley, 1991), and under field conditions (Favre et al., 2002). Dissolution and reduction of crystalline Fe(III) minerals is accelerated by chelation with carboxylate ligands in the presence of Fe(II) (Zinder et al., 1986; Blesa et al., 1987; Phillips et al., 1993; Kostka and Luther, 1994). Therefore as soil reduction proceeds and carboxylates formed in oxidation of organic matter accumulate in solution together with Fe2+ , dissolution and reduction of crystalline Fe(III) will commence. Dissolution of oxyhydroxide coatings will therefore lag behind the initial reduction of Fe(III). Reduction of Structural Fe. There may also be changes in charge due to reduction of structural Fe(III). Virtually all soil clay minerals contain some iron in their crystal structures and reduction of this structural Fe by chemical or microbial reducing agents, with the iron remaining octahedrally coordinated in the clay structure, is well documented (Stucki, 1988; Stucki et al., 1997). The extent of reduction, whether by microbes or chemical reducting agents, can be as much as 90 % of the octahedral Fe(III) in a few days (Kostka et al., 1999). The rate is enhanced by the presence of organic chelating agents that commonly occur in sediments and flooded soil solutions, and under such conditions Fe(III) reduction may lead to partial dissolution of the clay (Kostka et al., 1999). As structural Fe(III) is reduced, the negative charge on the clay will increase. It is found experimentally that the increase in negative charge is not directly equivalent to the amount of Fe(III) reduced, and the more reduced the clay is the smaller is the change in charge. An example is shown in Figure 3.9. The mechanism behind this is uncertain but involves dehydroxylation of the clay structure during reduction and sorption of metal cations from the solution (Stucki et al., 1997; Drits and Manceau, 2000). The extent of dehydroxylation and sorption varies with the extent of reduction, hence the change is nonlinearly related to the amount of Fe reduced. For example, for a nontronite: M[Si7 Al]Fe(III)4 O20 (OH)4 + mM+ + nH+ + pe− −−−→ M1+m [Si7 Al]Fe(III)4−p Fe(III)p O20 (OH)4−n + nH2 O

(3.45)

73

The Solid Surfaces in Submerged Soils

Surface charge (mmolc g−1)

1.5 1.4 1.3 1.2 1.1 1.0 0.9 0.8 0.7 0.0

0.5

1.0

1.5

2.0

Amount of Fe(III) reduced (mmol g−1)

Figure 3.9 Relation between surface charge and reduction of structural Fe in a dioctahedral smectite. Points are experimental data; lines are theoretical relations discussed in the text (Drits and Manceau, 2000). Reproduced by permission of Clay Minerals Society

where M is a sorbed cation and m, n and p are coefficients. The solid line in Figure 3.9 shows the fit to Equation (3.45) and the dotted line shows the expected relationship if only dehydroxylation occurs. If the generic reaction is simplified to [clay—Fe(III)OH] + nH+ + e− −−−→ [clay—Fe(II)OH1−n ](1−n)− + nH2 O (3.46) then using the upper value n = 0.75 (Figure 3.9), the complete reaction with simultaneous oxidation of organic matter (as for Reaction 3.44) is: 4[clay—Fe(III)OH] + CH2 O −−−→ 4[clay–Fe(II)OH0.25 ]0.25− + H+ + CO2 + H2 O

(3.47)

In Reaction (3.47), for each mol of Fe reduced the surface negative charge increases by 0.25 molc and 0.25 mol of H+ are released. For moderate reduction the changes are completely reversible but they are progressively less so with more thorough reduction (Stucki et al., 1984; Komadel et al., 1995; Gates et al., 1996). There are concomitant changes in the clay’s physical and chemical properties, including its surface area, swelling behaviour, and capacity to sorb cations. Changes in pH-dependent Charge. Changes in pH with soil reduction will cause changes in the charges on inorganic –OH functional groups and organic matter. From Equation (3.41), the increase due to organic functional groups will be approximately 0.5 mmolc g−1 organic C per unit pH increase, or, for a soil with 1 % organic C, 5 mmolc kg−1 soil. This is small compared with the changes due to dissolution of oxide coatings and reduction of structural Fe, which are of the order of several tens of mmolc kg−1 soil. But it may be important in highly

74

Interchange of Solutes between Solid, Liquid and Gas Phases

weathered soils with low cation exchange capacity. The changes due to edge –OH groups on kaolinites will be of a similar magnitude. The increase in oxide charge per unit increase in pH will be of the order of 5 mmolc kg−1 oxide at the ionic strength and pH of typical flooded soil solutions. However, if the surface oxide coatings dissolve in the process of reduction, this will be of no consequence. Formation of New Solid Phases. Once a sufficient concentration of dissolved constituents has been reached following submergence and soil reduction, new solid phases will precipitate. The nature of these compounds is discussed in detail in Chapter 4. In neutral soils with smectite or vermiculite in the clay fractions, the changes in redox and bases status following soil flooding may cause synthesis of materials similar to smectite with Fe2+ in the octahedral sheet. Other cations, e.g. Zn2+ , may also become entrapped. In acids soils, particularly those with kaolinite clay minerals, soluble Fe2+ concentrations tend to rise to high levels because of low CEC and because conditions do not favour precipitation of Fe(II) oxides or carbonates or synthesis of silicates. When a reduced soil is re-oxidized, Fe2+ changes into Fe(OH)3 . The original Fe oxides are thus distributed differently, generally with a higher specific surface and activity. In high-activity clay soils, this may increase the stability of the structure established just before flooding. In low-activity clay soils the effects of alternate reduction and oxidation are less clearly beneficial, partly because of leaching of nutrients.

3.6.1 ORGANIC MATTER IN SUBMERGED SOILS In general the organic matter in soils tends towards a steady state in which additions from net primary production balance the various processes removing it, and the organic matter attains a level and composition characteristic of the prevailing conditions. As discussed in Chapter 1, net primary productivity of wetlands is in general far greater than that of uplands and rates of organic matter loss due to decomposition, run-off, leaching and erosion tend to be less. In general, the yield of energy per unit of carbon oxidized is smaller in anaerobic fermentation than in aerobic respiration, and so, other things being equal, rates of decomposition are less. Hence in the temperate zone wetland soils often have large organic matter contents or are peaty. However, in the tropics, peat bogs and fens are rare (Table 1.1) and wetland rice soils tend not to have particularly large organic matter contents. In data assembled by Greenland (1997), the mean level of organic carbon in the topsoils of wetland rice soils from across tropical Asia was 2 %, and after excluding acid peaty soils the mean was 1 %. This compares with a range of 1.27–1.81 % for Oxisols and Ultisols of the Cerrado region of Brazil (Sanchez,

75

The Solid Surfaces in Submerged Soils

1981) and 2.78–4.80 for Oxisols and Ultisols of the humid tropical forest zone of Sumatra (van Noordwijk et al., 1997). Under intensive multiple-rice cropping, the soil organic mater may increase until a new steady state level is reached after some years. Table 3.12 gives an example for a rice–rice system over five crops in 2 years compared with a maize–rice system. However, very large soil organic matter contents do not develop. Evidently rates of decomposition are greater than expected for simple anaerobic decomposition. Figure 3.10 shows comparable rates of organic matter decomposition in soils that were kept continuously flooded or well-drained under otherwise similar, tropical conditions for 3–4 years (Neue and Scharpenseel, 1987). Clearly, Table 3.12 Carbon balances for rice–rice and maize–rice cropping systems over five consecutive crops in 3 years at IRRI, Laguna, Philippines. Except for short stubble all straw was removed from the fields and no organic manures were applied. Values are t C ha−1 ± SEs Cropping system

Rice–rice

N fertilizer (kg ha−1 ) a

Initial SOC Final SOC Change in SOC (% change) C from crop residues C mineralized (% of crop residues)

Maize–rice

0–0

190–100

0–0

190–100

19.13 ± 0.83 20.97 ± 0.49 +1.84 ± 0.44 (+10) 4.93 ± 0.18 3.09 ± 0.56 (63)

19.41 ± 0.27 22.15 ± 0.54 +2.74 ± 0.67 (+14) 7.72 ± 0.21 4.98 ± 0.47 (64)

19.22 ± 0.79 19.01 ± 0.40 −0.22 ± 0.50 (−1) 4.15 ± 0.28 4.37 ± 0.56 (105)

19.38 ± 0.97 19.83 ± 0.43 +0.46 ± 0.96 (+2) 7.09 ± 0.18 6.63 ± 0.87 (94)

a

SOC, soil organic carbon. Source: Witt et al. (2000). Reproduced with kind permission of Kluwer Academic Publishers.

(a) well drained

(b) submerged 100 Tropept Tropept Humult

30

10

Aquept Aquoll Aquult

30

10

14C

remaining (%)

100

3

0

10

20

3 0 10 20 30 40 50 Time (months after straw incorporation)

30

40

50

Figure 3.10 Decomposition of 14 C-labelled rice straw in (a) well-drained upland soils and (b) continuously submerged lowland soils under tropical conditions (adapted from Neue and Scharpenseel, 1987). Reprinted with permission from Elsevier

76

Interchange of Solutes between Solid, Liquid and Gas Phases

other factors are at work. In addition to temperature and aeration, and the quantity and nature of organic matter inputs, other factors influencing decomposition include the soil pH, which may be more favourable following submergence; the level and balance of nutrients, which may also be more favourable; and the communities of micro- and macro-fauna that together bring about the decomposition. Submerged soils are never wholly anoxic and contain aerated zones at the interfaces between the soil and floodwater and the soil and plant roots. Burrowing oligochaete worms transport fresh and partially decomposed organic matter between the soil and floodwater, and the activities of organisms in the soil and floodwater are thereby linked. Hence the soil–floodwater system as a whole behaves quite differently from its component parts in isolation. Decomposition processes are discussed further in Chapter 5. Studies by Olk and others of long-term (≤ 30 years) changes in organic matter in soils under intensive wetland rice cultivation have shown a gradual accumulation of less humified and more phenolic material (Olk et al., 1996, 1998, 1999; Mahieu et al., 2000a, b, 2002). These authors compared the chemical composition of organic matter from four soils with different histories of cropping and submergence: (a) one crop of upland rice annually without soil submergence; (b) one wetland rice crop and one soybean crop annually; (c) two wetland rice crops annually; and (d) three wetland rice crops annually. With increasing intensity of wetland rice cropping there were large increases in the proportions of less humified material in the soil organic matter, measured as diester P, amide N and phenolic C in nuclear magnetic resonance spectra. There were also positive correlations with visible light absorption and concentrations of free radicals, both of which are indices of humification, and negative correlations with the concentration of H, a negative index of humification. The authors speculate that slower lignin decomposition under restricted O2 supply in submerged soil leads to incorporation of phenolic compounds into young soil organic matter as it is turned over. Since phenolic compounds can react strongly with nitrogenous compounds, they further speculate that the mineralization and immobilization of N in intensively cropped rice soils may be adversely affected by accumulation of phenolic material.

3.7 SOLID–SOLUTION INTERACTIONS 3.7.1 ADSORPTION Adsorption depends on the interaction of ions and uncharged solutes with the functional groups on soil surfaces. It is in some ways analogous to complexation of ions with ligands in solution (Section 3.1), with the difference that the surface functional groups are stationary and their properties depend to a greater extent on interactions with neighbouring groups. Two types of surface complex are distinguished:

77

Solid–Solution Interactions

(a) inner-sphere complexes in which the adsorbed species is bound directly to the surface functional group, with no intervening water molecules; and (b) outer-sphere complexes in which at least one water molecule remains between the absorbed species and the surface. In inner-sphere complexes the bonding is covalent or ionic and the reactivity of the surface is altered by the interaction; in outer-sphere complexes the bonding is largely electrostatic and the reactivity of the surface is largely unaltered. Innersphere complexation is usually slower than outer-sphere and is often not readily reversible. It can occur regardless of the net surface charge and is little influenced by the ionic strength of the external solution. If an ion is adsorbed without forming a surface complex, neutralizing surface charge in only a delocalized way, then it is said to be part of the diffuse ion swarm. Such ions are free to move about in the soil solution near the surface. Figure 3.11 shows surface complexes on an inorganic hydroxyl surface. It shows the distinction between inner- and outersphere complexes, depending on the presence of water molecules between the surface and complexed species. The region of the diffuse ion swarm begins at the outer edge of the water molecules solvating ions in outer-sphere complexes. There are a number of more loosely defined terms for different types of adsorption that are related to the form of surface complexation. Specifically adsorbed ions are held in inner-sphere complexes whereas non-specifically adsorbed ions are in outer-sphere complexes or the diffuse-ion swarm. Readily exchangeable Oxygen Metal Inner-sphere complexes

H X

e.g. M = Zn, Pb, Cd X = P, As, Si M

H+ Outer-sphere complexes e.g. M = K, Ca, Mg, Fe X = Cl, NO3, HCO3

X−

H Water molecules

H M+ Solid−water interface

Figure 3.11 Complexes formed between solutes and hydroxyl groups on oxides and layer silicate edges (adapted from Sposito, 1984b). Reproduced by permission of Oxford University Press

78

Interchange of Solutes between Solid, Liquid and Gas Phases

ions are those that can be replaced easily by leaching with an electrolyte solution. This is an empirical definition, but only fully solvated ions can be readily exchangeable and therefore must be either in the diffuse-ion swarm or in outersphere complexes. Adsorption interacts strongly with complexation in solution. Table 3.13 indicates the range of complexes between metal ions and inorganic and organic ligands in soil solutions. In a submerged soil the organic ligands present include acetate, formate and propionate at concentrations of 10–40 mM in the early stages following submergence though less than 1 mM after 3–4 weeks. In addition concentrations of amino acids, phenolic acids and larger molecular weight humic acids may reach a few hundred µM, though transiently. Figure 3.12 shows the calculated effects of realistic concentrations of acetate, formate, propionate, glutamate, glycine, benzoate and phenylacetate on Fe(II), Mn(II) and Zn(II) species. The figure shows that for Fe(II) and Mn(II) the free ion dominates at all pHs, except for Fe above pH 9 where hydroxy complexes are important. Complexes with acetate are also significant at pHs above about 5, and FeHCO3 + above pH 6 and MnGlu above pH 5. Complexes with formate, propionate or either of the phenolic acids are unimportant at all pHs. The picture is more complicated for Zn(II) with many more significant species. The free ion dominates at pH ≤ 7.5 but complexes with acetate, HCO3 − , glutamate and especially CO3 2− are important at various pHs. Hydroxy complexes are only important at pH>9. Figure 3.13 shows the solubility of Zn2+ in soil at four Zn levels and different pHs. The figure shows that the soil solution is under-saturated with respect to likely pure Zn precipitates up to high pHs, and there is a marked minimum in solubility at near neutral pH. The explanation involves cation exchange and specific adsorption reactions, trace amounts of Zn2+ being sorbed preferentially over the main exchanging cations, and complexation reactions between Zn2+ and organic ligands in solution. The negative charge on soil surfaces increases as the pH increases, tending to increase sorption of Zn2+ on variable-charge surfaces. But at near neutral pH the concentration of dissolved organic matter in solution Table 3.13 The main species of trace metals in soil solutions Metal Mn(II)

Acid soils

Alkaline soils

Mn2+ , MnSO4 0 , Orga 0

2+

Mn2+ , MnSO4 0 , MnCO3 0 , MnHCO3 + +

Fe(II)

Fe , FeSO4 , FeH2 PO4

Ni(II)

Ni2+ , NiSO4 0 , NiHCO3 + , Org 2+

Cu(II)

Org, Cu

Zn(II)

Zn2+ , ZnSO4 2+

FeCO3 0 , Fe2+ , FeHCO3 + , FeSO4 0 NiHCO3 0 , NiHCO3 + , Ni2+ CuCO3 0 , Org ZnHCO3 + , ZnCO3 0 , Zn2+ , ZnSO4

0

+

Cd(II)

Cd , CdSO4 , CdCl

Pb(II)

Pb2+ , Org, PbSO4 0 , PbHCO3 +

Cd2+ , CdCl+ , CdSO4 0 , CdHCO3 + PbCO3 0 , PbHCO3 + , Pb(CO3 )2 2− , PbOH+

a Org, organic complexes, e.g. with fulvic acids. Source: adapted from Sposito (1983). Reproduced by permission of Elsevier.

79

Solid–Solution Interactions 10−2 Fe2+ 10−3

FeAc

10−4

FeHCO3+

Fe(OH)3− +

FeOH

10−5

Fe(OH)2 FeGly

10−6

Concentration (M)

10−3

Mn2+ MnGlu

10−4

MnAc

10−5 MnHCO3+ MnOH+

10−6 10−7 10−7 Zn2+ ZnAc

−8

10

ZnCO3

ZnHCO3+

ZnProp ZnForm 10−9 ZnGlu

10−10

ZnOH+

ZnGly

Zn(OH)2 Zn(OH)3−

−11

10

4

5

6

7 pH

8

9

10

Figure 3.12 Distributions of Fe(II), Mn(II) and Zn(II) species in a simulated flooded soil solution. Total concentrations in mM are Fe(II) = 2.5, Mn(II) = 0.25, Zn(II) = 0.0001, CT = 20, acetate = 10, formate = 1, propionate = 1, glutamate = 0.1, glycine = 0.1, benzoate = 0.1, phenylacetate = 0.1. Species accounting for 0 then (Ox) > (Red) and the oxidant is the dominant species, and vice versa. Hence a plot of pe versus pH with (Ox) = (Red) has slope m/n and intercept pe, and for points above the line the oxidant is dominant and for points below the reductant is dominant. pe is taken as the dependent variable, plotted on the ordinate, because pH is often controlled by processes in addition to redox reactions and is therefore more properly the independent variable. Figure 4.1 gives examples for biological redox couples important in natural systems. Figure 4.1(a) shows the diagram for the H2 O–O2 and H2 O–H2 couples. The respective lines are pe = 20.75 − 14 log PO2 − pH with PO2 = 1 atm, and pe = − 14 log PH2 − pH with PH2 = 1 atm. These are upper limits for the partial pressures of O2 and H2 in natural waters. For points above the upper line, H2 O is an effective reductant, producing O2 ; for points below the lower line H2 O is an effective oxidant, producing H2 . The region between the lines, where O2 acts as an oxidant and H2 as a reductant, covers most circumstances in natural systems.

−10

−5

0

5

10

15

4 5 6 7 8 9 10

PH2 > 1 atm

H2O

PO2 > 1 atm

(b)

NH3

4 5 6 7 8 9 10

NH4+

N2

NO3−

(c)

NO2−

4 5 6 7 8 9 10 pH

NH4+

NH3

NO3−

(d)

HS−

4 5 6 7 8 9 10 11

H2S

S(s)

SO42−

(e)

CO32−

HCO3−

4 5 6 7 8 9 10 11

CH4

CO2

Figure 4.1 pe–pH diagrams for important biological redox couples in natural systems. (a) H2 O–O2 . (b) The nitrogen system considering only stable equilibria: the only oxidation states involved are (−III), the elemental state and (+V). (c) The nitrogen system treating NH4 + , NH3 , NO3 − and NO2 − as metastable with respect to N2 which is treated as redox-inert. (d) The SO4 2− –S(s)–H2 S(aq) system, [total soluble S] = 10−2 M. (e) The carbon system ignoring elemental C. After Stumm and Morgan (1996). Reproduced by permission of Wiley, New York

pe

(a) 20

100

101

Thermodynamics and Kinetics of Redox Reactions

Figure 4.1(b) shows the diagram for the stable equilibria in the nitrogen system. According to this diagram N2 should be largely oxidized to NO3 − in most natural waters. The fact that it is not and N2 is known to persist in oxic waters indicates that a complete redox equilibrium does not exist; only a partial equilibrium is attained under the mediation of microbes. Figure 4.1(c) shows the diagram for the nitrogen system with N2 treated as redox-inert and NH4 + , NH3 , NO3 − and NO2 − as metastable with respect to N2 . This diagram more correctly represents conditions in natural systems, with NH4 + as the stable species under mildly reducing conditions and NO3 − under oxic conditions. This example illustrates the difficulty in choosing the correct redox couples to represent real systems in pe–pH diagrams. Some independent insight into the system is generally required to choose the correct couples. Figure 4.1(d) and (e) show diagrams for sulfur and carbon systems. The situation is further complicated for redox reactions involving several solid phases. An example is the Fe–CO2 –H2 O system shown in Figure 4.2. This shows that ferrihydrite, Fe(OH)3 , can be formed over a wide range of pe and pH, though the pe range is increasingly restricted under increasingly acid conditions, and Fe2+ is then the stable form. Siderite, FeCO3 , and the hypothetical Fe(II) hydrous oxide Fe(OH)2 may be formed under moderately reducing conditions but only at pH > 7. Elemental Fe is only stable under very strongly reducing conditions, outside the range in which water is stable. In real systems the situation

25 FeOH2+

20 Fe3+

PO2 > 1 atm

15

Fe(OH)4−

pe

10 Fe2+

5

Fe(OH)3(amorph, s)

0

PH2 > 1

−5 −10 −15

FeCO3(s) Fe(OH)2(s)

Fe(s) 0

2

4

6

8 pH

10

12

14

Figure 4.2 pe–pH diagram for the Fe–CO2 –H2 O system. [Fe(II)] = 1 mM, [Fe(III)] = 0.01 mM, CT = 5 mM. Amorphous Fe(OH)3 is ferrihydrite, FeCO3 siderite, and Fe(OH)2 a hypothetical Fe(II) hydrous oxide. The details of the construction of this diagram are explained in Stumm and Morgan (1996, Chapter 8). Reproduced by permission of Wiley, New York

102

Reduction and Oxidation

may be complicated by the presence of other redox species with which Fe reacts, such as sulfide, and by the slow kinetics of redox and precipitation reactions and the need for microbial mediation. Thus for example siderite is rarely found in soils though in many cases it is the thermodynamically favoured phase as shown in Figure 4.2. These points are discussed further in later sections. Because of the sensitivity of pe to pH it is often convenient to compare peo values ‘corrected’ to pH 7 and termed peo∗ , where: peo∗ = peo −

(Red) m 1 log −7 n (Ox) n

(4.22)

As discussed earlier, the concentration-dependent term in Equation (4.22) will often be small in comparison with the pH term and can be ignored. For couples in which the concentration term is more important, such as Fe(OH)3 –Fe2+ , peo∗ values can be calculated for representative concentrations. Table 4.2 gives peo∗ values for important redox couples in natural systems arranged in order of decreasing peo∗ with strong oxidants at the top and strong reductants at the bottom. From such a table it is possible to infer which couples will react when present together and which will have the oxidizing role and which the reducing role. The table shows for example that Fe(OH)3 can readily oxidize organic matter ‘CH2 O’ to form CO2 and Fe2+ but it cannot oxidize N2 to NO3 − . However note that the peo∗ value for the Fe(OH)3 –Fe2+ couple is sensitive to the value of (Fe2+ ).

4.1.5 ENERGETICS OF REACTIONS MEDIATED BY MICROBES Most redox reactions in vitro reach equilibrium only extremely slowly with half times of the order of months or years, even though they may be highly favoured thermodynamically. This is illustrated by the persistence of N2 in oxic systems even though its oxidation to NO3 − is strongly favoured (Table 4.1). However, microbes in soil and water are capable of catalysing particular reactions from which they obtain energy for metabolism. The half times of such microbially catalysed reactions are of the order of hours or days. The amounts of energy consumed or produced in redox reactions, and hence the efficiency with which they can be exploited by microbes, can be calculated from thermodynamic data. This gives surprisingly good insights into the dynamics of microbial communities in natural systems without detailed knowledge of the biochemical and physiological pathways involved. For example, the sequence of reduction reactions that occur in submerged soils following exclusion of O2 matches the order of decreasing free energy change for the corresponding redox reactions. Note that organisms cannot carry out gross reactions that are thermodynamically impossible: they do not oxidize substrates or reduce oxidants per se, but merely catalyse the process by mediating the electron transfers occurring. The energy produced or consumed in a given redox reaction is calculated as follows.

103

Thermodynamics and Kinetics of Redox Reactions Table 4.2

Equilibrium constants of reduction half-reactions at pH 7 and 25 ◦ C

1 O (aq) + H+ + e− 4 2 1 NO3 − + 65 H+ + e− 5

= 12 H2 O

1 NO3 − + 54 H+ + e− 4 1 MnO2 (s) + 2H+ + e− 2 1 Mn3 O4 (s) + 4H+ + e− 2 + −

=

=

MnOOH(s) + 3H + e 1 NO3 − 2 1 NO3 − 8 1 NO2 − 6 1 CH2 O 4

+

+H +e + +

5 + H 4 4 + H 3 +

−

30.79

8.33

25.33

8.02a

14.15

7.15

14.90

6.15

15.14

5.82

6.94

−0.06

16.54

−1.46a

3.99

−3.01

5.25

−3.50

4.25

−3.63

2.87

−4.13

4.63

−4.70

11.31

−6.69a

= =

1 NO2 − 2 1 NH4 + 8 1 NH4 + 6

+ + +

1 CH4 (g) 4 2+

Fe(OH)3 (s) + 3H + e

= Fe

1 CH2 O + H+ + e− 2 1 SO4 2− + 45 H+ + e− 8

= 12 CH3 OH

1 SO4 2− + 89 H+ + e− 8 1 CO2 (g) + H+ + e− 8

=

1 N 6 2

= +

=

=

α-FeOOH(s) + 3H + e

−

1 CO2 (g) 4 1 CO2 (g) 4

1 H O 2 2 3 H O 8 2 1 H O 3 2

+

1 H O 4 2

+ 3H2 O

1 H S(g) + 21 H2 O 8 2 1 HS− + 21 H2 O 8 1 CH4 (g) + 41 H2 O 8 1 NH4 + 3 2+

= Fe

+ 2H2 O

= 21 H2 (g)

H+ + e − +

12.65 10.06

= Mn2+ + 2H2 O =

+ 43 H+ + e−

21.05 18.81

= 32 Mn2+ + 2H2 O

+ e− −

13.75

9.67a

+e

+

20.75

21.82

=

+H +e

peo∗

= 12 Mn2+ + 2H2 O

−

−

1 N (g) + 53 H2 O 10 2 1 N O(g) + 85 H2 O 8 2

peo

+ H + e−

=

+ H+ + e−

=

1 (glucose) + 41 H2 O 24 1 CH2 O + 21 H2 O 4

0.00

−7.00

−0.20

−7.20

−1.20

−8.20

For reductive dissolution of Mn and Fe oxides peo∗ values are calculated with (Mn2+ ) = 0.2 mM and (Fe2+ ) = 1 mM to represent conditions in submerged soil solutions; in other couples reactants are given unit activities.

a

Consider the reaction Ox1 + Red2 = Red1 + Ox2 for which the reduction half reactions are: Ox1 + ne− = Red1 Ox2 + ne− = Red2 with equilibrium constants K1 and K2 , i.e. pe1 = 1/n log K1 and pe2 = 1/n log K2 . Therefore
Go = −2.303RT log K = −2.303RT log

K1 = −2.303RT n(peo1 − peo2 ) K2 (4.23)

104

Reduction and Oxidation

The free energy change for the reaction is
G =
Go + 2.303RT log

(Ox2 )(Red1 ) (Ox1 )(Red2 )

Combining Equations (4.23) and (4.24) gives
 (Ox2 )(Red1 ) o o
G = −2.303RT n(pe1 − pe2 ) − log (Ox1 )(Red2 )

(4.24)

(4.25)

As discussed earlier, the dependence of pe on the concentrations of reductants and oxidants is often small in comparison with its dependence on pH. The term in the square brackets in Equation (4.25) can therefore be replaced by peo∗ terms giving for the approximate standard free energy change: o∗
Go∗ ≈ −2.303RT n(peo∗ 1 − pe2 )

(4.26)

Figure 4.3 shows oxidation and reduction reactions used by microbes as energy sources on a scale of peo∗ (data from Table 4.1). The free energy changes for the different complete reactions can be read from the
Go∗ scale, in accordance with Equation (4.26). The energy expended by microbes in elaborating carbon, for example through fixation of CO2 , and other elements, for example nitrogen through fixation of atmospheric N2 , can be calculated in a similar way. Such calculations indicate the maximum energy available from a reaction or the minimum required to carry it out. The true gain to a microbe is smaller, or the cost larger, because of the energy required for cell maintenance and reproduction and other processes. The energetic efficiencies of biochemical processes are typically of the order of 30–40 %. Nonetheless the ecological succession of microbes in response to the stepwise oxidation of reduced compounds and exhaustion of oxidants can be predicted from such calculations. Thus the succession of aerobic organisms, denitrifiers, manganese reducers, iron reducers, sulfate reducers and methanogenic bacteria following submergence of a soil directly matches the order of decreasing peo∗ for the corresponding redox couples in Figure 4.3(b): O2 –H2 O, NO3 − –N2 , MnO2 (s)–Mn2+ , Fe(OH)3 (s)–Fe2+ , SO4 2− –HS− and CH2 O–CH4 . Microorganisms and organisms in general can be classified according to the principal sources of their energy, carbon and electrons. A hierarchical classification is not possible because all combinations of these three occur. Thus all three can be separate, as for green plants which obtain their energy from sunlight, carbon from CO2 and electrons by oxidizing water to O2 ; and all three can be the same, as for the majority of bacteria which use organic compounds as their sources of energy, carbon and electrons. Organisms that obtain their carbon from inorganic compounds, mainly CO2 , are called autotrophs. They are subdivided into photoautotrophs which obtain energy from sunlight—for example, green plants and photosynthetic bacteria—and chemoautotrophs which obtain energy from chemical processes—for example, in Figure 4.3(a), nitrifying bacteria, which oxidize NH4 + to NO3 − , sulfur oxidizing

130

120

H2 H+

H2S SO42−

100

110

Fe2+ Fe(OH)3

NH4 NO2

−

+

NO2

NO3−

−

90

80

70

60

50

40

30

20

10

Oxidations 15 14 13 12 11 10 9 8 7 6 5 4 3 2 1 0 −1 −2 −3 −4 −5 −6 −7 −8 −9 −10 O2 H2O

Reductions (b)

CH2O CO2

Oxidations 15 14 13 12 11 10 9 8 7 6 5 4 3 2 1 0 −1 −2 −3 −4 −5 −6 −7 −8 −9 −10 H+ H2

CO2 CH4

SO42− HS−

CH2O CH3OH

Fe(OH)3 Fe2+

NO3 NO2− NO3− NH4+

−

MnO2 Mn2+

NO3− N2O

NO3− N2

O2 H2O

Reductions

10

20

30

40

50

60

70

80

90

100

110

120

130

−∆Go* (kJ mol−1)

peo*

peo*

Figure 4.3 Free energy changes in redox reactions mediated by microbes. (a) Oxidation of reduced inorganic compounds linked to reduction of O2 . (b) Oxidation of organic matter ‘CH2 O’ linked to reduction of various organic and inorganic oxidants. pH = 7 and unit oxidant and reductant activities except (Mn2+ ) = 0.2 mM and (Fe2+ ) = 1 mM

−∆Go* (kJ mol−1)

(a)

105

106

Reduction and Oxidation

bacteria, which oxidize reduced S compounds to SO4 2− , and Mn(II) and Fe(II) oxidizing bacteria, which produce insoluble Mn and Fe oxides, though it is not certain that useful energy is derived from this. Examples of autotrophs using different electron sources in fixing CO2 are green plants which derive electrons from the oxidation of water, sulfide oxidizers which oxidize H2 S(g) to colloidal S, and ammonium oxidizers which oxidize NH4 + to NO2 − . Organisms that obtain carbon from ingested organic compounds are called heterotrophs, and most also derive their energy and electrons from these organic compounds. Examples are fungi, protozoa and most bacteria. A wide range of organic and inorganic oxidants are used as end electron acceptors in oxidizing the organic compounds, as in the reactions shown in Figure 4.3(b). Also a wide range of organic compounds are oxidized. The resulting free energy changes may differ substantially from those in Figure 4.3(b) for oxidation of the average compound ‘CH2 O’. For example the oxidation of glucose yields about 54 kJ more energy per mole of C than oxidation of acetate. This makes an increasingly significant difference the lower the peo∗ of the oxidizing couple. Thus it makes only a small difference for O2 or NO3 − reduction (12 and 15 %, respectively), but a large difference for SO4 2− reduction (69 %). An important component of the overall efficiency of energy production by microbes is the location of the linked couples and the resulting need to transport reactants and products across cell membranes. In denitrification and SO4 2− reduction, because all of the NO3 − and to a lesser extent the SO4 2− are dissolved in the soil solution, they are readily imported into the cell and their reduction linked directly to the oxidation of organic compounds via electron transfer systems. But in Mn and Fe reduction, the oxides are only sparingly soluble, and so the concentrations of Mn(III, IV) and Fe(III) in solution are small, even when large concentrations of the solid oxides are present. This presented a problem in establishing that Mn and Fe reduction was directly linked to microbial respiration in natural systems, rather than being an indirect effect through abiotic reactions involving side products of respiration. The evidence for the direct involvement of microbes is discussed in Chapter 5. 4.2 REDOX CONDITIONS IN SOILS This topic has a long history of research (Harrison and Aiyer, 1920; Sturgis, 1936; Pearsall and Mortimer, 1939; Shioiri, 1943; De Gee, 1950; Takai, 1952; Ponnamperuma, 1955; Baas-Becking et al., 1960; Jeffrey, 1961; Patrick, 1966; Ponnamperuma, 1972; Yu, 1985; Kyuma, 2003). The following factors result in conditions differing from those in simple aquatic systems: • The soil has a structure and contains a network of pores filled to a varying extent with water, and the soil is overlain by a layer of standing water of varying depth and degree of oxygenation. The filling and emptying of the pores is often very dynamic changing from complete saturation to near emptiness

Redox Conditions in Soils

•

• •

•

107

and vice versa within a matter of days. Redox conditions are correspondingly dynamic. Transport of solutes and gases through the soil is much slower than through soilfree water because of the restricted cross-sectional area for transport through the soil pore network and because of adsorption and reaction on soil surfaces (Chapter 2). Redox conditions are therefore closely linked to transport processes. Mineral surfaces have a much greater influence through sorption and precipitation of solutes and direct mediation of redox reactions. The soil contains organic matter which is humified to a varying extent and inputs of fresh organic matter are often much larger than in aquatic systems because of greater net primary productivity. The organic matter both provides substrates for microbial processes and participates in sorption and other reactions. The micro- and macrobiological ecologies are different.

In this section the redox conditions developing in soils following submergence are described and the processes governing these conditions are analysed in terms of the soil chemistry and microbiology discussed so far. 4.2.1 CHANGES WITH DEPTH IN THE SOIL The floodwater standing on the soil surface is usually sufficiently shallow, well mixed by wind and thermal gradients, and oxygenated by photosynthetic organisms that it is essentially aerobic. However transport of O2 into the underlying soil is too slow for more than a thin layer to be aerobic. In this layer the concentrations of Mn2+ , Fe2+ and other reduced species are negligible, and CO2 is the main end product of microbial respiration. In the underlying anaerobic soil, only a few millimetres away, the concentrations of Mn2+ , Fe2+ and the various organic products of anaerobic respiration can be very large. Thus conditions change dramatically over a very short distance. The distribution of reduced species with depth follows a characteristic pattern reflecting the succession of terminal electron acceptors used by microbes—O2 , NO3 − , Mn(III, IV), Fe(III), SO4 2− and organic C in fermentation reactions. Sorption, precipitation and dissolution reactions also influence the distribution. Figure 4.4 shows profiles of EH and extractable Mn2+ , Fe2+ and S2− in soil columns following flooding, illustrating some of these effects. A steady state develops over time and the profiles of the reduced species in the soil then reflect the profile of EH being progressively deeper for the less-easily reduced species. Thus the profile of Mn2+ in the figure extends closest to the surface, Mn reduction taking place at the highest EH , and the S2− profile extends least close to the surface. The depth to which O2 penetrates the soil, as indicated by the depth at which EH begins to decrease, is about 10 mm under the conditions of the figure. This depth depends on such factors as the oxygenation of the floodwater, the

25 −100 0

20

15

10

5

13 wk

4 wk

1 day

100 200 300 400 500 0 EH (mV)

10 wk

2 wk

2

10

1 wk

4 6 8 [Mn2+] (µmol g−1)

1 day

13 wk

8 wk

6 wk

10 wk

2 wk

0

4 wk

6 wk

10 20 30 [Fe2+] (µmol g−1)

1 day

1 wk

2 wk

13 wk

40

8 wk

7 wk

10 wk

0

3 wk

8 wk 11 wk

250 500 750 100012501500 [S2−] (cpm g−1)

2 wk

13 wk

5 wk

Figure 4.4 Depth distribution of EH , Mn2+ , Fe2+ , S2− at different times following flooding (Patrick and DeLaune, 1972). Reproduced by permission of Soil Sci. Soc. Am.

Depth in soil (mm)

0

108

109

Redox Conditions in Soils

amount and quality of organic matter present, and the concentrations of easily reducible Fe(III) and other reductants. I now summarize the changes in electrochemical conditions that occur in the soil with time following submergence as they affect the profiles and dynamics of redox species. 4.2.2 CHANGES WITH TIME Reduction of a submerged soil proceeds roughly in the sequence predicted by thermodynamics: −

O2 + CH2 O −−−→ CO2 + H2 O +

4NO3 + 5CH2 O + 4H −−−→ 2N2 + 5CO2 + 7H2 O

2MnO2 + CH2 O + 4H+ −−−→ 2Mn2+ + CO2 + 3H2 O +

2+

4Fe(OH)3 + CH2 O + 8H −−−→ 4Fe

+ CO2 + 11H2 O

(4.27) (4.28) (4.29) (4.30)

SO4 2− + 2CH2 O + 2H+ −−−→ H2 S + 2CO2 + 2H2 O

(4.31)

2CH2 O −−−→ CH4 + CO2

(4.32)

and Typically O2 becomes undetectable within a day of submergence and then NO3 − is reduced. Reduction of NO3 − will not occur until the O2 concentration reaches a very small value. Likewise, whilst NO3 − is being reduced, the pe is poised in the range 3–6 and reduction of Mn and Fe are prevented. However NO3 − will be exhausted within a matter of days and then reduction of Mn and Fe may proceed. In the absence of O2 Fe(III) is generally the main oxidant in the soil, its concentration typically exceeding concentrations of NO3 − , Mn(III, IV) or SO4 2− by at least an order of magnitude (Chapter 3). Between 1 and 20 % and sometimes as much as 90 % of the free Fe(III) in the soil is reduced to Fe(II) over 1–2 months of submergence (Ponnamperuma, 1972; van Breemen, 1988). Some of the structural Fe(III) in soil clays is also reduced (Stucki et al., 1997). The course of soil reduction and the changes in pe and pH are therefore generally dominated by the reduction of Fe(III). Changes in pe, pH and Alkalinity It is difficult to obtain reliable measurements of EH and hence pe in soils. Strictly, only measurements made with the electrodes in soil solution extracts rather than directly in soil are thermodynamically meaningful, and these are also subject to various errors, particularly due to the presence of mixed redox systems. Nonetheless it is a useful parameter and is the only single electrochemical property that can distinguish submerged soils from well-drained ones. Figure 4.5 shows changes in pe, pH and Fe2+ in the soil solution of four representative soils following flooding (IRRI, 1964). The figure shows that in all the soils there is a minimum in pe after a few days followed by an increase,

−1

0

1

2

3

4

5

6

0

2

4

6

8

10 21 26 30

1.67 2.78 0.30 1.2

10 12 14 16

1.5 4.1 1.5 5.1

Soil Org C Active (%) Fe (%)

5.2

5.4

5.6

5.8

6.0

6.2

6.4

6.6

6.8

7.0

7.2

0

4

6

8 10 12 14 16

Time (weeks after flooding)

2

[Fe2+] in solution (mM)

pH

0

2

4

6

8

10

12

0

2

4

6

8

10 12 14 16

Figure 4.5 Changes in pe, pH and Fe2+ in soil solutions of various soils following submergence at 25 ◦ C. pe values are calculated from measured EH (V) values in soil solutions. Data from IRRI (1964). Reproduced by permission of IRRI

pe

7

110

111

Redox Conditions in Soils

and this is characteristic of most soils following flooding (Ponnamperuma, 1972). The minimum can be less than zero and can be accompanied by evolution of H2 gas. It is due to fermentation reactions starting as soon as O2 and NO3 − are used up but before populations of Mn and Fe reducing bacteria are established. As discussed in Section 5.3, the low solubility of Mn(III, IV) and Fe(III) oxides may initially limit the rate of their reduction. Organic acids produced in fermentation reactions will help dissolve Mn(III, IV) and Fe(III) from oxide particles and thereby facilitate the establishment of the Mn and Fe reducers. As Mn and Fe reduction then proceed, the pe will increase to values corresponding to the Mn and Fe couples involved, and then gradually decline. Simultaneously H+ ions are consumed in Reactions (4.28)–(4.31) and the pH tends to increase. Initially the pH of aerobic soils may decrease following submergence because CO2 formed in aerobic respiration escapes from the soil only very slowly, and it therefore accumulates. As CO2 continues to accumulate during anaerobic respiration and fermentation, large partial pressures develop, typically in the range 5 to 20 kPa. The accumulation of CO2 lowers the pH of alkaline soils and curbs the increase in pH of acid soils. As a result the pHs of all soils tend to converge following submergence in the range 6.5–7. As the partial pressure of CO2 increases, the concentration of HCO3 − in the soil solution increases and therefore the concentrations of balancing cations in solution increase. Changes in alkalinity and concentrations of cations in solution following submergence are shown in Figure 4.6. The NH4 + , Mn2+ and especially Fe2+ ions formed in soil reduction displace exchangeable cations into solution.

25 Concentration in soil solution (mmolc L−1)

Alkalinity 20 Ca2+ + Mg2+ + NH4+ + Na+ + K+

15

Fe2+ + Mn2+

10

5

0

0

2 4 6 8 10 12 14 16 Time (weeks after submergence)

Figure 4.6 Changes in alkalinity and concentrations of cations in the soil solution following submergence (Ponnamperuma, 1972)

112

Reduction and Oxidation

Also, the changes in pH will cause changes in the charges of variable-charge clays and organic matter, thus the cation exchange capacity of acid soils will tend to increase and that of alkaline soils decrease.

Changes in Fe Large concentrations of Fe2+ develop in the soil solution in the weeks following flooding, often several mM or tens of mM (Figure 4.5). Calculations with chemical equilibrium models show that the ion activity products of pure ferrous hydroxides, carbonates and other minerals are often exceeded 100-fold (Neue and Bloom, 1989). Evidently precipitation of these minerals is inhibited, probably as a result of adsorption of foreign solutes, such as dissolved organic matter and phosphate ions, onto nucleation sites (Section 3.7). However, once a sufficient supersaturation has been reached there is a rapid precipitation of amorphous solid phases, which may later re-order to more crystalline forms. Only a small part of the Fe(II) formed in reduction remains in solution; the bulk is sorbed in exchangeable forms or, eventually, precipitated. The identities of the solid phases that form remain a mystery. Direct identification is difficult because Fe(II) and Mn(II) solid phases are readily oxidized by O2 and it is therefore necessary to maintain scrupulously anoxic conditions to ensure that the material examined actually represents that in anoxic soil. An alternative is to make indirect assessments through measurements of pe, pH and [Fe2+ ] in solution, but these too are difficult (see section on measurement of redox potential in soil). Some of the well-known solid phases that might form are shown in Table 4.3. None of these appears to be quantitatively important, at least in the first few Table 4.3 at 25 ◦ C

Some possible mineral phases in reduced soils and their equilibrium constants

Compound Mn(II) hydroxide Rhodocrosite Hauerite Fe(II) hydroxide Fe(II)Fe(III) hydroxide Siderite Vivianite Pyrite Source: a Calculated from
Gof values. b Stumm and Morgan (1996). c Lindsay (1979). d Arden (1950).

Equilibrium

log K

Mn(OH)2 (s) + 2H+ = Mn2+ + 2H2 O MnCO3 (s) = Mn2+ + CO3 2− MnS2 (s) = Mn2+ + S2 2− Fe(OH)2 (s) + 2H+ = Fe2+ + 2H2 O Fe3 (OH)8 (s) + 2H+ = Fe2+ + 2Fe(OH)3 + 2H2 O FeCO3 (s) = Fe2+ + CO3 2− Fe3 (PO4 )2· 8H2 O(s) = 3Fe2+ + 2H2 PO4 − + 8H2 O FeS2 (s) = Fe2+ + S2 2−

15.13a −10.39b −14.79c 11.67a −10.60d −10.45b 3.11c −26.93c

Redox Conditions in Soils

113

weeks or months following submergence. The large increases in CO2 pressure as dissolved Mn(II) and Fe(II) accumulate would suggest Mn and Fe carbonates should be precipitated. However it is unlikely that simple Mn and Fe carbonates are formed because Mn2+ and Fe2+ ions have similar radii (0.083 and 0.078 nm, respectively) and can readily substitute for each other in crystal lattices. Rhodocrosite (MnCO3 ) and siderite (FeCO3 ) are end members of a continuous series of solid solutions of Fe(II)–Mn(II) carbonates (Deer et al., 1992). Iron–manganese minerals also readily incorporate Mg2+ (radius 0.072 nm) and to a lesser extent Ca2+ (0.1 nm) and other divalent cations. It is therefore likely that various solid solutions are formed. There is evidence that mixed Fe(II)–Fe(III) hydroxides are formed. These can be produced easily in vitro by partial oxidation of pure Fe(II) hydroxy salts and they have some of the observed properties of the solid phase Fe(II) found in reduced soils, including the grayish-green colours characteristic of reducing conditions in soils. This material is ‘green rust’ and has the general formula Fe(II)6 Fe(III)2 (OH)18 with Al3+ partly substituted for Fe3+ and Cl− , SO4 2− and CO3 2− substituted for OH− . Once precipitation begins, a quasi-steady state will eventually be attained in which the soil pe and pH are poised by the redox and precipitation equilibria operating. In the transition to the steady state, protons will be provided by dissociation of acids in the soil solution—e.g. H2 CO3 derived from CO2 –and by reactions with the soil exchange complex. The course of reduction and the eventual steady state will depend on these reactions and it is therefore necessary to allow for them in predicting what the steady state conditions will be. In the following section I describe a simple model for calculating the changes in pe, pH and concentrations of inorganic reductants during soil reduction, allowing for the effects of pH buffering and cation exchange, and the characteristics of the mineral phases formed. The approach is based on that of van Breemen (1988) for partial redox equilibrium in soil without pH buffering and cation exchange. 4.2.3 CALCULATED CHANGES IN pe, pH AND Fe DURING SOIL REDUCTION Consider an idealized soil containing ferric hydroxide and readily decomposable organic matter. The following conditions hold: • the soil is initially saturated with the atmospheric partial pressure of O2 but otherwise closed to exchange of O2 ; • the partial pressure of CO2 is constant; • the soil exchange complex is initially saturated with divalent cations M2+ , i.e. H+ is treated as non-exchangeable and there are no other monovalent cations; • the soil reaches a steady state following reduction in which the soil solution is in equilibrium with Fe(OH)3 and Fe3 (OH)8 .

114

Reduction and Oxidation

Following flooding, O2 dissolved in the soil solution is consumed according to Reaction (4.27). There is no pH change, the CO2 pressure being constant, and pe is poised by the O2 –H2 O couple: O2 (aq) + 4H+ + 4e− = H2 O i.e. pe = 21.45 + 41 log[O2 ]L − pH

(4.33)

Once all the O2 has been used up Fe(OH)3 is reduced according to Reaction (4.30) and the pe is poised by the Fe(OH)3 –Fe2+ couple: Fe(OH)3 (s) + 3H+ + e− = Fe2+ + 3H2 O i.e. pe = 16.54 − log[Fe2+ ]L − 3pH

(4.34)

Once pe falls sufficiently, precipitation of Fe3 (OH)8 commences and Fe3 (OH)8 is formed at the expense of Fe(OH)3 according to the reaction 3Fe(OH)3 + CH2 O + 8H+ → Fe3 (OH)8 + CO2 + 11H2 O The pe is now poised by the Fe(OH)3 –Fe3 (OH)8 half reaction: 3Fe(OH)3 (s) + H+ + e− = Fe3 (OH)8 + H2 O i.e. pe = 1.46 − pH

(4.35)

Hence for a given generation of Fe2+ in Fe(OH)3 reduction, and for a specified initial soil CEC and concentration of non-carbonate anions in the soil solution ([X− ]L ), we have five unknowns: the soil pH and the concentrations of Fe2+ and M2+ in the soil solid and solution; and these can be found from the following five equations: (1) Equation (3.69) for the electrical neutrality of the solution, with [HCO3 − ] found from pCO2 and pH; (2) Equation (3.70) for the electrical neutrality of the solid, with changes in acidity in the solid related to changes in pH with the soil pH buffer capacity; (3) Equation (3.72) for divalent–divalent cation exchange; and (4) and (5) two equations like Equation (3.73) for conservation of M2+ and Fe2+ . These equations can be solved simultaneously with Equations (4.33)–(4.35) to obtain values of pe, pH, [O2 ] and [Fe2+ ] over the course of reduction. Figure 4.7 shows results for realistic flooded soil conditions, expressed in terms of the amounts of CH2 O oxidized in the different reactions. Figure 4.7(a) gives results in the absence of pH and cation buffering by the soil; Figure 4.7(b)–(d) gives results for different values of bHS , CEC and [X− ]L .

0.0

0.5

1.0

1.5

2.0

2.5

3.0

3.5

4.0

0

2

O2

4

6

8 10 12 14 16

pe

pH

Fe2+

bHS = 0, CEC = 0, [X−]L = 0

0

(b)

2

4

O2 6

0

(c)

2

4

O2 6

8 10 12 14 16

pe

pH

Fe2+

bHS = 50, CEC = 35, [X−]L = 20

CH2O oxidized (mmol kg−1)

8 10 12 14 16

pe

Fe2+

pH

bHS = 15, CEC = 35, [X−]L = 20

0

(d)

2

O2 4

6

pe

pH

−2

0

2

4

6

8

10

12

14

16

18

8 10 12 14 16

Fe2+

bHS = 50, CEC = 10, [X−]L = 10

4.0

4.5

5.0

5.5

6.0

6.5

7.0

7.5

8.0

pH pe

Figure 4.7 Calculated changes in pe, pH, [O2 ] and [Fe2+ ] in an idealized soil during reduction. Mineral phases Fe(OH)3 and Fe3 (OH)8 . Parameters for pH buffering and cation exchange differ between (a)–(d) as indicated. Units of bHS (soil pH buffer power), mmol kg−1 pH−1 ; CEC (initial cation exchange capacity), cmolc kg−1 ; and [X− ]L (concentration of non-carbonate anions), mM. CO2 pressure = 10 kPa, θ = 0.6, ρ = 1.0, initial pH = 4.5

[Fe2+], [O2] in solution (mM)

(a)

115

116

Reduction and Oxidation

It will be seen that pe is initially buffered at about 15 until the O2 is exhausted, and it then falls rapidly to the point where it is buffered by the Fe(OH)3 –Fe2+ couple. Ferrous ions are now released into the solution and protons removed from it, and the changes in pe and pH now depend on the buffering of Fe2+ and H+ by the soil solid. As a result of buffering, the increases in Fe2+ and pH as Fe(OH)3 is reduced are more gradual. Comparing Figure 4.7(b) and (c), for which the CEC and [X− ]L are the same but bHS different, the effect of increasing bHS is to further slow the increase in pH, and the pH at the steady state when Fe3 (OH)8 is formed is smaller and the pe and concentration of Fe2+ in solution correspondingly larger. Also a much larger quantity of CH2 O is consumed in reaching the steady state. From the stoichiometry of Reaction (4.30), the amount of exchangeable Fe2+ formed in mmolc kg−1 is roughly four times the amount of CH2 O oxidized. The effect of varying CEC can be seen by comparing Figure 4.7(c) and (d). With decreasing CEC at constant bHS , [X− ]L and PCO2 , the concentration of Fe2+ at steady state is increased and the pH correspondingly decreased and pe increased. In summary, the calculations predict that: (1) O2 will disappear rapidly after flooding; (2) dissolved Fe2+ will appear in the soil solution and its concentration increases to a constant steady-state level; (3) an initially low pH will increase to between 6.5 and 7 at the steady state; (4) pe will decrease from about 15 to near 0; (5) the rates of change in Fe2+ , pH and pe and their steady-state values, and the amounts of organic matter oxidized in reaching the steady state, depend on pH buffering and cation exchange by the soil. These predictions can be compared with the results for real soils shown in Figure 4.7. In the real soils the ranges of pe and pH are similar and a steady state is attained in which the concentrations of Fe2+ in solution are similar to those predicted with the model. However the large peak in Fe2+ concentrations in some soils before the steady state is reached is not predicted. The peak occurs because precipitation of ferrous carbonate is slow and may be inhibited by interfering solutes in the soil, resulting in supersaturation with respect to the expected solid phases. Note that although the pe–pH–[Fe2+ ] relationships shown in Figure 4.7 are consistent with control by the Fe(OH)3 –Fe3 (OH)8 system, in fact various other reduced Fe solid phases are possible and as discussed above it is difficult to establish unequivocally which phase controls Fe2+ solubility in reduced soils. 4.2.4 MEASUREMENT OF REDOX POTENTIAL IN SOIL In principle the redox potential provides a simple means of gauging a soil’s redox status. However in practice it is difficult to make reliable measurements.

Redox Conditions in Soils

117

Stumm and Morgan (1996) discuss the problems for simple aquatic systems and van Breemen (1969), Ponnamperuma (1972), McBride (1994) and Patrick et al. (1996) discuss the additional problems for soil systems. I here give the main points. Measurements of EH are usually made with a platinum electrode placed in the soil solution together with a reference half cell electrode of known potential. The platinum electrode transfers electrons to and from the soil solution without reacting with it. Reducing half reactions in the soil tend to transfer electrons to the platinum electrode and oxidizing half reactions to remove them. At equilibrium no electrons flow and the electric potential difference between the half cell comprising the platinum electrode and the soil solution and the half cell comprising the reference electrode is recorded. The first problem to mention is that thermodynamically meaningful measurements of EH must be made on soil solution extracts and not directly on the soil itself. Although EH values measured in soil following reduction may show the expected qualitative trends and expected differences between soils, they are not satisfactory for quantitative interpretation. Hence duplicate measurements of EH in soil can vary by as much as 100 mV and values are generally far too low in terms of the Fe2+ and Mn2+ concentrations measured (IRRI, 1964). This is firstly because the measurement indicates the potential in the immediate vicinity of the electrode and not that of the whole soil, and there may be large microscale variations in EH especially near the surfaces of bacterial cells. Secondly they are subject to liquid junction potential errors. It is therefore necessary to make measurements in solution withdrawn from the soil ensuring the minimum of gas exchange during sampling. It is particularly important that no O2 is allowed to enter the solution and that no CO2 is lost. This may be achieved by withdrawing the solution through porous tubing into previously evacuated tubes. Apart from these sampling errors there are a number of intrinsic errors in the measurement of EH in soil solutions. The measurement depends on there being no net flow of current through the circuit made by the platinum electrode and reference electrode. However the current in one direction, called the exchange current, i0 , is not zero. Its value for each half-reaction varies with the electrode potential and with the concentrations of the oxidant and reductant. Figure 4.8 shows this schematically for the Fe2+ –Fe3+ couple. As can be seen from the figure, an infinitesimal shift in the electrode potential from its equilibrium value will make the half reactions proceed in one direction or the other and a net current will flow through the circuit. The equilibrium potential of the system can be found from the potential at which no net current flows. How precisely and reproducibly this measurement indicates the equilibrium potential depends on how steeply the net current deviates from zero near the equilibrium potential. The greater the exchange current, i0 , the more steeply the net current varies with the potential. This in turn depends on the redox couple operating and its concentration. Modern instruments will give reliable measurements for i0 values greater than about 0.1 µA. Figure 4.8 shows that for the Fe2+ –Fe3+ couple, i0 ≈ 100 µA

118

Reduction and Oxidation net current

(a) −100

(b) −100

Fe3+→ Fe2+ −i 0

−50

−50 Fe3+→ Fe2+

← Potential (mV) 525 Current (µA)

net current

0

0 475

500

525

475 Fe2+→ Fe3+

50

+i 0

Fe2+→ Fe3+

50

100

100

(c) −50

−i 0

0 525 50

net current

Fe3+→ Fe2+

475

500

450 +i 0 425

+i 0 Fe2+→ Fe3+

Figure 4.8 Electrode current versus electrode potential curves for the Fe2+ –Fe3+ couple in water at pH 2 with (a) [Fe3+ ] = [Fe2+ ] = 1 mM; (b) [Fe3+ ] = [Fe2+ ] = 0.1 mM; (c) [Fe3+ ] = 0.1 mM, [Fe2+ ] = 1 mM. Electrode area = 1 cm2 (Stumm and Morgan, 1996). Reproduced by permission of Wiley, New York

for [Fe3+ ] = [Fe3+ ] = 10−3 M (Figure 4.8a). If the concentration of both ions is 10-fold smaller, i0 and the slope are 10-fold smaller (Figure 4.8b). However if the concentration of only one of the ions is decreased the drop in i0 is not as great (Figure 4.8c); note also that the equilibrium potential is shifted. If [Fe3+ ] = [Fe3+ ] = 10−7 M, i0 ≈ 0.1 µA and measurements are no longer reliable. In practice the limiting concentration is nearer 10−5 M because of the effects of trace impurities. The value of i0 will increase with the surface area of the electrode. However the benefit of this tends to be offset by greater effects of impurities. In the case of the O2 –H2 O couple, the net current is virtually zero over a wide range of electrode potentials as shown in Figure 4.9(a). This makes it extremely difficult to determine the equilibrium potential for the O2 –H2 O couple, and so EH measurements in aerated soils are not reliable. A further problem, particularly in soil systems, is that several redox systems may be present, in which case the apparent equilibrium potential may be the result

119

Transformations of Nutrient Elements Accompanying Changes in Redox (a)

(b) H2O→H2

−1

Current (mA)

O2→H2O

O2→H2O Fe3+→Fe2+

0

0 +1 H2O→O2

1

−1 0 Potential (V)

Em

Eeq

Potential (V)

Fe2+→ Fe3+

Figure 4.9 Electrode current versus electrode potential curves for solutions containing O2 : (a) in otherwise pure water; (b) in the presence of Fe2+ . In (a) the net current is close to zero over a wide range of potential, so it is difficult to locate the equilibrium potential. In (b) the measured equilibrium potential is a mixed potential, Em , obscuring the true equilibrium potential of the system, Eeq (Stumm and Morgan, 1996). Reproduced by permission of Wiley, New York

of the combined exchange currents of two or more redox couples. Figure 4.9(b) illustrates this for the Fe2+ –Fe3+ system in the presence of trace concentrations of dissolved O2 . The measured equilibrium potential, Em , at which the net current is zero may be the potential at which the rate of reduction of O2 at the electrode equals the rate of Fe2+ oxidation. This would be likely if the concentration of Fe2+ greatly exceeded that of Fe3+ , as in general it will in submerged soils. The two couples are not in equilibrium with each other and the measured potential is termed a mixed potential. The mixed potential does not represent either of the individual couples operating and is therefore difficult to interpret. Many redox couples do not react reversibly at electrode surfaces. Examples are CO2 –CH4 and NO3 − –N2 . This too complicates interpretation. These factors rather constrain the usefulness of EH measurements in soil solutions. Inferences about the thermodynamics of redox processes in soils that rely heavily on measurements of redox potential should be treated with caution. Nonetheless soil EH measurements provide a ready measure of redox status, for example in experiments in which constant EH and pH are required (Patrick et al., 1973). 4.3 TRANSFORMATIONS OF NUTRIENT ELEMENTS ACCOMPANYING CHANGES IN REDOX These are briefly discussed here in the context of redox chemistry. More complete discussions are given in Chapters 5–8.

120

Reduction and Oxidation

4.3.1 TRANSFORMATIONS OF CARBON In broad terms the decomposition of organic matter under anaerobic conditions is expected to be slower than under aerobic conditions because the free energy changes for the reactions involved are much smaller (Table 4.1 and Figure 4.3). For example, for the aerobic decomposition of ‘CH2 O’, 1 ‘CH2 O’ 4

+ 14 O2 = 14 CO2 (g) + 41 H2 O


Go = −119 kJ mol−1 at pH 7, whereas for its anaerobic decomposition in methanogenesis, 1 ‘CH2 O’ = 41 CO2 (g) + 41 CH4 (g) 4
Go = −17.7 kJ mol−1 at pH 7. Consequently the microbes mediating the decomposition derive less energy and produce fewer cells per unit of carbon metabolized. The accumulation of organic matter in marshes and peat bogs illustrates this point. (But note the rarity of tropical wetland soils with large organic matter contents, discussed in Section 3.7.) The most striking difference between anaerobic and aerobic decomposition is in the nature of the end products. In aerobic decomposition the main products are CO2 , NO3 − , SO4 2− and resistant residues; in anaerobic decomposition they are CO2 , H2 , CH4 , N2 , NH4 + , H2 S and various partially decomposed and humified residues. The decomposition proceeds in two stages. The first involves formation of organic acids, particularly acetic, propionic and butyric, plus various aliphatics and phenolics, some of which are toxic to plants. The second involves conversion of organic acids to gaseous products and follows a characteristic pattern. In the first few days, H2 formed in fermentation reactions may be evolved together with CO2 . Nitrogen gas is also evolved, formed in denitrification of NO3 − . As inorganic redox couples then begin to buffer the redox potential, H2 evolution ceases and CO2 is the main end product of carbohydrate metabolism. This continues until the pe and pH reach values at which methanogenesis is possible, typically 1 or 2 weeks after submergence. The concentration of CH4 in the soil solution and in gas bubbles then exceeds the concentration of CO2 several-fold as a result of solubility and precipitation effects. Although there is wide variation in the composition of gases formed between soils, this general pattern is always seen. At higher temperatures, CO2 and CH4 are formed sooner and at greater rates. Also, at higher temperature and pH, the ratio of CH4 to CO2 in the soil gases changes in favour of CH4 because of solubility and precipitation effects and the higher optimal temperatures for methanogens. 4.3.2 TRANSFORMATIONS OF NITROGEN The main transformations of N are summarized in Figure 4.10. In the absence of oxygen, mineralization of organic N proceeds only as far as NH4 + , and NH4 +

121

Transformations of Nutrient Elements Accompanying Changes in Redox

NO2

Oxidation state of N

NO2−

+2

NO

+1

N2O

0

N2

denitrification

+3

N fixation −3

NH4+

pe

Oxidation

+4

12

Reduction

NO3−

nitrification

+5

−4

immobilization Organic N mineralization

Figure 4.10 Nitrogen transformations in submerged soils on a redox scale (McBride, 1994). Reproduced by permission of Oxford University Press

accumulates in the soil solution and exchange complex. Because of the low N requirement of anaerobic metabolism, subsequent immobilization by microbes tends not to be important, or, if it occurs—as when organic matter with a wide C:N ratio is present—the immobilization is temporary. Further transformations of N take place at the oxic interfaces between the soil and floodwater and root and soil where NH4 + diffusing in from the neighbouring anoxic soil may be nitrified to NO3 − . Subsequently, NO3 − diffusing out into the anoxic soil may be denitrified to N2 . This process results in significant losses of N from wet soils but its importance in submerged soils is unclear (Section 5.3). Under strongly reducing conditions (pe < −4) reduction of N2 to NH4 + is thermodynamically possible. The net reaction is 1 N (g) 6 2

+ 31 H+ + 14 ‘CH2 O’ = 31 NH4 + + 41 CO2 (g)


Go = −14.3 kJ mol−1 at pH 7. However this reaction has a very large activation energy because of the energy required to break the N≡N triple bond (942 kJ mol−1 ). Therefore only highly specialized ‘nitrogen fixing’ organisms are capable of maintaining sufficiently reducing conditions in their cells to mediate the reaction. The niches in submerged soils in which nitrogen fixers may operate are discussed in Chapter 5. Most of the mineralizable N in the soil is converted to NH4 + within a few weeks of submergence if the temperature is favourable and the soil not strongly acid or deficient in other nutrients. The concentration of NH4 + in the soil solution typically reaches 0.1 to 5 mM buffered by from 5 to 20 times this concentration

122

Reduction and Oxidation

Concentration of P in soil solution (µM)

(c)

(b) 8

6 25 18 23

4

21 29 28 1

2

0

27 0

2

4

6

8

Concentration of SO42− in soil solution (mM)

Concentration of NH4+ in soil solution (mM)

(a)

10 12 14 16

12 16.5 mM at 0.5 wk

10

39

8 6 1

4

23

21 2 14

26

0

0

2

4

6

8

10 12 14 16

140 120 100 1 80 60 25

40 20 0

0

26 27 14 2 4 6 8 10 12 14 16 Time (weeks after flooding)

Soil

pH

1 14 18 21 23 25 26 27 28 29 39

7.6 4.8 5.6 4.6 5.7 4.8 7.6 6.6 4.9 5.8 8.1

Org C Active (%) Fe (%) 2.3 2.8 6.0 4.1 8.0 4.4 1.5 2.0 2.9 7.7 2.0

0.18 2.13 0.27 2.78 0.47 0.18 0.30 1.60 4.70 1.80 -

Figure 4.11 Changes in (a) NH4 + , (b) SO4 2− and (c) P in the soil solution of various soils following flooding (modified from IRRI, 1964, 1965). Reproduced by permission of IRRI

of NH4 + on the soil exchange complex. Figure 4.11 shows changes in NH4 + in solution following submergence of a range of soils.

4.3.3 TRANSFORMATIONS OF SULFUR The stable form of sulfur under moderately strong reducing conditions (pe < −3) is hydrogen sulfide, H2 S, which is readily soluble and under non-acid conditions

Transformations of Nutrient Elements Accompanying Changes in Redox

123

dissociates to HS− (pK = 7.02). For the reduction of SO4 2− the net reaction is 1 ‘CH2 O’ 4

+ 18 SO4 2− + 81 H+ = 14 CO2 (g) + 81 HS−

and
Go = −20.5 kJ mol−1 at pH 7. H2 S and HS− are also produced in the hydrolysis of the S-containing amino acids. The HS− formed further dissociates to S2− (pK = 13.9). However in most submerged soils the concentration of Fe2+ in the soil solution is sufficient that virtually all S2− is precipitated as amorphous ferrous sulfide and very small concentrations of H2 S and HS− remain in solution. The relations between the SO4 2− –HS− and Fe(OH)3 –Fe2+ systems at neutral pH are shown in Figure 4.12. Amorphous ferrous sulfide may gradually crystallize as mackinawite (FeS). Under some circumstances pyrite is then formed, e.g. FeS(s) + S(s) → FeS2 (s), leading to potential acid sulfate soils (Section 7.3). There may be a cycling of S compounds of different oxidation state between anaerobic and aerobic zones in the soil, such as at the soil—floodwater interface. In reduced lake and marine sediments this leads to accumulation of insoluble sulfides as SO4 2− carried into the sediment from the water above is immobilized. Such deposits function as sinks for heavy metals. Plants absorb S through their roots as SO4 2− ; H2 S is toxic to them. Therefore HS− must be oxidized to SO4 2− in the rhizosphere before it is absorbed. Figure 4.12 shows changes in the concentration of SO4 2− in the soil solution following submergence of a range of soils. In neutral and alkaline soils concentrations of SO4 2− greater than 10 mM may decrease to 0 within 6 weeks of submergence. In acid soils the concentration of SO4 2− in solution may initially increase following submergence and then slowly decline over several months. FeS

FeCO3

Fe(OH)3

log concentration (M)

0 −2

H2S + HS−

SO42−

−4 −6 −8

Fe2+

−10 −12 −8

−6

−4

−2

0

2

4

6

8

10

pe

Figure 4.12 Concentration–pe diagram for FeS, FeCO3 and Fe(OH)3 at pH = 7, CT (total carbonate carbon) = 5 mM and [SO4 2− ] + [H2 S(aq)] + [HS− ] = 1 mM (modified from Stumm and Morgan, 1996). Reproduced by permission of Wiley, New York

124

Reduction and Oxidation

The initial increase occurs because SO4 2− sorbed on variable charge clays and oxides is desorbed as the pH increases. The rate of subsequent reduction will be low if the pH remains below 5.5, the optimal range of pH for SO4 2− reducing bacteria being greater than this.

4.3.4 TRANSFORMATIONS OF PHOSPHORUS Phosphorus is often the most limiting nutrient in natural wetlands. Because of its association with soil Fe, its solubility changes markedly during reduction and oxidation. In general it is not itself reduced and remains in the +5 oxidation state, though production of phosphine gas (PH3 ; +3 oxidation state) at rates ≤ 6.5 ng m−2 h−1 has been reported in laboratory experiments with brackish and saline marsh soils (Devai and Delaune, 1995). Review articles on transformations of P in submerged soil include Patrick and Mahapatra (1968), Kirk et al. (1990a) and Willett (1991). Typically when a soil is submerged the concentrations of water- and acidsoluble P increase, reach a peak or plateau, and then decrease (Figures 4.11c and 4.13). For the soils shown in the figures, the peak P concentrations in solution were smallest for acid soils high in active Fe and greatest for a sandy soil low in Fe. The increases in acid-soluble P were greatest in an alkali soil low in active

Concentration of acid-soluble P in soil (mmol kg−1)

1.2

1.0

26

0.8

0.6

27 18

0.4

21 0.2

28 14

0.0

0

2 4 6 8 10 Time (weeks after submergence)

12

Figure 4.13 Changes following flooding in the concentration of P soluble in an acetate buffer at pH 2.7. Numbers next to curves identify soils; properties given in table in Figure 4.11 (modified from Ponnamperuma, 1985). Reproduced by permission of IRRI

Transformations of Nutrient Elements Accompanying Changes in Redox

125

Fe, intermediate in sandy loams high in organic C and low in active Fe, and least in acid clays high in active Fe. The increases in soluble P are particularly linked to the transformations of Fe and changes in pH. The main processes are: • reduction of Fe(III) compounds holding P on their surfaces and within their crystal lattices; • dissolution of Ca-P compounds in alkaline soils as the pH decreases and desorption of P held on variable-charge surfaces in acid soils as the pH increases; • displacement of sorbed P by organic anions and chelation of metal ions that would otherwise immobilize P; and • mineralization of organic P. Subsequent decreases in solubility may be due to re-sorption or precipitation on clays and oxides as soil conditions continue to change, and decomposition of organic anions chelating P or chelating Al and Fe with which it would otherwise react. Following submergence soils often release more P to solutions low in P but adsorb more P from solutions high in P. This apparent paradox can be explained by the reduction of Fe(III) oxides to poorly ordered gel-like Fe(II) compounds with large surface areas. Phosphorus solubilized in soil reduction is sorbed on the amorphous surfaces and desorbed when P is removed from the soil solution; but fresh P added to the soil is removed from solution by sorption onto the Fe(II) surfaces. Consequently many soils do not show significant increases in P solubility during flooding (Willett, 1991), and with prolonged flooding the P may become re-immobilized in less soluble forms. Gradual immobilization of P with prolonged anaerobicity is shown in Figure 4.14, which gives change in labile P over 3 years of double rice cropping of a perennially wet soil (B¨ucher, 2001). The labile P declines even in plots that received sufficient P fertilizer to more than off-set crop removals. Periodic drying of the soil during the fallow periods tended to increase labile P in the soil, but not in years when the soil remained anaerobic during the fallow (following the 1998 and 1999 wet season crops). The effect was greatest when tillage was delayed until the end of the fallow, resulting in more-reducing conditions in the soil, and it carried through to the succeeding rice crop. Supporting laboratory and greenhouse studies showed that changes in soil Fe with reduction and oxidation were responsible for the changes in P. Rapid drying and oxidation of the soil can also result in the P becoming very insoluble (Brandon and Mikkelsen, 1979; Willett, 1979; Sah et al., 1989; Huguenin-Elie et al., 2003). Re-oxidized Fe(II) compounds may be precipitated in poorly crystalline forms with large specific surface areas, on and in which P becomes immobilized. Hence upland crops grown in rotation with rice frequently suffer P deficiency even though crops on similar soils not used for rice grow healthily. The problem is in part also due to disruption of mycorrhizal networks during flooding (Ilag et al., 1987; Ellis, 1998; Miller, 2000).

NK plots

1999 DS

End of fallow 2 Dat Time

Fallow

1999 WS Fallow 2000 DS

42 Dat

21 Dat

End of fallow

2 Dat

Mid-fallow 63 Dat Harvest

42 Dat

21 Dat

End of fallow 2 Dat Mid-fallow

63 Dat Harvest

42 Dat

21 Dat

Mid-fallow

4 Dat

Resin-extractable P (mg kg−1 dry soil)

Figure 4.14 Changes in labile soil P (extractable with HCO3 − -form anion exchange resin) during 3 years of wetland rice cropping as affected by timing of tillage (early, late = start, end of fallow), incorporation of previous crop’s straw, and application of P (20 kg ha−1 in NPK plots). The overall P balances over 3 years were +37 and +7 kg P ha−1 in the NPK plots with and without straw, and −90 and −115 kg P ha−1 in the PK plots. DS, WS, dry, wet season; DAT, days after transplanting (B¨ucher, 2001). Reproduced by permission

0

5

10

0 15

5

21 Dat

Early tillage, no straw

Late tillage, Late tillage, with straw no straw

Harvest

10

Mid-fallow

15

42 Dat

Early tillage, with straw

Fallow

63 Dat Harvest

1998 WS

End of fallow 2 Dat

NPK plots

63 Dat

Fallow

21 Dat

1998 DS

42 Dat

20

Fallow

63 Dat Harvest

25

126

127

Oxidation of Reduced Soil

4.4 OXIDATION OF REDUCED SOIL When a spadeful of wet, anaerobic soil is brought to the surface and allowed to dry, air enters through drying cracks and the soil tends to become uniformly oxidized and turn a uniform brown. Whereas when oxidation occurs without drying—as, for example, near a root releasing O2 into wet soil—it is far less uniform and reddish-brown ferric oxide deposits form on and near the oxidizing source. The difference depends on the relative rates of movement of O2 into the soil and of ferrous iron and other reductants in the opposite direction, and the rates of reaction. Figure 4.15 indicates the range of rates of O2 consumption in different soils. Oxygen is consumed in oxidation of inorganic reductants, such as Fe(II), as well as in oxidation of organic matter by microbes. Bouldin (1968) and Howeler and Bouldin (1971) compared measured rates of O2 movement into anaerobic soil cores with the predictions of various models, and obtained the best fits with a model allowing for both microbial respiration and abiotic oxidation of mobile and immobile reductants; abiotic oxidation accounted for about half the O2 consumed. The kinetics of the abiotic reactions are complicated. They often depend on the adsorption of the reductant on solid surfaces as, for example, in

[O2]/[O2]initial

1

0.1

pH Org C [Fe2+] (%) (µmol g−1) 6.2 6.6 5.9 6.8 5.6 5.6 7.6

0.01

0

1.60 2.30 0.82 1.71 0.72 1.01 0.54

20

42.3 39.6 33.6 18.9 26.9 5.3 16.5

40 Time (h)

60

80

Figure 4.15 Rates of oxygen consumption by shaken suspensions of anaerobic soils. Points are measured data, lines are fits to two first-order rate equations. The apparent rate constant for the initial reaction is common to all soils; that for the main reaction varies 30-fold between the soils and is well correlated with [Fe2+ ] (Reddy et al., 1980). Reproduced by permission of Soil Sci. Soc. Am.

128

Reduction and Oxidation

the autocatalysis of Fe2+ oxidation by adsorption of Fe2+ on ferric hydroxide formed in the reaction. The adsorption is likely to be pH-dependent, a decrease in pH tending to decrease sorption and increase the concentration of Fe2+ in solution. Hence there may be complex interactions between the mobility of Fe2+ , the rate of oxidation, and pH changes caused by the reaction. Such interactions can produce banded distributions of iron around an O2 source, as found, for example, by Saleque & Kirk (1995) for the distribution of iron near rice roots and calculated by Kirk et al. (1990) with a model of the coupled diffusion and reaction of O2 , Fe2+ and acidity in soil. This is an example of the Liesegang phenomenon (Stern, 1954; Keller, 1980). 4.4.1 KINETICS OF Fe2+ OXIDATION Aqueous Solution The reaction between Fe2+ and O2 to form insoluble ferric hydroxide can be written 4Fe2+ + O2 + 10H2 O = 4Fe(OH)3 + 8H+ (4.36) Equation (4.36) shows that two H+ ions are produced for each mole of Fe2+ oxidized, i.e. the reaction is accompanied by acidification. In aqueous solution, the rate is found to be very sensitive to pH and at near neutral pH the reaction is accelerated 100-fold if the pH is raised by one unit. The following empirical rate law applies in the pH range 5–8 (Stumm and Lee, 1961; Wehrli, 1990) −d[Fe(II)]/dt = k[O2 ][OH− ]2 [Fe(II)]

(4.37)

where k ≈ 2 × 1014 mol3 dm−9 s−1 at 25 ◦ C and [Fe(II)] is the sum of the concentrations of Fe(II) species present—Fe2+ and its hydroxy complexes, FeOH+ and Fe(OH)2 , for which the formation constants are 10−4.5 mol−1 dm3 and 10−7.4 mol−2 dm6 , respectively. Therefore [Fe(II)] ≈ [Fe2+ ], but the pH dependence of the rate is due to the parallel oxidation of the three species. At [O2 ] = 0.28 mM (i.e. in equilibrium with atmospheric PO2 ), the half time for the reaction is 0.34 h at pH 7 and 143 days at pH 5. As discussed in Section 4.1, most redox reactions reach equilibrium only slowly if they are not catalysed. Oxidation of Fe2+ is catalysed by adsorption of Fe2+ onto Fe(OH)3 formed in the reaction, so Equation (4.36) only holds for the initial rates of reaction. Tamura et al. (1976) studied the oxidation of a solution of Fe2+ at different controlled pHs near neutral and with varying additions of Fe(OH)3 . The reaction obeyed the rate law −d[Fe2+ ]/dt = k[O2 ][OH− ]2 [Fe2+ ] + kS [O2 ][Fe2+ ]ad

(4.38)

where [Fe2+ ] is the concentration in solution, [Fe2+ ]ad the concentration adsorbed on Fe(OH)3 and kS the rate constant for oxidation of adsorbed Fe2+ (= 73 mol−1

Oxidation of Reduced Soil

129

dm3 s−1 , with all concentrations in mol dm−3 suspension). Adsorption is described by [Fe2+ ]ad /[Fe2+ ] = K[Fe(III)]/[H+ ] (4.38a) where [Fe(III)] is the concentration of Fe(OH)3 and K = 10−14.3 . Other metal oxidation reactions catalysed by sorption onto oxide surfaces are described in Section 7.3.

Soil A similar catalysis occurs on soil surfaces. Ahmad and Nye (1990) and Kirk and Solivas (1994) studied the kinetics of Fe2+ oxidation in soil suspensions by measuring changes in extractable Fe2+ in the whole soil and in solution during oxidation at constant [O2 ]. They found that 75 % of the initial Fe2+ was oxidized rapidly (t1/2 ≈ 2 h) and the remainder only very slowly (t1/2 ≈ 8 days). In the soils studied, the pH fell from near neutral to less than 5 over the course of the fast reaction. Measurements of the fast reaction at constant pH (Figure 4.16) showed that the oxidation of adsorbed Fe2+ was much faster than solution Fe2+ , and that the adsorbed Fe2+ was oxidized at a rate that was nearly independent of pH. Figure 4.16 shows that the overall rate of oxidation is more dependent on the concentration of sorbed Fe2+ ([Fe2+ ]S ) than the concentration in solution ([Fe2+ ]L ). Thus, although at pH 6.5 [Fe2+ ]L drops to one-tenth of its initial value within 1 h, d[Fe2+ ]/dt does not decrease to nearly the same extent. The figure also shows that oxidation of sorbed Fe2+ , indicated by the slopes of the lines in Figure 4.16(c), is surprisingly little influenced by pH and roughly follows first-order kinetics. The overall rate equation at constant pH is therefore −d[Fe2+ ]/dt = RkL [O2 ]L [Fe2+ ]L + kS [O2 ]L [Fe2+ ]S

(4.39)

where kL and kS are the rate constants for the reactions in the solution and solid and R the solution to solid ratio. Values of kS , calculated from the data in Figure 4.16(c) range from 0.19 mol−1 dm3 s−1 at pH 6.5 to 0.15 mol−1 dm3 s−1 at pH 5. Initially, most of the readily oxidizable Fe(II) is sorbed on the soil exchange complex. As the soil is oxidized, the Fe(OH)3 formed provides fresh sorption sites, as well as possibly blocking some of the original sites. Ahmad and Nye (1990) estimated the importance of the freshly formed Fe(OH)3 in sorbing Fe2+ compared with the original soil exchange complex, and found that the importance of the freshly formed Fe(OH)3 was much greater at higher pH, consistent with the expected greater pH-dependence of sorption on Fe(OH)3 surfaces. They also found that the oxidation of Fe2+ when sorbed on a mixture of soil exchange and Fe(OH)3 sites was much slower than on Fe(OH)3 in the absence of soil, described by Equations (4.38) and (4.38a).

130

Reduction and Oxidation (b) 100

(a) 100 90

4.5

4.5

30

5.0 5.0

6.0

60

10 [Fe2+]L (mM)

[Fe2+] (mmol kg−1)

75

6.5

45

5.5 3 1

6.0 6.5

0.3 0.1 30

0

50

100 150 Time (min) (c)

200

250

0

50

200

250

90 75 6.0

60 [Fe2+]s (mmol kg−1)

100 150 Time (min)

45 30

5.0 4.5

15

0

50

100 150 Time (min)

200

250

Figure 4.16 Changes in concentrations of Fe2+ in (a) whole soil, (b) soil solution and (c) soil solid during oxygenation of reduced soil suspensions at different pHs. [Fe2+ ]S was calculated from [Fe2+ ]–R[Fe2+ ]L (Kirk and Solivas, 1994). Reproduced by permission of Blackwell Publishing

A possible explanation is that access of O2 to the exchange or Fe(OH)3 sites where the Fe2+ is adsorbed is restricted. Possibly Fe(OH)3 is precipitated between clay lamellae at the oxidation sites and it partially blocks the original exchange sites. This mechanism would also imply a wide range of reaction rates between soils, with kS values ranging by perhaps an order of magnitude, as in Figure 4.15. In summary, the reaction can be represented by the following simplified scheme: kS

Fe2+ L ⇀ ↽ Fe2+ S −−−→ Fe(OH)3

131

Oxidation of Reduced Soil

in which the exchange of Fe2+ between the solid and liquid is rapid and the overall rate depends only on the approximately first-order oxidation of sorbed Fe2+ . Although this rate is independent of pH, the distribution of Fe2+ between the solid and liquid is not and it is therefore necessary to allow for pH changes in calculating the rate.

4.4.2 SIMULTANEOUS DIFFUSION AND OXIDATION IN SOIL Kirk et al. (1990b) and Kirk and Solivas (1994) used the above understanding of oxidation kinetics to develop a model of soil oxygenation. The model allows for the diffusion of O2 into the soil, the diffusion of Fe2+ towards the oxidizing surface, the rate of formation and concentration profile of the Fe(OH)3 formed, and the diffusion by acid–base transfer of the acidity formed: H3 O+ diffusing away from the zone of acidification and HCO3 − (derived from CO2 ) towards it. The principal equations are as follows, expressed in planar geometry so as to be able to test the predictions against experimentally measured reactant profiles. (1) For the diffusion and reaction of O2 :   ∂ ∂[O2 ]L ∂[O2 ] = DLO θf − 41 S1 − S2 ∂t ∂x ∂x

(4.40)

where [O2 ] and [O2 ]L are the concentrations of O2 in the whole soil and solution, respectively, S1 is the rate of Fe2+ oxygenation, S2 is the rate of O2 consumption in microbial respiration, and the other parameters are as defined in Chapter 2. (2) For the diffusion and reaction of Fe2+ :   ∂[Fe2+ ]L ∂[Fe2+ ] ∂ (4.41) DLI θf − S1 = ∂t ∂x ∂x where [Fe2+ ] and [Fe2+ ]L are the concentrations of mobile Fe2+ in the whole soil and solution, respectively. (3) For the diffusion and reaction of soil acidity (Section 2.2)   ∂[HS] ∂pH ∂ − + 2.303θf (DLH [H3 O ]L + DLC [HCO3 ]L ) + 2S1 =− ∂t ∂x ∂x (4.42) where [HS] is the concentration of titratable soil acid. Applying Equation (4.39) to the structured-soil system, and ignoring the slow oxidation of Fe2+ in solution, gives S1 = ρkS [O2 ]L [Fe2+ ]S

(4.43)

0

10

20

30

40

50

60

70

0

2

4

6

8

3 days 6 days 9 days

10

120 100 9 days 80 60 40 20 3 days 0 0 2 4 6 8 10 mm

0 2 4 6 8 10 Distance from surface exposed to O2 (mm)

0

25

50

75

100

125

[Fetotal] (mmol kg−1)

[Fe(III)] (mmol kg−1)

4.0

4.5

5.0

5.5

6.0

6.5

(c) 7.0

0

2

4

6

8

10

Figure 4.17 Profiles of (a) Fe(II), (b) Fe(III) and (c) pH in columns of reduced soil exposed to O2 at one end for different times. Points are experimentally measured; lines are predicted using the model described in the text with independently estimated parameter values (Kirk and Solivas, 1994). Reproduced by permission of Blackwell Publishing

[Fe(II)] (mmol kg−1)

(b)

pH

(a) 80

132

133

Oxidation of Reduced Soil

where ρ is the soil bulk density. The rate of microbial O2 consumption is described by a Michaelis–Menten type equation: S2 = ρvmax [O2 ]L /(KM + [O2 ]L )

(4.44)

Kirk and Solivas (1994) measured profiles of Fe(II) and Fe(III) concentrations and pH in columns of reduced soil exposed to O2 at one end and compared the results with the predictions of the model using independently measured parameter values. The agreement between the observed and calculated results, shown in Figure 4.17, is good. The measured profiles of [Fe(II)] (Figure 4.17a) are scattered, probably because of the spatial variability inherent in soil reduction and the clustering of microbes around favourable microsites. There was much less scatter in the Fe(OH)3 and pH profiles which are the result of abiotic reactions. The zone of Fe(II) depletion extends further than the zone of Fe(OH)3 accumulation, as expected because Fe2+ is mobile but Fe(OH)3 is not. As a result, Fe(OH)3 accumulated in the oxidation zone close to the source of O2 , as shown in the inset in Figure 4.17(b). The good agreement between observed and calculated results and the fact that the model contains no arbitrary fitting parameters show that the important processes are well understood and that the model provides a satisfactory description of the system. It can therefore be used to explore other conditions through a sensitivity analysis (Figure 4.18). The figure shows that over the range of parameter values expected for submerged soils, substantial amounts of iron are transferred towards the O2 -exposed surface leading to a well-defined zone of Fe(OH)3 accumulation. For a given soil Fe(II) content, the accumulation is sensitive to the soil Fe2+ buffer power, the oxidation rate constant and the soil bulk density. The fall in pH in the oxidation zone is sensitive to the initial soil pH, the soil pH buffer power, and the partial pressure of CO2 . By contrast if the soil dries to any extent resulting in partially air-filled pores, the penetration of O2 increases dramatically: for an air-space of just 1 % of total 5.5 5.0

4 [Fe]

3 r 2 1 0 0.01

1

10

100

PCO2

r 4.5 [H+] 4.0 3.5

ks

0.1

bHs

3.0 0.01

ks 0.1

1

10

100

Spread oxidation front (cm)

5

0.14 [Fe]

bFe pH at soil surface

Fe transferred (mol cm−2)

6

0.12

ks

0.10 r 0.08

[Fe]

0.06 0.04 0.02 0.01

0.1

1

10

100

Multiple of standard parameter value

Figure 4.18 Sensitivity of the model used for the calculations in Figure 4.17 to its parameters: [Fe] is the initial concentration of mobile Fe2+ , bFe is the soil Fe2+ buffer power, bHS is the soil pH buffer power, kS is the Fe2+ oxidation rate constant and ρ is the soil bulk density. Standard values as for calculations in Figure 4.17

134

Reduction and Oxidation

porosity, Kirk et al. (1990b) calculated a four-fold increase in the penetration of O2 . There is then little accumulation of iron in the oxidation zone because the O2 diffuses so much faster than the Fe2+ that almost no Fe2+ can move towards the oxidation front before it is oxidized. The generation of acidity is correspondingly dispersed through the soil. These conclusions are discussed further in Chapter 6 in relation to the rhizospheres of wetland plants.

5 Biological Processes in the Soil and Floodwater

The soil and floodwater in wetlands are busy with life, and this drives the biogeochemistry. The remarkable long-term productivity of wetland rice systems depends on the fixation of carbon and nitrogen from the atmosphere by organisms in the soil and water, for which conditions are optimal. For example, in a long-term experiment at the International Rice Research Institute in the Philippines in which three crops of rice have been grown each year for 30 years without additions of fertilizers or manures and with complete removal of the rice straw, grain yields have remained nearly constant at 3 to 3.5 t ha−1 per crop or a total of 9 to 10 t ha−1 per year (Dobermann et al., 2000). No other intensive agricultural system without artificial inputs of nutrients comes close to this level of productivity. The accumulation of nitrogen by crops in this experiment has remained constant at about 50 kg N ha−1 per crop, largely due to additions from biological fixation in the floodwater and floodwater–soil interface (Ladha et al., 2000). Comparable rates of nitrogen fixation are attained in other fluxial wetland systems (Table 1.5). This chapter describes the important micro- and macrobiological processes in submerged soil and the overlying floodwater. Processes in plants and their rhizospheres are discussed in Chapter 6. The microbiological processes are discussed first and then the additional complexities caused by macrobiological processes and the particular ecology of the floodwater–soil system. 5.1 MICROBIOLOGICAL PROCESSES Descending through the soil from the floodwater there is a gradient of redox potential and a sequence of zones characterized by progressively more-reduced electron acceptors. Figure 5.1 shows hypothetical concentration profiles of redox species with depth. At sufficient depth the only electron acceptors are CO2 and H+ , and this zone is dominated by fermentation and methanogenesis. At intermediate depths there are successive zones of sulfate reduction, iron reduction, manganese reduction and denitrification. The microbes mediating these processes are largely prokaryotic; populations of fungi and other eukaryotes that are important in digesting organic matter under aerobic conditions are much less significant in anaerobic soil. The Biogeochemistry of Submerged Soils Guy Kirk  2004 John Wiley & Sons, Ltd ISBN: 0-470-86301-3

136

Biological Processes in the Soil and Floodwater Concentration O2 NO3−

NH4+

Depth

Mn2+ Fe2+

CH4

Figure 5.1 Indicative concentration profiles of redox species with depth in submerged soil

5.1.1 PROCESSES INVOLVED IN SEQUENTIAL REDUCTION The sequence of reactions by which organic matter is oxidized following submergence loosely follows the predictions of thermodynamics—i.e. in the order of decreasing free energy change—as described in Chapter 4. However, rates of reduction vary greatly between soils and there are complicated interactions between the microbial processes involved. Hence it is difficult to predict a priori, for example, how long after submergence a given soil will become methanogenic and what the rate of methane production will be. The free energy change for a particular redox reaction varies with pe, pH, and the concentrations of reductants and oxidants according to Equation (4.26):
 (Ox2 )(Red1 ) 0 0
G = −2.303RT n(pe1 − pe2 ) − log (Ox1 )(Red2 ) In this equation, the value of (Red) is a function of the nature of the reductant, its solubility, the crystallinity of solid phases containing it, effects of solubilizing agents, transport limitations, and other factors. Likewise the value of (Ox) is a function of various factors. As discussed in the previous chapter, most redox reactions are very slow and the prevailing conditions are therefore sensitive to catalysis. Three types of catalysis are involved: • Abiotic, for example by adsorption of reactants onto mineral surfaces, distinguished from biotic catalysis by the absence of a temperature optimum.

Microbiological Processes

137

Abiotic catalysis is generally less important than biotic but may be important. Examples are Mn(III,IV) and Fe(III) reduction by microbial metabolites, and Fe(II) oxidation which is catalysed by sorption onto soil particles. • Abiontic, involving free extracellular enzymes or solubilizing agents, enzymes bound to soil surfaces, enzymes within dead or non-proliferating cells, or enzymes associated with dead cell fragments. Extracellular enzymes are important in the initial stages of organic matter oxidation, in which polysaccharides and proteins are hydrolysed to soluble compounds that can be absorbed by microbial cells and further oxidized in biotic processes. • Above all, biotic catalysis by microbes is important. Biotic catalysis is complicated. Different communities of microbes deal with different parts of the sequence of processes degrading organic matter. Anaerobic decomposition involving organic electron acceptors (i.e. fermentation) generally occurs concurrently with respiration involving inorganic electron acceptors, and both produce intermediates that act as both oxidants and reductants. There are often syntrophic relationships between microbes in which the metabolisms of two or more organisms are linked and mutually beneficial. For example, in methanogenesis, oxidation of fatty and amino acids to H2 , CO2 and acetate is endergonic under standard conditions (i.e. PH2 = 1 atm), but a sufficiently small concentration of H2 is maintained locally by methanogens that utilize H2 (Conrad et al., 1986; Zehnder and Stumm, 1988; Krylova and Conrad, 1998). Likewise there are antagonisms between microbes, for example where one microbe maintains the concentration of a substrate below the threshold of a competitor, such as in the inhibition of methanogens by SO4 2− reducers competing for H2 (Achtnich et al., 1995). There are also specific inhibitory effects through particular metabolites, such as in the inhibition of methanogens by denitrifiers (Roy and Conrad, 1999). Hence the initial microbial populations, growth rates and community structures may all be important in the overall course of reduction. The main pathways of organic matter oxidation in anaerobic soil are as follows. In the initial stages, fermenting bacteria excrete extracellular enzymes that hydrolyse polysaccharides and proteins to soluble compounds. These may then be absorbed by microbial cells and converted to alcohols, fatty acids and H2 . If inorganic electron acceptors are available, the alcohols and fatty acids are completely oxidized to CO2 in sequential reduction reactions. If inorganic electron acceptors are not available—whether because they have been exhausted or because they are otherwise inaccessible—communities of fermenting bacteria decompose the alcohols and fatty acids to acetate, H2 and CO2 . These then serve as substrates for methanogenic archaea. Sugar monomers may also be directly converted to acetate by homacetogenic bacteria. Likewise proteins are hydrolysed to amino acids by extracellular enzymes, and the amino acids then ultimately oxidized to acetate, H2 , NH4 + and CO2 . Figure 5.2 shows the sequential reduction of inorganic electron acceptors and production of CO2 , CH4 and intermediaries in two representative soils from a

Acetate (µmol g−1)

0

2

4

6

8

10

12

14

0

0

40

60

80

100 120

Time after submergence (days)

20

NO3−

Fe(II)

0.0

0.2

0.4

0.6

0.8

1.0

1.2

1.4

−100

0

0

1

2

3

4

5

5.0

5.5

6.0

6.5

7.0

7.5

275 250 225 200 175 150 125 100 75 50 25 0

0

200

400

600

800

0

2

4

6

8

10

12

14

16

0

40

60

80

100 120

Fe(II)

EH

pH

Time after submergence (days)

20

NO3−

SO42−

Acetate

CH4

H2

CO2

0.0

0.5

1.0

1.5

2.0

−100

0

100

200

300

400

500

600

0

1

2

3

4

5

0

1

2

3

4

5

5.5

6.0

6.5

7.0

7.5

Figure 5.2 Sequential reduction of electron acceptors and accumulation of CO2 and CH4 in two rice soils. The soils were submerged and incubated at 30 ◦ C in sealed bottles (Yao et al., 1999). Reproduced with kind permission of Kluwer Academic Publishers

0

25

50

75

100

125

150

175

200

EH

100

400

200

200

600

pH

400

500

0

10

20

30

40

300

SO42−

Acetate

H2

CH4

CO2

50

800

1000

1200

CH4 or CO2 (kPa)

Acid-soluble Fe(II) (µmol g−1)

H2 (kPa)

EH (mV) SO42− (µmol g−1)

16

pH NO3− (µmol g−1)

CH4 or CO2 (kPa) Acetate (µmol g−1) Acid-soluble Fe(II) (µmol g−1)

H2 (kPa)

EH (mV) SO42− (µmol g−1)

Soil No. 14 (low labile C, high Fe)

pH NO3− (µmol g−1)

Soil No. 8 (high labile C, low Fe)

138

Microbiological Processes

139

sample of 16 rice soils studied by Yao et al. (1999). Three distinct phases can be distinguished: (1) an initial reduction phase lasting 19–75 days in the 16 soils, during which most of the inorganic electron acceptors are depleted and the rate of CO2 production, given by the slope of the CO2 accumulation line in the figure, is maximal; (2) a methanogenic phase starting after 2–87 days and lasting 38–68 days, during which the rate of CH4 production is maximal; and (3) a pseudo steady-state phase during which rates of CH4 and CO2 production and concentrations of H2 and acetate are roughly constant. The line of H2 accumulation in the figure is informative because H2 is turned over rapidly as it is produced in fermentation and consumed in Fe(III) and SO4 2− reduction and methanogenesis. Hence there are peaks in H2 pressure in the early stages of Fe(III) and SO4 2− reduction and again at the transition from Fe(III) and SO4 2− reduction to methanogenesis. Because consumption tends to increase with the concentration of H2 but production is independent of it, there is a point at which consumption equals production, characterized by H2 concentrations in the nM range. Acetate is also produced in fermentation and consumed in methanogenesis, but its turnover is slower and larger concentrations build up. Figure 5.3 compares the quantities of electrons consumed in reduction of inorganic electron acceptors and methanogenesis in the 16 soils with those donated in the oxidation of organic matter to CO2 . At the end of the initial reduction phase, the former exceeded the latter in nine of the soils, probably in part because CO2 was precipitated in carbonates and in part because some of the organic carbon was converted to forms more oxidized than that in CO2 . However by the end of the incubation the electron balance was zero in all but three of the soils. At the end of the incubation, only 6–17 % of the organic carbon in the soils was released as gases: 61–100 % as CO2 , 500 µm diameter)

Amphipods, copepods

Chironomid insect larvae, gastropods, isopods

Nematodes, amphipods, copepods, oligochaetes

Crayfish, clams, oligochaetes, gastropods, isopods, midge larvae Oligochaetes including earthworms, crayfish

Nematodes, oligochaetes, ‘terrestrial’ invertebrates (mitesacari, springtailscollembola)

Nematodes, amphipods, oligochaetes Nematodes, amphipods, oligochaetes Nematodes, amphipods, oligochaetes Nematodes, amphipods, copepods, oligochaetes, polychaetes

Oligochaetes, polychaetes, midge larvae Oligochaetes, fiddler crabs, snails, mussels Oligochaetes, fiddler crabs, mud crabs, periwinkles, snails, mussels, oysters, clams Fiddler crabs, oysters, barnacles

Source: adapted from Craft (2001). Reproduced by permission of Lewis Publishers.

160

Biological Processes in the Soil and Floodwater

Table 5.5 Invertebrates in ricefield soil and floodwater Densities (number m−2 )

Microcrustacea Ostracods

Copepods Cladocerans Insect larvae Chironomids Mosquitoes Molluscs Snails

Oligochaetes Tubificids

Min.

Max.

Mean

0

98 000

6000

0 0

40 000 33 000

33 000 900

0

10 000

600

0

7000

170

0

1000

200

0

40 000

10 000

Comments

       

Stimulated by factors that increase primary production, such as N and P fertilizer. Filter bacteria and algae from water

 

Feed on epipelic algae at soil surface and on floating algae

      



Inhibited by high acidity, N fertilizer and pesticides; stimulated by high organic matter. Graze on epipelic and floating algae, and on algae epiphytic on plant stems Stimulated by factors that increase primary production and bacterial decomposers; inhibited by high soil bulk density

Source: Roger (1996), Simpson et al. (1993a, 1994a,b).

organic matter and applications of N fertilizer, presumably through the effects of these on primary production and bacterial decomposers on which oligochaetes feed. Populations of insect larvae and molluscs are successional over the season following the cycles of algal populations. But no general trends have been established for the dynamics of oligochaetes, though this may reflect the paucity of data (Roger, 1996). The effects of macrofauna on the soil biogeochemistry can be summarized (Aller, 1994): • manipulation of particles: exposure of substrate resulting in increased decomposition; • grazing: consumption of microbes, stimulation of microbial growth, increased mineralization; • excretion of substrate and nutrients: stimulation of microbial growth, increased mineralization;

Macrobiological Processes

161

• construction of burrows: synthesis of refractory or inhibitory structural products; • irrigation of burrows: increased transfer of soluble oxidants and nutrients, increased re-oxidation and mineralization; • transport of particles: transfer between major redox zones, increased re-oxidation and mineralization. Bioturbation Oligochaetes feed with their heads downward in the burrow and posterior ends upward in the water. They ingest fine soil particles, extract carbon and minerals, and deposit residues in faeces on the soil surface. The faeces may subsequently fall into the burrow and be mixed. The net effect is a loosening and mixing of the soil to depth. The burrows of the species found in ricefields may be several centimetres deep and a millimetre or so in diameter. Deeper and wider burrows are formed by species found in other wetlands (Table 5.4). Once the burrows are constructed, the worms tend to remain in them and maintain a supply of oxygen from the overlying water by waving their posteriors in the water and moving their bodies in a peristaltic motion. Thereby the water in the burrows is mixed with the overlying water and solutes diffusing into a burrow are rapidly transferred to the surface and vice versa. The calculations in Section 2.4 show the great sensitivity of solute transport and mixing to the geometry, density and activity of the burrowing animals. The effects of oligochaetes on the soil or sediment depend on the particular circumstances. Limnologists and oceanographers consider oligochaetes to be agents of aeration, increasing the depth of the oxidized layer and stimulating mineralization and nitrification–denitrification (Fry, 1982; Aller, 1994). For example, Davis (1974) found that the oxidized layer (EH > 200 mV) in profundal lake sediments was increased by 0.3–1.6 cm by tubificid populations of 800 m−2 . By contrast in ricefields, where primary production and the amounts of organic matter in the floodwater may be much greater, the effect can be to enhance the incorporation of organic matter into the soil and so to make the surface soil on average more reduced, in spite of oxygenation of the solution in the burrows. Kikuchi and colleagues (Kikuchi et al., 1975; Kikuchi and Kurihara, 1977, 1982) found that with realistic densities of tubificids and organic matter in ricefields, the oxidized layer at the soil surface disappeared altogether. They found that weed growth was diminished because seeds were moved to a depth at which the O2 concentration was too low for germination, and as a consequence oxygenation of the soil by weeds decreased and populations of aerobes in the soil decreased and anaerobes increased. The concentrations of NH4 + , ortho-P and acid soluble Fe in the floodwater increased and the concentration of NO2 − + NO3 − decreased (Figure 5.11). In practice redox conditions in the burrows will oscillate as the oxygenation of the floodwater varies over the diurnal cycle. Aller (1994) found in a wide range of organic matter-rich sediments containing burrowing invertebrates that

162

Biological Processes in the Soil and Floodwater

acid soluble Fe (µM)

ortho P (µM)

NO3− + NO2− (µM)

NH4+ (µM)

600 with tubificids without

400 200 0 500 400 300 200 100 0 10 8 6 4 2 0 750 600 450 300 150 0 150

175

200

225

250

Day of year

Figure 5.11 Effects of tubificids on concentrations of N species, P and Fe in the floodwater of unplanted microplots in ricefields. Species B. sowerbyi, density 1000 m−2 (Kikuchi and Kurihara, 1982). Reproduced with kind permission of Kluwer Academic Publishers

solid particles constantly cycled between oxic and anoxic zones but typically spent 10- to 100-times longer under anoxic than oxic conditions. Cyclic redox patterns were also common within individual burrows and were accompanied by rapid switching of metabolic processes. Even brief, periodic re-exposure of organic matter to O2 resulted in more complete decomposition than under constant conditions or unidirectional redox change. Redox oscillation apparently results initially in net remineralization of existing microbial biomass followed by stimulated renewed synthesis. Aller (1994) found that some properties, such as the accumulation of P in the sediment, were comparable under fully oxic and oscillating redox conditions but differed under continuously anoxic conditions. This is another mechanism by which the operation of the floodwater–soil system as a whole is not a simple sum of its component parts.

Is Biodiversity Important?

163

5.3 IS BIODIVERSITY IMPORTANT? Submerged and non-submerged soils have huge biological diversity and contain orders of magnitude more biological species than above ground habitats and aquatic systems (Liesack et al., 2000; Conrad and Frenzel, 2002; Usher et al., 2004). The origins of this diversity are in the physical and chemical heterogeneity of soils at micro- and macro-scales, and intense competition for substrate. In submerged soils there are extreme gradients of redox potential from oxic to anoxic zones due to the slow transport of O2 through the soil, and gradients of substrate from rich to poor zones due to the non-uniform distribution of plant debris and root exudates. Heterogeneity also arises from soil physical structure and the labyrinthine network of soil pores which constrain the movements of both organisms and substrates. Although submerged soils have weak macro-structure, especially rice soils that have been deliberately puddled, they retain considerable micro-structure in water-stable micro-aggregates and a corresponding network of pores. These factors in combination result in a near infinite combination of opportunities and constraints for different organisms. But does all this biodiversity have any consequences for soil processes at the macro-scale? This is a seemingly straightforward question, but the answer has been surprisingly elusive. Progress has been hampered by the absence of suitable experimental methods for analysing biological diversity and its relation to soil functions. Three types of method are used (Ritz, 2004): (1) genotypic analysis, which assesses the basic genetic information about the community of microbes present; (2) phenotypic analysis, which assesses the expression of the genetic information, i.e. the living form of the microbial community; and (3) functional analysis, which assesses the processes that the microbial community is actively or potentially engaged in. The greatest diversity is revealed in genotypic analysis, but there is a corresponding lack of discrimination. Analyses of soil DNA often do not show clear differences between soils from widely differing environments, including in submerged soils (Liesack et al., 2000; Reichardt et al., 2000). Phenotypic analysis, such as by assaying membrane-bound phospholipids from living microbes, is more discriminatory, and there is now good evidence for phenotypic ‘signatures’ in soil microbial communities, modified by the environment the microbes are operating in, including for submerged soils (Reichardt et al., 1996). In functional analysis, the actual or potential activities of the microbial community are measured. Techniques for this have been developed based on the ability of soil communities to utilize different C-containing compounds using Biolog plates. The results often match phenotypic analysis and the expected effects of soil and environmental differences. However the method is biased towards those microbes that thrive under the particular conditions of the assay in vitro (Preston-Mafham

164

Biological Processes in the Soil and Floodwater

et al., 2002). A solution is to measure the utilization of substrates added directly to soil, and practical methods for doing this are being developed (Degens and Harris, 1997). Through such techniques evidence is emerging for the importance of biodiversity for macro-scale soil processes (Usher et al., 2004). Though there is evidently a great deal of redundancy in most microbial populations, there are threshold levels of biodiversity below which important soil functions are impaired. For example, the decomposition of recalcitrant organic matter can only be achieved by consortia of organisms operating together, such as in the anaerobic decomposition of organic matter in submerged soils in sequential reduction reactions mediated by microbes (Section 5.1.1). Also the growth and activity of individual organisms is necessarily constrained by the nature of the prevailing community, which is important, for example in the persistence of rare organisms in the soil and management of soil-borne diseases. Given that biodiversity is important, it is important to understand how soil management affects it. But as yet there is not much information on this. It has been assumed that intensification of rice production and more widespread use of fertilizers and pesticides in the past few decades will have diminished the diversity of microbes and invertebrates in ricefields compared with those under traditional practices. Roger et al. (1991) compared the diversity of arthropods in farmers’ fields in the Philippines and at the International Rice Research Institute, and found the greatest diversity in fields at the Institute, where there has been heavy use of fertilizers and pesticides for many years, and the least in fields in the Ifugao rice terraces at Banaue, where there has been little use of fertilizers and pesticides. This goes against the often hypothesized trend of intensification reducing biodiversity. Simpson et al. (1993b, 1994a, b) measured the effects of fertilizer and pesticides on populations of algae and invertebrates in ricefields, and found complicated interactions. Whereas N fertilizer inhibits N2 fixation by cyanobacteria, P fertilizer stimulates it, and the overall productivity of the floodwater is generally increased by fertilization. Likewise pesticides have various effects. Part of the community of organisms responsible for mineralizing organic matter may be killed by pesticide, but the subsequent collapse of predators may allow other mineralizing organisms to bloom. Several insecticides reduce the numbers of ostracods that graze on N2 -fixing cyanobacteria, and so N2 fixation is enhanced.

6 Processes in Roots and the Rhizosphere

Though wetland plants have the advantage of an assured water supply, they must contend with various difficulties that their dryland counterparts are largely spared. First, because of the very slow diffusion of respiratory gases through water and submerged soil compared with dryland soil, the roots must aerate themselves internally by forming internal gas channels to the air above. Second, the roots must exclude toxic products of anaerobic metabolism in the soil, such as organic acids and ferrous iron, or tolerate large concentrations of these toxins internally. Third, they must contend with the altered forms and solubilities of nutrients in the soil under anaerobic conditions, for example the predominance of ammonium rather than nitrate as the plant-available form of nitrogen. That wetland plants are capable of surmounting these difficulties is shown by the great productivity and biodiversity of wetland systems. This chapter discusses the various processes and mechanisms involved in this. 6.1 EFFECTS OF ANOXIA AND ANAEROBICITY ON PLANT ROOTS Generally, in plant cells well supplied with O2 , energy is provided for growth and metabolism by the oxidation of glucose in the three stages shown in Figure 6.1: (1) glycolysis, in which 1 mol of glucose is converted to 2 mol of pyruvate yielding 2 mol of ATP (the main form in which energy is transported and utilized in plants) and 2 mol of NADH2 (reduced NAD which acts as a universal reducing agent in non-green plant tissues); (2) the Krebs cycle, in which 1 mol of pyruvate is completely oxidized to CO2 yielding 1 mol of ATP and 5 mol of NADH2 ; and (3) the mitochondrial cytochrome chain, in which 1 mol of NADH2 generates 3 mol of ATP. The net result is that complete aerobic respiration of 1 mol of glucose yields 38 mol of ATP. However in the absence of O2 , anaerobic glycolysis—fermentation—produces only 2 mol of ATP per mol of glucose consumed. In the absence of O2 the mitochondrial cytochrome chain ceases to operate and as a result NADH2 accumulates The Biogeochemistry of Submerged Soils Guy Kirk  2004 John Wiley & Sons, Ltd ISBN: 0-470-86301-3

166

Processes in Roots and the Rhizosphere

(a)

2 ATP

2 ADP

Glucose

2 (1,3-Diphosphoglycerate) 2 ADP

2 PGA

2 NAD

2NADH2

2 Ethanol

2 ATP

3 2 Acetaldehyde

2 PEP 2 ADP

1

2 ATP 2 Pyruvate

2 Lactate 2 2 CO2

Krebs cycle (b)

Isocitrate

Pyruvate NAD

CO2 a-Ketoglutarate NAD, ADP

Citrate NADH2 CO2 Acetyl-CoA

Oxaloacetate

NADH2

ATP e

ADP FP

e

FPH2

Malate

NAD ADP

NADH2, ATP CO2 Succinate FP Fumarate

NADH2

(c)

NAD NADH2

cytb

ATP e

cytc

ADP e

cyta

ATP e

O2

cyta3 OH−

Figure 6.1 Pathways involved in glucose oxidation by plant cells: (a) glycolysis, (b) Krebs cycle, (c) mitochondrial cytochrome chain. Under anoxic conditions, Reactions 1, 2 and 3 of glycolysis are catalysed by lactate dehydrogenase, pyruvate decarboxylase and alcohol dehydrogenase, respectively. ATP and ADP, adenosine tri- and diphosphate; NAD and NADH2 , oxidized and reduced forms of nicotinamide adenine dinucleotide; PGA, phosphoglyceraldehyde; PEP, phosphoenolpyruvate; Acetyl-CoA, acetyl coenzyme A; FP, flavoprotein; cyt, cytochrome; ε, electron. (Modified from Fitter and Hay, 2002). Reprinted with permission from Elsevier

and the Krebs cycle is suppressed. This leads to an accumulation of acetaldehyde—the first end-product of fermentation (Figure 6.1a); synthesis of alcohol dehydrogenase catalysing the conversion of acetaldehyde to ethanol; and consumption of NADH2 as acetaldehyde is reduced to ethanol and hence continuing production of ATP and pyruvate. So fermentation can continue to generate ATP

Effects of Anoxia and Anaerobicity on Plant Roots

167

for as long as carbohydrate reserves last. However, because the efficiency of this process is much less than the efficiency of aerobic respiration—only 2 mol of ATP are produced per mol of glucose consumed compared with 38—the rate of fermentation must increase sharply under anoxia if the cell energy supply is to be maintained. This can lead to rapid exhaustion of plant reserves under prolonged anoxia. In addition anoxic cells must contend with toxic products of fermentation, particularly ethanol and lactate. Hence the plant will go to some lengths to avoid anoxia in its active tissues.

6.1.1 ADAPTATIONS TO ANOXIA The most important adaptation plants make to anoxic soil conditions is the development of highly porous tissue in the root cortex called aerenchyma (Figure 6.2) (Jackson and Armstrong, 1999). The development of aerenchyma may occur both through closely regulated separation and expansion of cells, or, more usually, through programmed cell death, also under tight regulation in response to external stimuli. The result is a continuous pathway of gas channels between the base of the root and the tip. This both permits gas transport between the plant’s aerial parts and respiring root tissues, and lessens the amount of respiring tissue per unit root volume. In addition, the root wall layers become partially suberized along part of the root length, resulting in decreased permeability to gases and hence less loss of O2 to the anaerobic soil outside. The mechanisms by which the aerenchyma remains gas-filled rather than waterfilled are not fully understood but appear to involve metabolic control (Raven, 1996). The gas-filled state is favoured by inward gradients of water potential created by evapo-transpiration and by barriers to water movement in the apoplasm such as exodermis (van Noordwijk and Brouwer, 1993). Thus the root acts as a moderately gas-tight pipe conveying O2 down from the shoots to the elongating and actively respiring tip and venting CO2 and other respiratory gases in the opposite direction. Figure 6.3 shows changes in root porosity and respiration rate along the length of maize roots grown in anoxic media. Metabolic adaptations in the root to provide alternative respiratory pathways are far less important. Where these do occur, they are only of short-term use. Indeed, in plants that tolerate prolonged soil submergence, root tissues are often particularly sensitive to anoxia (Vartapetian and Jackson, 1997). Without morphological adaptations and a continuous supply of O2 to the root tip, survival is limited. That said, some rice genotypes will survive several days of anoxia resulting from complete submergence of the plant following flash flooding, which is a widespread phenomenon in rainfed rice systems. The tolerance appears to depend on (a) the water being sufficiently clear and with a sufficient dissolved CO2 content that the plants can continue to photosynthesize and produce carbohydrates; (b) cessation of growth so as to preserve carbohydrates for maintenance processes; and (c) increased alcoholic fermentation to maintain glycolysis, NAD

168

Processes in Roots and the Rhizosphere (a)

(b) P

E CC

SC

CC

0.1 mm

RH (c)

0.1 mm (d)

CC P

P

AE

branch roots aerenchyma

central cylinder

E 0.1 mm

SC

0.1 mm

crack

br

an

ch

ro

ots

(e)

aerenchyma 0.1 mm

Figure 6.2 Cross-sections of primary rice roots. (a) Radial section close to tip showing intercellular spaces (I), central cylinder (CC), and rhizodermis (RH). (b) and (c) Radial sections of younger (39 days) and older (72 days) basal parts showing exodermis (E), schlerenchymatous cylinder (SC), parenchymatous or cortical cells (P) and aerenchyma (AE). (d) and (e) Axial sections of mature root (72 days) showing break through of lateral roots (Butterbach-Bahl et al., 2000). Reproduced by permission of verlag

recycling and ATP synthesis (Setter et al., 1997). However these adaptations would not serve a vigorously growing root system in normal circumstances. Transport of gases through the aerenchyma may occur by diffusion and, where pressure gradients develop, by convection. Pressurized flow is important in wetland plants with root systems permitting a throughflow of gases, but is insignificant in other plants (Beckett et al., 1988; Skelton and Alloway,

169

Effects of Anoxia and Anaerobicity on Plant Roots 60

250

Aerenchyma

200

40

150

Stele

Wall (estimated)

30

Total

100 50 0

50

Cortex 0

20 10

Cortex wall

100

200

Aerenchyma (%)

Respiration rate (ng O2 mL−1 s−1)

300

300

400

0 500

Distance from apex (mm)

Figure 6.3 Aerenchyma development and changes in respiration rate along the length of maize roots grown in anoxic media (adapted from Armstrong et al., 1991a). Reproduced by permission of Backhuys publishers

1996). In throughflow systems atmospheric gases are driven or sucked into the above-ground parts of the plant and then vented from some other point on the above-ground parts as an O2 -depleted and CO2 -enriched exhaust. There are various possible sources of positive pressure—e.g. humidity-induced diffusion and thermal transpiration—and of negative pressure—e.g. wind (Venturi forces), the greater solubility of CO2 than O2 (140-fold at 25 ◦ C and pH 7), differences in gas velocities, and thermo-osmosis (references in Jackson and Armstrong, 1999). Resistance to pressure flow is inversely proportional to the fourth power of the radius of the conducting vessel, and so large pore-diameters in the diaphragm partitions of leaf sheath, stem and rhizome are an essential prerequisite for efficient pressurized flow. A well known example of a pressurized flow system is the water lily (Dacey, 1980, 1981). Pressurized flow could in principle occur in a non-throughflow root system, such as that of rice, driven by dissolution of respiratory CO2 produced from gaseous O2 . However, Beckett et al. (1988) have shown that convection by this means will always be subordinate to diffusion in non-throughflow systems and will only ever have a minor effect. Hence diffusion is the principle means of gas transport. The effectiveness of the internal O2 transport by diffusion or convection depends on the physical resistance to movement and on the O2 demand. The physical resistance is a function of the cross-sectional area for transport, the tortuosity of the pore space, and the path length. The O2 demand is a function of rates of respiration in root tissues and rates of loss of O2 to the soil where it is consumed in chemical and microbial reactions. The O2 budget of the root therefore depends on the simultaneous operation of several linked processes and these have been analysed by mathematical modelling (reviewed by Armstrong

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Processes in Roots and the Rhizosphere

et al., 1991b, 2000). In the following section I describe the model developed by Armstrong and Beckett (1987). This accounts for the most important processes within the root and has been corroborated for various wetland species using measurements of O2 gradients within roots with microelectrodes.

6.1.2 ARMSTRONG AND BECKETT’S MODEL OF ROOT AERATION To summarize, the main factors influencing the O2 budget of a non-throughflow root in anoxic soil are as follows. (1) The extent of aerenchyma development by the degradation of the primary root cortex. (2) Rates of respiration in different root tissues. The formation of aerenchyma decreases the respiratory O2 demand per unit root volume because there is less respiring root tissue. Also, some plants can tolerate a degree of anoxia in parts of the root, which substantially reduces the O2 demand per unit root volume. (3) The permeability of the root wall to gases. Sub-apical parts of the root can have permeabilities several orders of magnitude smaller than those in the region of the tip. (4) The proportion of fine lateral roots branching off the primary root. Having high surface area to volume ratios, laterals tend to be O2 -leaky. For simplicity, the effects of lateral roots are not dealt with explicitly in Armstrong and Beckett’s model, but they are dealt with in Section 6.2. In the model, the internal structure of the root is described as three concentric cylinders corresponding to the central stele, the cortex and the wall layers. Diffusivities and respiration rates differ in the different tissues. The model allows for the axial diffusion of O2 through the cortical gas spaces, radial diffusion into the root tissues, and simultaneous consumption in respiration and loss to the soil. A steady state is assumed, in which the flux of O2 across the root base equals the net consumption in root respiration and loss to the soil. This is realistic because root elongation is in general slow compared with gas transport. The basic equation is

 d[O2 ]G d[O2 ]L 1 d d (6.1) DG θG fG + rDL − Rroot − Rsoil = 0 dz dz r dr dr where the first term represents axial diffusion through the cortical gas spaces, the second term radial diffusion through root tissues, and the third and fourth terms the rates of O2 consumption in tissue respiration and loss to the soil, respectively. Here DG and DL are the diffusion coefficients of O2 in air and water, respectively, θG is the gas content of root by volume, fG is the impedance factor for diffusion in the cortical gas spaces, r is the radial distance, z is the axial distance and [O2 ]G

Architecture of Wetland Plant Root Systems

171

and [O2 ]L are the concentrations of O2 in the gas space (mol per unit volume gas space) and in root tissue (mol per unit volume root), respectively. The boundary conditions for solving Equation (6.1) are: (a) at the root base, [O2 ]G is the ambient value in the atmosphere; and (b) at the root apex, [O2 ]G is the minimum value required for root respiration [≈ 30 µmol dm−3 (gas space)]. The equations are solved numerically. 6.2 ARCHITECTURE OF WETLAND PLANT ROOT SYSTEMS In dryland plants the size of the root system compared with the shoot system is generally governed by the plant’s water requirements except under quite severe nutrient deficiency (Tinker and Nye, 2000). However, in wetland plants in submerged soil, the free availability of water means that the size of the root system is more often likely to be governed by nutrient requirements. The length densities of wetland root systems may be comparable to those of dryland plants: length densities of rice roots are typically 20–30 cm cm−3 in the topsoil (Matsuo and Hoshikawa, 1993). A large proportion of the length may be as fine roots. In rice in submerged soil short fine laterals, 1–2 cm long and 0.1–0.2 mm in diameter, develop as branches along the primary roots once the primary roots are a few cm long. These are much less aerenchymatous than the primary roots (porosities of 1–2 % compared with ≤50 %) and they do not develop secondary thickenings in their walls to the same extent (Matsuo and Hoshikawa, 1993). They may themselves be branched producing up to sixth order laterals. They account for a small part of the root mass but the bulk of the external surface, and they are plumbed directly into the main water and solute transport vessels in the stele of the primary root (as can be seen in Figure 6.2). The structure of the rice root is therefore apparently dominated by the need for internal gas transport. On the face of it, this structure may conflict with the needs for efficient nutrient absorption (Kirk and Bouldin, 1991). The development of gas-impermeable layers in the root wall seems likely to impair the ability of those parts of the root to absorb nutrients, and the disintegration of the cortex might impair transport from the apoplasm to the main solute transport vessels in the stele, though these points are uncertain (Drew and Saker, 1986; Kronzucker et al., 1998a). It seems likely that the short fine lateral roots are responsible for the bulk of the nutrient absorption by the root system and compensate for any impairment of nutrient absorption by the primary roots as a result of adaptations for internal aeration. The question arises: what combination of fine laterals and aerenchymatous primary roots provides the greatest absorbing surface for a given root mass? Not having impermeable wall layers and having a large surface area to volume ratio, the laterals will leak O2 more rapidly than the adjacent primary root. A related question is therefore how the O2 budget of the root system is affected by the combination of primary roots and laterals. Armstrong et al. (1990, 1996) modelled O2 release from adventitious and lateral roots of the rhizomatous wetland

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species Phragmites australis, and found that for the appropriate combination of root types, properties and dimensions, and a large but realistic soil O2 demand, the ratio of O2 consumption in root respiration to that in loss to the soil was 13:1 for adventitious roots but 0.15:1, i.e. reversed, for laterals. Evidence for preferential loss of O2 from laterals in rice includes measurements of Fe oxide coatings on roots placed in deoxygenated agar containing Fe(II) (Trolldenier, 1988); changes in redox potential as roots grew across rows of Pt electrodes in anaerobic soil (Flessa and Fischer, 1993); and the abundance of methane oxidizing bacteria, which are obligate aerobes, along rice lateral roots in anaerobic soil (Gilbert et al., 1998). Although O2 leakage compromises the root’s internal aeration, some leakage is desirable for a number of purposes. These include oxidation of toxic products of anaerobic metabolism in submerged soil such as ferrous iron (van Raalte, 1944; Bouldin, 1966; van Mensvoort et al., 1985); nitrification of ammonium to nitrate, there being benefits in mixed nitrate–ammonium nutrition (Kronzucker et al., 1999, 2000); and mobilization of sparingly soluble nutrients such as P (Saleque and Kirk, 1995) and Zn (Kirk and Bajita, 1995) as a result of acidification due to iron oxidation and cation–anion intake imbalance.

6.2.1 MODEL OF ROOT AERATION VERSUS NUTRIENT ABSORPTION Kirk (2003) has developed a simple model to compare root requirements for aeration with those for efficient nutrient acquisition in rice. The main features of the rice root system are summarized in Figure 6.4. The model considers roots in the anoxic soil beneath the floodwater—soil interface, receiving their oxygen solely from the aerial parts of the plant. Structure of the Root System The distribution of primary roots beneath a hill of plants is approximately hemispherical with the individual roots randomly distributed with respect to the vertical and horizontal directions. Thus if there are N primary roots per hill, the length of primary roots per unit soil volume, LVP , at any distance r from the centre of the hill is N dN/dr = (6.2) LVP (r) = dV /dr 2πr 2 About each primary root there is a cylinder of laterals, increasing in density with distance from the root base (Figure 6.5). The laterals may develop up to sixth-order branches. A simple equation to describe this is: LVL (r) = LVL max

r2 (k + r)2

(6.3)

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Architecture of Wetland Plant Root Systems

Superficial roots in floodwater and oxic soil Floodwater Oxic surface soil Primary roots (with laterals) in anoxic soil Anoxic soil

Fine roots penetrating plough pan Plough pan Oxic subsoil

Figure 6.4 Root system of the rice plant (Kirk, 2003). Reproduced by permission of Blackwell Publishing

where LVL is the length density of laterals in the cylinder of soil occupied by them, k is a coefficient, equivalent to the distance at which LVL (r) = 0.25LVL max , and r0 < r ≤ rlat . If the cylinder has outer radius x and inner radius aP (i.e. the radius of the primary root), and x and aP are constant along the root length, then the total length density of primary and lateral roots at distance r from the centre of the hill is   r2 N 2 2 LV (r) = − a )L 1 + π(x (6.4) VL max P 2πr 2 (k + r)2 Equation (6.4) gives reasonable fits to measured profiles of LV with depth in the field. Structure of an Individual Root and its Laterals The porosity of the cortex, permeability of the root wall and the coverage of the root with laterals vary along the root length, with a much smaller porosity,

174

Processes in Roots and the Rhizosphere radius of hill

zone of laterals

r0

rlat r1

r2 rmax

zone of decreasing porosity zone of root tip

Figure 6.5 Idealized primary root and its cylinder of laterals. The parallel lines indicate the increasing length density of laterals along the primary root. The branching of the laterals is not represented (Kirk, 2003). Reproduced by permission of Blackwell Publishing

more-permeable wall and no laterals in the region of the tip. Where the laterals emerge from the primary root, there are generally cracks in the epidermis a few µm wide and apparently directly connected to the primary root aerenchyma (Butterbach-Bahl et al., 2000). It seems likely these will be important in gas transfer, though there are no direct measurements showing this. In practice leakage of O2 from the cracks and axial gradients of O2 within laterals will lead to gradients of O2 release along laterals. However, for the intended purpose of the model an elaborate treatment of these effects is not necessary; it is sufficient that the loss of O2 increases with the density of laterals and a constant leakage along the length of laterals is assumed. Figure 6.5 defines for the purposes of the model the distances at which the porosity and coverage with laterals change. It is assumed that, because of the changes in wall permeability along the root, nutrients are only absorbed by the primary root in the zones beyond the laterals (rlat < r < rmax ) and by the laterals. This is also the surface across which O2 leaks.

Architecture of Wetland Plant Root Systems

175

Transport and Consumption of O2 in the Roots and Losses to the Soil To avoid unduly complicating the model, radial diffusion within the root is not allowed for. Equation (6.1) therefore reduces to:
 d d[O2 ]G (6.5) DG θG fG − Rroot − Rsoil = 0 dr dr where r is the distance from the root bases (not the radial distance across the root as in Equation 6.1). Rroot at a particular distance along the root is the sum of the respiration in the primary root and in any laterals emerging from it. Hence, if the rate of respiration per unit root mass is Q, Rroot = ρ(1 − θG )Q +

ρ(1 − θG2 )QπaL2 π(x 2 − aP2 )LVL πaP2

(6.6)

Likewise Rsoil at a particular distance is the sum of the rates of loss from the primary root and from the laterals. Hence, if FO2 is the flux across unit root surface, Rsoil =

2πaL FO2 π(x 2 − aP2 )LVL 2πaP FO2 + πaP2 πaP2

(6.7)

It is assumed that the primary root wall is completely impermeable to O2 in the zone covered with laterals. In fact the root wall is not completely impermeable in this zone but the resulting flux is small compared with that from the rest of the root system and no serious error arises from ignoring it. It is also assumed that the flux from the laterals and the primary root in the zone beyond the laterals is constant. In fact the sink for O2 in the surrounding soil will vary in a complicated way with soil conditions and time, and there will be differences along the root length. However to some extent these differences cancel each other (Kirk, 2003) and the additional complexity involved in allowing for them is unjustified. The same boundary conditions apply as for Armstrong and Beckett’s model, and the equations are solved numerically.

Model Calculations Figure 6.6 shows results for a realistic set of standard parameter values. The maximum primary root length is 27.3 cm declining to 17.7 dm as the coverage with laterals increases from +200 mV. The results suggest the possibility of using management practices to maintain the redox potential in a range where both N2 O and CH4 emissions are low. 8.2.3 DIFFERENCES BETWEEN RICE PRODUCTION SYSTEMS Bronson et al. (1997a,b) made continuous measurements of CH4 and N2 O emissions from ricefields over a period that included two dry season and one wet season irrigated rice crops and the two intervening fallow periods. The soil was clayey and poorly drained. Figure 8.7 shows that during the growing seasons, N2 O fluxes were generally barely detectable although small emissions (≤ 3.5 mg N m2 day−1 ) occurred after N fertilizer applications. Methane fluxes, on the other hand, were substantial throughout the rice-growing seasons. The total emission of CH4 over the season decreased three- to four-fold when N was supplied as (NH4 )2 SO4 rather than urea at 200 kg N ha−1 , but emission of N2 O was 2.5-fold greater with (NH4 )2 SO4 . Mid-season drainage suppressed CH4 emission by ≤ 60 %, but markedly increased N2 O emissions. Figure 8.8 shows the results for the fallow periods. These lasted 5 to 11 weeks and were weedy. The soil was generally aerobic, and moderate amounts of NO3 − accumulated (7–20 kg N ha−1 ). Moderately high, continuous N2 O emissions occurred, apparently during nitrification of mineralized organic N in the topsoil and possibly also during denitrification in the wet subsoil. The flux of N2 O was greatest immediately after rainfall and after the field was flooded for rice at the end of the fallow, as a result of denitrification of accumulated NO3 − .

251

N2O flux (mg nm−2 day−1)

CH4 flux (mg cm−2 day−1)

Nitrogen Oxides

800

35 (a) straw

(b) urea 28

600

continuous flooding

drainage (arrows)

400

21 14

200

7 0

0

6

2.0 (c) straw

(d) urea

1.5

4

1.0 2

0.5 0.0 0

20

40

60

0 80 100 0 20 Days after transplanting

40

60

80

100

N2O flux (mg nm−2 day −1)

CH4 flux (mg cm−2 day−1)

Figure 8.7 Emissions of CH4 and N2 O during a rice crop with different water, straw and fertilizer managements. Single upward arrow = drainage; double downward arrow = flood irrigation (Bronson et al., 1997a). Reproduced by permission of Soil Sci. Soc. Am. 5

5 Geen manure Straw

4

4

3

3

2

2

1

1

0

0

80

80

60

60

40

40

20

20

0

0

6

12

18

24

0 30 36 0 6 Days after harvest

Urea Ammonium sulfate

12

18

24

30

36

Figure 8.8 Emissions of CH4 and N2 O during a fallow between rice crops. Single arrow = rainfall; double arrow = flood irrigation (Bronson et al., 1997b). Reproduced by permission of Soil Sci. Soc. Am.

252

Trace Gases

Little CH4 was emitted during the fallows. This study demonstrates that rice soils in the fallow periods can be significant sources of N2 O. A common cropping sequence in the rainfed lowlands is wet season rice followed by a dry season upland crop on residual soil moisture or supplemental irrigation, followed by a 60- to 70-day fallow during the dry-to-wet transition. Alternate soil wetting and drying in this system create particular difficulties for the conservation of nitrogen in the soil (Buresh et al., 1993a; George et al., 1993, 1994, 1995). Soil N mineralized and nitrified at the onset of rains in the fallow may be lost by leaching and by denitrification when the soil becomes submerged. Commonly high-value vegetable crops are grown in the dry season with heavy applications of fertilizers, leaving substantial amounts of residual nitrate in the soil. This situation leads to large losses of N before the wet season rice is established. Studies of N balances in an intensified rainfed lowland system of this sort in the Philippines have shown N losses of up to 550 kg ha−1 year−1 through nitrate leaching and denitrification (Tripathi et al., 1997).

8.3 AMMONIA 8.3.1 GLOBAL BUDGET Ammonia has a lifetime of only a few hours to a few days in the atmosphere. It and its reaction products are transported through the atmosphere and deposited on terrestrial surfaces elsewhere. It is the main gaseous alkaline species in the atmosphere and neutralizes a large part of the acid produced in oxidation of sulfur and nitrogen oxides, probably up to a half though its dry-deposition is much faster than that of NOx and SO2 (Dentener and Crutzen, 1994). Dry- and wet-deposition of ammonia contribute to soil acidification because 2 mol of H+ are produced in the nitrification of 1 mol of NH4 + . Also a large part of the ammonia deposited on moist forest soils may be re-emitted as N2 O (Section 8.2). Table 8.5 shows a global inventory of ammonia emissions compiled by Bouwman et al. (1997). The main sources are the excreta of domestic animals (40 %), use of nitrogen fertilizers (17 %), the oceans (15 %) and biomass burning (11 %). About half of the global emission comes from Asia, and 70 % is from food production. Europe, the Indian subcontinent and eastern China have the largest emission rates, reflecting the densities of domestic animals and the types and intensities of fertilizer use. Anthropogenic emissions have probably increased three-fold since 1950 in line with the increase in global population and food production. Ammonia volatilization from fertilizers is a function of the type of fertilizer, soil conditions, meteorological conditions–temperature, wind speed, precipitation–and fertilizer management. Table 8.6 shows the global use of nitrogenous fertilizers and the corresponding NH3 emissions based on empirical emission factors for different fertilizer types in temperate and tropical conditions (Bouwman

Ammonia

253

Table 8.5

Estimates of global ammonia emissions (Tg N year−1 ) from different sources

Reference Animals Cattle including buffaloes Pigs Horses, mules, asses Sheep, goats Poultry Wild animals Total animals Others Synthetic fertilizers Undisturbed ecosystems Crop plants Biomass burning Human excrement Sea surface Fossil fuel combustion Industry Total emission

Schlesinger and Hartley (1992)

Dentener and Crutzen (1994)

Bouwman et al. (1997)

19.9 2.0 1.8 4.1 2.4 32.3

14.2 2.8 1.2 2.5 1.3 2.5 24.5

14.0 3.4 0.5 1.5 1.9 0.1 21.7

8.5 10 — 5 4 13 2.2 — 75

6.4 5.1 — 2.0 — 7.0 — — 45.0

9.0 2.4 3.6 5.7 2.6 8.2 0.1 0.2 53.6

a

a

Included in undisturbed ecosystems. Source: Bouwman et al. (1997). Reproduced by permission of American Geophysical Union.

Table 8.6 Global use of nitrogenous fertilizers and corresponding NH3 emissions based on empirical emission factors for different fertilizer types and uses Type of nitrogenous fertilizer Urea Ammonium bicarbonate Ammonium nitrate NPK Anhydrous ammonia Nitrogen solutions Calcium ammonium nitrate Ammonium phosphates Ammonium sulfate Others Total

Global use

Emission (Gg N year−1 )

(Tg N year−1 ) 29.2 9.5 8.2 6.6 5.2 4.2 4.1 3.7 2.6 3.7 77.0

Temperate

Tropical

Total

1632 802 25 40 18 11 9 35 34 20 2626

4137 1189 141 219 190 93 72 113 169 85 6409

5769 1991 166 259 209 104 82 147 203 105 9035

Source: Bouwman et al. (1997). Reproduced by permission of American Geophysical Union.

et al., 1997). Seventy per cent of the emission is from developing countries in the tropics; of this 65 % is from urea and 19 % from the volatile hydrolysis product of urea, ammonium bicarbonate, which is widely used in China. Based on data for rice area and yield by country (IRRI, 2002), the approximate relation

254

Trace Gases

between rice yield and N fertilizer use (Figure 7.1), and the emission factors used by Bouwman et al., I estimate a total emission of NH3 from wetland rice of very roughly 3.6 Tg N year−1 , of which 1.2 Tg N year−1 is from China and 0.6 from India. This compares with a total global emission from N fertilizer of 9.0 Tg N year−1 . Clearly wetland rice is an important source of NH3 .

8.3.2 PROCESSES GOVERNING AMMONIA EMISSIONS FROM RICE Urea is the main form of N fertilizer used in rice, together with, in China, ammonium bicarbonate. At least two applications are generally made by broadcasting the fertilizer onto the floodwater: the first 14–21 days after planting the crop and a second at the maximum tillering stage 45–55 days after planting. The first is subject to high rates of loss by volatilization (De Datta and Patrick, 1986; De Datta, 1995). Losses are smaller once the crop canopy and root system are established, because turbulence and hence gas exchange at the water surface are less and absorption by the crop is greater. Rates of volatilization during the early period measured by bulk aerodynamic and micrometeorological methods often account for 30–40 % and sometimes as much as 60 % of the fertilizer applied (Simpson et al., 1984; Cai et al., 1986; Fillery et al., 1986 Freney et al., 1990; De Datta et al., 1989). Losses during the later stages are typically less than half this, depending on how well matched the application is with crop demand. Urea broadcast into the ricefield floodwater is hydrolysed to ammonium, bicarbonate and hydroxyl ions; the reaction is catalysed by the enzyme urease: CO(NH2 )2 + 2H2 O + H+ −−−→ 2 NH4 + + HCO3 − One mol of H+ is consumed in this reaction for every 2 mol of NH4 + formed. In subsequent volatilization of NH3 , 1 mol of H+ is produced for every mol of NH4 + converted to NH3 : NH4 + −−−→ NH3 + H+ Because urease activities are much greater in the soil than in the floodwater, the NH4 + is largely formed in the soil as the urea moves downward by mass flow and diffusion. The NH4 + , H+ and other reactants will also move between the floodwater and soil–both upward and downward–with NH3 being lost from the floodwater by volatilization. The recovery of N in the crop therefore depends on the rate of movement of urea and its reaction products through the soil and on the rate at which the roots remove N from the downward moving pool. Rachhpal-Singh and Kirk (1993a,b) developed a model of these processes based on equations for the transport and reaction of urea, ammoniacal species (NH4 + , NH3 , NH4 OH), carbonate species (H2 CO3 , HCO3 − , CO3 2− ) and mobile acid–base pairs (H2 CO3 –HCO3 − , HCO3 − –CO3 2− , NH4 + –NH3 , NH4 + –NH4 OH, H2 O–OH− ). The equations are of the form of Equation (2.6)

5

4

3

2

1

0

2

60 50 40 30 20 10 0

0

3

2

6

soil)

5

10

dm−3

4

4 6 8 Time (days)

[Urea-N] (mmol

1

%N volatilized

7

0

2

[NH4-N] (mmol

1 dm−3

3 soil)

4

5

6.8

7.0

7.4 Soil pH

7.2

7.6

2 days 5 days 10 days

7.8

Figure 8.9 Profiles of urea-N, ammoniacal-N and pH with depth following broadcast application of urea on ricefield floodwater, and the corresponding rates of NH3 volatilization (calculated with the model of Rachhpal-Singh and Kirk, 1993a,b)

Depth in soil (cm)

0

255

256

Trace Gases

with terms for N, C and base added or removed in urea hydrolysis, organic C and N mineralization, and root uptake. The dynamics of CO2 in the floodwater and the coupled transfer of CO2 and NH3 across the air–water interface (Section 3.5) are allowed for. The model shows that cumulative volatilization of NH3 is sensitive to the initial distribution of urea in the soil, its rate of hydrolysis, and the rate of absorption of N by rice roots. It is largely insensitive to other parameters. For example, it might be thought that addition of organic matter to the soil to acidify the floodwater should lessen NH3 volatilization. However, the model shows that although increased CO2 production affects the diurnal change in floodwater pH, it little affects the daily average pH and hence NH3 volatilization. This is because the relative rates of movement of carbonate species and acidity between the soil and floodwater are such that the increased alkalinization of the floodwater resulting from increased CO2 loss is not matched by an equal inflow of acidity from the soil. The model shows that the spread of urea and NH4 + into the soil is typically only a centimetre or two in a week (Figure 8.9). The recovery of broadcast fertilizer N in the crop must therefore depend entirely on the superficial root system in the soil–floodwater interface. The good recovery of broadcast fertilizer N obtained if the fertilizer is added when the crop demand is maximal (Peng and Cassman, 1998) therefore indicate rapid uptake by roots in the soil–floodwater interface. 8.4 SULFUR COMPOUNDS 8.4.1 GLOBAL BUDGET Submerged soils are important sinks for atmospheric sulfur (Howarth et al., 1992). Sulfate washed into wetlands or deposited from the atmosphere is largely reduced to sulfide by sulfate-reducing bacteria. Subsequent precipitation with metals, especially as FeS, results in more or less permanent removal of the S from the global S cycle. Little sulfur is re-emitted from wetlands into the atmosphere. Table 8.7 gives estimates of global emissions of volatile sulfur compounds from different sources. Total emissions are in the range 98 to 120 Tg (S) year−1 ; 75 % is anthropogenic, mainly from fossil fuel combustion in the northern hemisphere. The main natural sources are the oceans and volcanoes. Wetlands and soils contribute less than 3 % of the total emission. 8.4.2 EMISSIONS FROM RICEFIELDS The main source of S emissions from ricefields is the burning of crop residues, during which most of the sulfur in the residues is converted to volatile oxides (Fox

H2 S

CH3 SCH3

Estimates of global sulfur emissions (Tg S year−1 ) CS2

OCS

SO2

SO4

Totala

b

Numbers in parentheses are fluxes from northern/southern hemispheres. Excluding contributions from sea salt. c Excluding contributions from soil dust. Source: Seinfeld and Pandis (1998). Reproduced by permission of Wiley, New York.

a

Fossil-fuel combustion + industry Total reduced S = 2.2 70 2.2 71–77 (68/6) (mid 1980s) Biomass burning